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micro-organisms

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advances in

geomicrobiology

CAMBRIDG

.Cambridge. org/9 /80521 862221

Micro-organisms and Earth systems -
advances in geomicrobiology

There is growing awareness that important environmental transformations are catalysed,
mediated and influenced by micro-organisms, and such knowledge is having an increasing
influence on disciplines other than microbiology, such as geology and mineralogy. Geo-
microbiology can be defined as the study of the role that microbes have played and are
playing in processes of fundamental importance to geology. As such, it is a truly inter-
disciplinary subject area, necessitating input from physical, chemical and biological
sciences. The book focuses on some important microbial functions in aquatic and terrestrial
environments and their influence on 'global' processes and includes state-of-the-art
approaches to visualization, culture and identification, community interactions and gene
transfer, and diversity studies in relation to key processes. Microbial involvement in key
global biogeochemical cycles is exemplified by aquatic and terrestrial examples. All major
groups of geochemically active microbes are represented, including cyanobacteria, bacteria,
archaea, microalgae and fungi, in a wide range of habitats, reflecting the wealth of diversity
in both the natural and the microbial world. This book represents environmental
microbiology in its broadest sense and will help to promote exciting collaborations between
microbiologists and those in complementary physical and chemical disciplines.

Geoffrey Michael Gadd is Professor of Microbiology and Head of the Division of Environmental
and Applied Biology in the School of Life Sciences at the University of Dundee, UK.

Kirk T. Semple is a Reader in the Department of Environmental Science at Lancaster University,
UK.

Hilary M. Lappin-Scott is Professor of Environmental Microbiology in the School of Biosciences,
University of Exeter, UK.

Symposia of the Society for General Microbiology

Managing Editor: Dr Melanie Scourfield, SGM # Reading, UK
Volumes currently available:

43 Transposition

45 Control of virus diseases

47 Prokaryotic structure and function - a new perspective

51 Viruses and cancer

52 Population genetics of bacteria

53 Fifty years of antimicrobials: past perspectives and future trends

54 Evolution of microbial life

55 Molecular aspects of host-pathogen interactions

56 Microbial responses to light and time

57 Microbial signalling and communication

58 Transport of molecules across microbial membranes

59 Community structure and co-operation in biofilms

60 New challenges to health: the threat of virus infection

61 Signals, switches, regulons and cascades: control of bacterial gene expression

62 Microbial subversion of host cells

63 Microbe-vector interactions in vector-borne diseases

64 Molecular pathogenesis of virus infections

SIXTY-FIFTH SYMPOSIUM OF THE
SOCIETY FOR GENERAL MICROBIOLOGY
HELD AT KEELE UNIVERSITY SEPTEMBER 2005

Edited by

G. M. Gadd, K. T. Semple & H. M. Lappin-Scott

micro-organisms
and earth systems

advances in
geomicrobiology

Published for the Society for General Microbiology

M.J

CAMBRIDGE

UNIVERSITY PRESS

CAMBRIDGE UNIVERSITY PRESS

Cambridge, New York, Melbourne, Madrid, Cape Town, Singapore, Sao Paulo

Cambridge University Press

The Edinburgh Building, Cambridge CB2 2RU, UK

Published in the United States of America by Cambridge University Press, New York

www.camb ridge, org

Information on this title: www.cambridge.org/9780521862226

This publication is in copyright. Subject to statutory exception and to the provision of
relevant collective licensing agreements, no reproduction of any part may take place
without the written permission of Cambridge University Press.

First published in print format 2005

ISBN-13 978-0-511-14130-0 eBook (NetLibrary)
isbn-io 0-511-14130-0 eBook (NetLibrary)

ISBN-13 978-0-521-86222-6 hardback
isbn-io 0-521-86222-1 hardback

Cambridge University Press has no responsibility for the persistence or accuracy of URLs
for external or third-party internet websites referred to in this publication, and does not
guarantee that any content on such websites is, or will remain, accurate or appropriate.

CONTENTS

Contributors vii

Editors' Preface xi

M . Wagner and M . W. Taylor

Isotopic-labelling methods for deciphering the function of uncultured

micro-organisms 1

L. A. Warren

Biofilms and metal geochemistry: the relevance of micro-organism-induced

geochemical transformations 1 1

N. D. Gray and I. M. Head

Minerals, mats, pearls and veils: themes and variations in giant sulfur bacteria 35

D. W. Hopkins, B. Elberling, L. G. Greenfield, E. G. Gregorich, P. Novis,
A. G. O'Donnell and A. D. Sparrow

Soil micro-organisms in Antarctic dry valleys: resource supply and utilization 71

V. R. Phoenix, A. A. Korenevsky, V. R. F. Matias and T. J. Beveridge

New insights into bacterial cell-wall structure and physico-chemistry: implications

for interactions with metal ions and minerals 85

J. Coombs and T. Barkay

Horizontal gene transfer of metal homeostasis genes and its role in microbial
communities of the deep terrestrial subsurface 1 09

L. G. Benning, V. R. Phoenix and B. W. Mountain

Biosilicification: the role of cyanobacteria in silica sinter deposition 131

K. H. Nealson and R. Popa

Metabolic diversity in the microbial world: relevance to exobiology 1 51

D. B. Nedwell

Biogeochemical cycling in polar, temperate and tropical coastal zones:

similarities and differences 173

G. M. Gadd, M. Fomina and E. P. Burford

Fungal roles and function in rock, mineral and soil transformations 201

K. Pedersen

The deep intraterrestrial biosphere 233

SGM symposium 65

vi Contents

J. A. Raven, K. Brown, M. Mackay, J. Beardall, M. Giordano, E. Granum,
R. C. Leegood, K. Kilminster and D. I. Walker

Iron, nitrogen, phosphorus and zinc cycling and consequences for primary

productivity in the oceans 247

J. R. Lloyd

Mechanisms and environmental impact of microbial metal reduction 273

M. Kriiger and T. Treude

New insights into the physiology and regulation of the anaerobic oxidation

of methane 303

N. Clipson, E. Landy and M. Otte

Biogeochemical roles of fungi in marine and estuarine habitats 321

P. C. Bennett and A. S. Engel

Role of micro-organisms in karstification 345

Index 365

SGM symposium 65

CONTRIBUTORS

Barkay, T.

Department of Biochemistry and Microbiology, Cook College, Rutgers University,
76 Lipman Dr., New Brunswick, NJ 08901, USA

Beardall, J.

School of Biological Sciences, Monash University, Clayton, VIC 3800, Australia

Bennett, P. C.

Department of Geological Sciences, The University of Texas at Austin, Austin, TX 787 1 2,
USA

Benning, L. G.

Earth and Biosphere Institute, School of Earth and Environment, University of Leeds, UK

Beveridge, T. J.

Department of Molecular and Cellular Biology, College of Biological Science,
University of Guelph, Guelph, Ontario, Canada N1 G 2W1

Brown, K.

Plant Research Unit, Division of Environmental and Applied Biology, School of Life
Sciences, University of Dundee at SCRI, Scottish Crop Research Institute, Invergowrie,
Dundee DD2 5DA, Scotland, UK

Burford, E. P.

Division of Environmental and Applied Biology, Biological Sciences Institute, School of Life
Sciences, University of Dundee, Dundee DD1 4HN, Scotland, UK

Clipson, N.

Department of Industrial Microbiology, University College Dublin, Belfield, Dublin 4, Ireland

Coombs, J.

Department of Biochemistry and Microbiology, Cook College, Rutgers University,
76 Lipman Dr., New Brunswick, NJ 08901, USA

Elberling, B.

Institute of Geography, University of Copenhagen, 0ster Voldgade 10, DK-1350,
Copenhagen K., Denmark

Engel, A. S.

Department of Geology and Geophysics, Louisiana State University, Baton Rouge,
LA 70803, USA

Fomina, M.

Division of Environmental and Applied Biology, Biological Sciences Institute, School of Life
Sciences, University of Dundee, Dundee DD1 4HN, Scotland, UK

Gadd,G. M.

Division of Environmental and Applied Biology, Biological Sciences Institute, School of Life
Sciences, University of Dundee, Dundee DD1 4HN, Scotland, UK

SGM symposium 65

viii Contributors

Giordano, M.

Department of Marine Science, Universita Politecnica delle Marche, 60131 Ancona,
Italy

Granum, E.

Department of Animal and Plant Sciences, University of Sheffield, Sheffield S1 2TN, UK

Gray, N. D.

School of Civil Engineering and Geosciences, Institute for Research on the Environment
and Sustainability and Centre for Molecular Ecology, University of Newcastle, Newcastle
uponTyneNEI 7RU, UK

Greenfield, L. G.

School of Biological Sciences, University of Canterbury, Private Bag 4800, Christchurch,
New Zealand

Gregorich, E. G.

Agriculture Canada, Central Experimental Farm, Ottawa, Canada K1 A 0C6

Head, I. M.

School of Civil Engineering and Geosciences, Institute for Research on the Environment
and Sustainability and Centre for Molecular Ecology, University of Newcastle, Newcastle
uponTyneNEI 7RU, UK

Hopkins, D.W.

School of Biological and Environmental Sciences, University of Stirling, Stirling FK9 4LA,
Scotland, UK

Kilminster, K.

School of Plant Biology, University of Western Australia, M090 35 Stirling Highway, Crawley,
WA 6009, Australia

Korenevsky, A. A.

Department of Molecular and Cellular Biology, College of Biological Science,
University of Guelph, Guelph, Ontario, Canada N1G 2W1

Kruger, M.

Federal Institute for Geosciences and Resources (BGR), Stilleweg 2, D-30655 Hannover,
Germany, and Max-Planck-lnstitute for Marine Microbiology, Celsiusstrasse 1, D-28359
Bremen, Germany

Landy, E.

School of Biomedical and Molecular Sciences, University of Surrey, Guildford GU2 7XH,
UK

Leegood, R. C.

Department of Animal and Plant Sciences, University of Sheffield, Sheffield S1 2TN, UK

Lloyd, J. R.

The Williamson Research Centre for Molecular Environmental Studies and the School
of Earth, Atmospheric and Environmental Sciences, University of Manchester, Manchester
M13 9PLUK

SGM symposium 65

Contributors ix

Mackay, M.

Plant Research Unit, Division of Environmental and Applied Biology, School of Life
Sciences, University of Dundee at SCRI, Scottish Crop Research Institute, Invergowrie,
Dundee DD2 5DA, Scotland, UK

Matias,V. R. F.

Department of Molecular and Cellular Biology, College of Biological Science,
University of Guelph, Guelph, Ontario, Canada N1 G 2W1

Mountain, B. W.

Institute of Geological and Nuclear Sciences, Wairakei Research Centre, Taupo, New Zealand

Nealson, K.H.

Department of Earth Sciences, University of Southern California, Los Angeles,
CA 90089-0740, USA

Nedwell, D. B.

Department of Biological Sciences, University of Essex, Colchester C04 3SQ, UK

Novis, P.

Manaaki Whenua - Landcare Research, PO Box 69, Lincoln 81 52, New Zealand

O'Donnell, A. G.

Institute for Research on Environment and Sustainability, University of Newcastle upon Tyne,
Newcastle upon Tyne NE1 7RU, UK

Otte, M.

Department of Botany, University College Dublin, Belfield, Dublin 4, Ireland

Pedersen, K.

Deep Biosphere Laboratory, Department of Cell & Molecular Biology, Goteborg University,
Box 462, SE-405 30 Goteborg, Sweden

Phoenix, V. R.

Department of Molecular and Cellular Biology, College of Biological Science,
University of Guelph, Guelph, Ontario, Canada N1 G 2W1

Popa, R.

Department of Earth Sciences, University of Southern California, Los Angeles,
CA 90089-0740, USA

Raven, J. A.

Plant Research Unit, Division of Environmental and Applied Biology, School of Life
Sciences, University of Dundee at SCRI, Scottish Crop Research Institute, Invergowrie,
Dundee DD2 5DA, Scotland, UK

Sparrow, A. D.

School of Biological Sciences, University of Canterbury, Private Bag 4800, Christchurch,
New Zealand, and Department of Natural Resources and Environmental Sciences, University
of Nevada, 1000 Valley Rd, Reno, NV 89512, USA

SGM symposium 65

Contributors

Taylor, M.W.

Department of Microbial Ecology, University of Vienna, Althanstr. 14, A-1 090 Vienna,
Austria

Treude, T.

Max-Planck-Institute for Marine Microbiology, Celsiusstrasse 1, D-28359 Bremen, Germany

Wagner, M.

Department of Microbial Ecology, University of Vienna, Althanstr. 14, A-1 090 Vienna,

Austria

Walker, D. I.

School of Plant Biology, University of Western Australia, M090 35 Stirling Highway, Crawley,
WA 6009, Australia

Warren, LA.

School of Geography and Earth Sciences, McMaster University, 1280 Main St. West,
Hamilton, Ontario, Canada L8S4K1

SGM symposium 65

EDITORS' PREFACE

The science of the environment encompasses a huge number of biological, chemical
and physical disciplines. For several years, scientists have been interested in large-scale
environmental processes/phenomena, such as soil formation, global warming and
global elemental cycling. Until recently, the role and impact of micro-organisms on
these 'global' environmental processes has been largely ignored or, at best, underesti-
mated. However, there is growing awareness that important environmental trans-
formations are catalysed, mediated and influenced by micro-organisms, and such
knowledge is having an increasing influence on disciplines other than microbiology,
such as geology and mineralogy. Geomicrobiology can be defined as the study of the
role that microbes have played and are playing in processes of fundamental importance
to geology As such, it is a truly interdisciplinary subject area, necessitating input from
physical, chemical and biological sciences, in particular combining the fields of
environmental and molecular microbiology together with significant areas of
mineralogy, geochemistry and hydrology. As a result, geomicrobiology is probably the
most rapidly growing area of microbiology at present. It is timely that this topic should
be the subject of a Plenary Symposium volume of the Society for General Microbiology
(SGM) to emphasize and define this important area of microbiological interest, and
help to promote exciting collaborations between microbiologists and other environ-
mental and Earth scientists.

This Symposium arose from the Environmental Microbiology Group of the SGM and
presents a snapshot of some key areas of geomicrobiology written by experts in their
respective fields. The book focuses on some important microbial functions in aquatic
and terrestrial environments and their influence on 'global' processes and includes
state-of-the-art approaches to visualization, culture and identification, community
interactions and gene transfer, as well as diversity studies in relation to key pathways.
Novel approaches for the study of diversity and function of microbial communities are
highlighted and applied to key environmental problems, such as community
interactions in biofilms, and the microbiology of surface, sub-surface and extreme
environments. Microbial involvement in key global biogeochemical cycles is exem-
plified by aquatic and terrestrial examples, and includes metal and mineral trans-
formations and development, element cycling in marine and estuarine systems, primary
production, and anaerobic methane oxidation. All major groups of geochemically
active microbes are represented, including cyano bacteria, bacteria, archaea, microalgae
and fungi, in a wide range of habitats both aquatic and terrestrial, aerobic and
anaerobic, and benign and extreme, reflecting the wealth of diversity in both the
natural and the microbial world. It should also be appreciated that many of the natural
processes discussed also have application in applied contexts such as agriculture and

SGM symposium 65

xii Editors' preface

plant productivity, environmental exploitation and resource utilization and microbial
treatment of pollution. This book will truly represent environmental microbiology in
its broadest sense and we hope that it will have broad appeal, not only to environmental
microbiologists, but also to environmental scientists, geologists, geochemists, Earth
scientists, ecologists and environmental biotechnologists. A beneficial outcome may
be the promotion of exciting collaborations between microbiologists and those in
complementary physical and chemical disciplines. Only through interdisciplinary
scientific approaches will the roles of micro-organisms in Earth systems be better
clarified, appreciated and understood.

We would like to thank all the authors for enthusiastically supporting this project
despite their obvious heavy work commitments. We also wish to thank Diane Purves
(University of Dundee) for expert assistance with manuscripts and author liaison, and
all at the SGM office, particularly Melanie Scourfield, Josiane Dunn and Janet Hurst,
for their efficient help and support with the Symposium organization and production of
this volume.

Geoffrey M. Gadd

Kirk T. Semple

Hilary M. Lappin-Scott

SGM symposium 65

Isotopic-labelling methods for
deciphering the function of
uncultured micro-organisms

Michael Wagner and Michael W. Taylor

Department of Microbial Ecology, University of Vienna, Althanstr. 1 4, A-1 090 Vienna, Austria

INTRODUCTION

With the benefit of hindsight, the last 20 years in microbial ecology will probably be
referred to as the census period that dramatically changed our perception of bio-
diversity within the three domains of life. Bacteria and archaea are no longer viewed as
groups of peculiar and morphologically simple organisms that show relatively little
diversification despite their long evolutionary history, but have now been recognized
to harbour a perplexing number of novel phylogenetic lineages (Rappe & Giovannoni,
2003). Current estimates assume that the number of prokaryotic species ranges in the
millions and thus vastly exceeds the fewer than 10000 described prokaryotic species
that have been isolated to date in pure culture (Curtis et ai, 2002). This dramatic para-
digm shift was only made possible by the development of cultivation-independent
molecular approaches for surveying microbial diversity in nature. Whilst it is now
evident that most prokaryotes cannot be cultured easily, due to their living in complex
communities and their intimate metabolic links with both their abiotic and biotic
environments, the powerful arsenal of techniques at our disposal enables us to see
beyond the 'cultured few' and gain valuable insights into the realm of uncultured micro-
organisms (Wagner, 2004). It is now relatively straightforward to determine the species
richness of natural microbial communities by comparative sequence analysis of
environmentally retrieved 16S rRNA gene sequences (Olsen et al., 1986; Schloss &
Handelsman, 2004). Furthermore, even the abundance of bacteria in environmental
samples can be determined easily by using quantitative PCR (Skovhus et al., 2004) or
hybridization formats such as fluorescence in situ hybridization (FISH) (Wagner et al.,
2003). Due to these technological advances, many novel and often numerically

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

M. Wagner and M. W. Taylor

important bacterial and archaeal species have been identified in various ecosystems.
However, for most of these players, their ecophysiology and contribution to the func-
tioning of the ecosystem remain hidden.

Given the above, one of the biggest challenges in contemporary microbial ecology is
to develop strategies that enable us (i) to directly investigate metabolic properties of
defined but uncultured micro-organisms, and (ii) to identify those uncultured organ-
isms that are responsible for defined processes within their natural environment. These
two criteria highlight important differences in approaches that are currently taken in
the study of microbial community function. In a microbial world dominated by un-
cultured representatives, those investigators who approach matters from a diversity
standpoint typically target organisms of interest (e.g. those that are highly abundant in
a particular ecosystem) and endeavour to find out what metabolic traits these species
possess. An alternative, equally valid approach is to recognize a process occurring in the
environment (e.g. sulfate reduction or denitrification) and try to establish which
organisms are responsible. Both approaches are served equally well by the methods
outlined in this chapter. Specifically, we provide an overview of a group of recently
developed methods that exploit the addition of isotope-labelled substrates to complex
microbial communities in order to bridge the gap between microbial community
structure and function.

IDENTIFICATION OF PLAYERS THAT INCORPORATE LABELLED
SUBSTRATES

One of the most elegant ways to better understand the ecophysiology of uncultured
bacteria is to expose them to isotope-labelled substrates and, subsequently, to link their
identification with a measurement of substrate incorporation into their biomass. This
concept was first realized by extracting phospholipid fatty acids from complex
microbial communities after incubation with 13 C-labelled substrates, and identifying
active groups of micro-organisms by detection of labelled signature compounds using
isotope-ratio mass spectrometry (Boschker et aL, 1998). However, this approach is not
suited for the identification of uncultured bacteria, because their phospholipid fatty-
acid composition is unknown. Consequently, several methods, which are presented in
more detail below, were developed to exploit other, more informative biomarkers or
even to allow simultaneous organism identification and detection of substrate
incorporation at the single-cell level. A common feature of all these methods is that it is
possible to fine-tune the experimental setup such that very specific questions regarding
the ecophysiology of selected micro-organisms can be answered. This is achieved by
performing parallel experiments using modified incubation conditions (pH, tempera-
ture, presence/absence of different electron acceptors, presence/absence of specific
inhibitors for selected microbial groups etc.) under which the microbial biomass

SGM symposium 65

Function of uncultured micro-organisms

encounters the labelled substrate. Alternatively, this modus operandi enables microbial
ecologists to hunt for novel micro-organisms responsible for defined functions in the
environment. However, with the exception of the so-called HetC0 2 -microautoradio-
graphy (MAR) approach (Hesselsoe etaL, 2005), all of these methods are dependent on
the availability of commercial isotope labelling for the compounds of interest. Further-
more, all mentioned techniques fail to detect micro-organisms that convert labelled
compounds without using them for the synthesis of biomass or storage compounds.

DNA stable-isotope probing (DNA-SIP) is based on the incorporation of stable iso-
tope-labelled substrates into the DNA of substrate-consuming micro-organisms
(reviewed recently by Dumont 6c Murrell, 2005). To date, DNA-SIP has only been
performed with 13 C-labelled substrates. In DNA-SIP, the heavy DNA of the substrate
consumers is separated from the light DNA of the other community members by
buoyant density-gradient centrifugation with added ethidium bromide. The identity of
the active bacteria is revealed subsequently by amplification, cloning and comparative
sequence analysis of 16S rRNA genes or selected functional genes from the heavy DNA.
Probing the light- and heavy-DNA fractions with specific PCR primers can be applied
to show rapidly whether a given bacterial player has incorporated the offered substrate.
One of the major advantages of DNA-SIP is that large genomic-DNA fragments of the
active community fraction can be isolated and subsequently analysed by the environ-
mental-genomics approach (DeLong, 2002, 2005). On the other hand, the DNA of
active community members will only be labelled sufficiently for DNA-SIP if the added
compounds are highly enriched in 13 C, and if relatively high concentrations of labelled
substrate are offered over extended time periods (up to several weeks) so that slow-
growing micro-organisms are also able to replicate their DNA and to divide. It is
essential to be aware that such incubation conditions might induce major shifts in the
composition of the resident microbial community, perhaps comparable to the biases
reported previously for enrichment cultures (Wagner et ai, 1993). Compared to the
rapid-growing r-strategists, K-strategists that are adapted to low substrate concen-
trations might thus be heavily underrepresented or even missing (if the applied
substrate concentration is inhibitory to them) in the heavy-DNA fraction. Further-
more, due to the long incubation times required for DNA-SIP, primary consumers will
transfer label to other community members by cross-feeding. Therefore, not all
genomes found in the heavy DNA are from bacteria capable of using the labelled
compound (although, on a more positive note, this effect could also be exploited to
obtain indications of metabolic interactions among different community members).

Several of the limitations of DNA-SIP can be overcome by RNA-SIP (Manefield et al.,
2002, 2005). In microbial cells, synthesis rates of RNA (including 16S rRNA) are
significantly higher than rates of DNA replication; consequently, if DNA is replaced by

SGM symposium 65

M. Wagner and M. W. Taylor

• • • • m .
mm*9 f • \m .

JO • #'S • •<••*••

•

•■

(a)

i •

4| ft * *

*"#

* + *#*

1 * #*-■

•

.-•

* -H

(b) #

Fig. 1. Isotope-array experiment to identify butyrate-oxidizing bacteria under aerobic conditions in
activated sludge from a nutrient-removal wastewater-treatment plant, (a) Fluorescence scan of the
array; (b) radioactivity scan of the same array. Pictures kindly provided by A. Loy, M. Schloter and M.
Hesselsoe.

RNA as a biomarker, much shorter incubation times are required to obtain sufficient
amounts of labelled material. 13 C-labelled RNA can be separated from light RNA
via caesium trifluoroacetate (CsTFA) density-gradient centrifugation, although it has
been observed that labelled RNA with a specific buoyant density is found over a number
of fractions in CsTFA gradients. To solve this problem, 16S rRNA from each density
fraction must be amplified separately by RT-PCR and subsequently analysed by a
fingerprinting method such as denaturing-gradient gel electrophoresis (DGGE).
Increasing band intensities in the fingerprint pattern of certain 16S rDNA fragments
from high-density fractions throughout the duration of the labelled-substrate pulse are
then used to demonstrate that the 16S rRNA of a particular bacterium has become
increasingly labelled during the duration of the experiment. Subsequently, the
respective DGGE band is cut from the gel, reamplified by PCR and sequenced in order
to identify the substrate-consuming micro-organisms. The disadvantage of this proce-
dure is that it is very time-consuming and dependent on the resolving power of DGGE,
which is insufficient in highly complex microbial communities. In future studies, the

SGM symposium 65

Function of uncultured micro-organisms

identification of bacteria with a specific function could be accelerated and improved in
resolution by hybridizing heavy DNA or RNA obtained by SIP to diagnostic rRNA-
targeted microarrays (Guschin et al., 1997; Loy et al., 2002, 2005; Wilson et al., 2002).

RNA as a biomarker is also exploited by the so-called isotope-array approach
(Adamczyk et al., 2003) but, in contrast to SIP, the microbial communities are incubated
with radioactively (and not lj C) -labelled compounds prior to nucleic-acid extraction.
To date, exclusively 14 C-labelled compounds have been used for isotope-array analysis.
Again, as RNA is targeted, only relatively short incubation times are required. After
community RNA extraction, the RNA is labelled covalently with a fluorescent dye and
hybridized to a suitable rRNA-targeted microarray by using standard procedures.
Visualization of the microarray by using a fluorescence scanner reveals the community
structure, whilst recording the radioactivity of each microarray spot by use of a beta-
imager shows which community member has incorporated radioactive label into its
rRNA and thus consumed the offered substrate (Fig. 1). Due to limited resolution of
commercially available beta-imagers, larger probe-spot sizes are required for optimal
results compared with standard microarrays. An advantage of the isotope array
compared with RNA-SIP is that it is PCR-independent and thus offers the potential for
quantitative analysis. The ratio of radioactive signal to fluorescence signal, which can
be determined easily for each hybridized probe spot, reflects the incorporation rate of
the labelled substrate into the RNA of the detected bacterial population (s) and can thus
be used to compare their metabolic activities within a complex microbial community.
On the other hand, and in contrast to SIP, the nature of this format makes it impossible
to identify novel substrate-consuming micro-organisms whose 16S rRNA sequences
are not yet known (i.e. you can only find what you are looking for) . In the future, the iso-
tope array should be adapted to the use of stable isotopes so as to avoid radioactive lab
work and also render it adaptable to the analysis of microbial communities colonizing
humans. For this purpose, beta-imaging must be replaced by detection methods such as
time-of-flight secondary-ion mass spectrometry.

In contrast to the aforementioned methods, structure and function of microbial
communities can be analysed by a combination of FISH and MAR at the single-cell
level (Lee et al., 1999; Ouverney & Fuhrman, 1999; Daims et al., 2001; Nielsen et al.,
2002, 2003b) (Fig. 2), as long as the environmental sample is suited for FISH analysis.
After incubation of the microbial communities with radioactively labelled substrate,
community members are identified by FISH with rRNA-targeted oligonucleotide
probes. In parallel, the incorporation of radioactive substrate into the biomass of
bacterial cells is visualized with a radiation-sensitive silver halide emulsion, which is
placed on top of the radiolabeled organisms and subsequently processed by standard
photographic procedures. Excited silver ions will precipitate as metallic silver and

SGM symposium 65

M. Wagner and M. W. Taylor

2-10 pm

cryo-
section

Cover slip

Silver-grain formation above cells that
took up radioactively labelled substrate

CLSM

Bacteria labelled by FISH

Fig. 2. Schematic representation of the FISH-MAR approach. CLSM, Confocal laser-scanning
microscope.

Methanol added

Control

Fig. 3. Application of FISH-MAR to investigate the ecophysiology of uncultured filamentous micro-
organisms. The filaments showed uptake of methanol in MAR experiments (a) and were identified
simultaneously by FISH as members of the ' Gammaproteobacteria' (b). Control experiments with
pasteurized biomass showed no unspecific uptake or binding of labelled methanol (c, d). Pictures
kindly provided by K. Stoecker and H. Daims.

SGM symposium 65

Function of uncultured micro-organisms

appear as black grains after development of the film (Fig. 3). As FISH requires perm-
eabilization of the bacterial cell envelope, radioactive substrates that are taken up by a
bacterial cell, but not incorporated into macromolecules, will not result in silver-grain
formation. It should, however, be noted that MAR cannot differentiate between con-
version of labelled organic substrates into intracellular storage products and active
substrate metabolism. The recently developed HetC0 2 -MAR approach (Hesselsoe
et al., 2005) allows this limitation to be overcome by utilizing the tendency of meta-
bolically active heterotrophic bacteria to assimilate significant amounts of C0 2 . Thus,
14 C0 2 is added as an activity marker to complex microbial communities and, after
suitable incubation periods, active autotrophic and heterotrophic community members
can be identified by FISH-MAR. HetCO^-MAR also avoids difficulties associated with
obtaining many isotope-labelled substrates, as this approach requires that only the C0 2
is labelled and other added substrates are not.

Major advantages of FISH-MAR are that incubation times are comparatively short,
because every labelled macromolecule of the monitored bacterial cells contributes to
the detection of substrate incorporation. Furthermore, the amount of substrate incor-
poration into specific uncultured bacterial cells can be quantified by counting the
number of silver grains on top of the cells if bacteria with known specific radioactivity
are used as an internal standard. This technique can be applied to measure in situ
substrate affinities {K s ) of uncultured bacteria and to study physiological differences
within naturally occurring single populations of bacteria (Nielsen et al., 2003a). In
future studies, it should be possible to modify the FISH-MAR method such that radio-
actively labelled substrates are replaced by stable isotope-labelled compounds and
the stable-isotope composition of FISH-identified microbial cells is measured by
secondary-ion mass spectrometry (Orphan et al., 2001, 2002).

The last few years have seen the development of an impressive battery of tools that
measure function of uncultured micro-organisms by the addition of labelled substrates.
After initial, proof-of-principle type studies, several of these tools are now applied
(almost) routinely in microbial ecology laboratories and they will be used increasingly
in time-course experiments to uncover metabolic networks in multi-component
microbial communities (Lueders et al., 2004). Together with quantitative FISH-based
analysis of co-occurring microbial populations (Daims et al., 2005), these approaches
will shed new light on the complex network of interactions existing in microbial
communities.

OUTLOOK

We have provided here a brief overview of cutting-edge methods for the analysis of
microbial-community function. As a field that necessarily emphasizes the importance

SGM symposium 65

8 M. Wagner and M. W. Taylor

of technological advances, microbial-community ecology should continue to benefit
significantly from the development and application of isotopic-labelling methods.
We predict that this area should prove a productive meeting ground for process- and
diversity-focused microbial ecologists.

The approaches outlined above should prove especially fruitful when used in combina-
tion with another emerging subdiscipline of microbial ecology, that of environmental
genomics. Much has been made recently of the power of environmental genomics,
exemplified by the massive sequencing efforts undertaken in the Sargasso Sea (Venter
et aL, 2004). However, such studies tell us more about potential, rather than actual,
community function because, for many of the retrieved genes, annotation provides no
indication of a possible function of the encoded protein. Even for those genes where
annotation is meaningful, no information on actual expression in the environment is
apparent. Environmental proteomics offer a partial solution to this quandary (Wilmes
& Bond, 2004), but this field is in its infancy and widespread application is surely
several years away. The real beauty of environmental genomics, then (at least from a
community-function perspective), may be its role in the formulation of hypotheses
(Wagner, 2005). Strong hints about organismal metabolism can be obtained via meta-
genome sequencing, hints that can be evaluated by using the isotopic-labelling
techniques described in this chapter. The synergy derived from mixing these approaches
could be profound: imagine a situation where one discovers a 'new' organism of
presumed importance in an environment (e.g. any of a dozen or more candidate phyla
with currently no cultured representatives), but nothing is known of its metabolism.
FISH-MAR, SIP and isotope arrays all offer potential avenues for investigation, yet
where would one start? Possible carbon sources alone number in the hundreds and it
is simply not feasible to test all of these with isotopic methods. However, if one has
genomic information for such organisms (as obtained by bacterial artificial chromo-
some or shotgun sequencing from a mixed sample), then valuable clues as to suitable
target substrates should be forthcoming. The need for environmental genomics is
arguably reduced when one approaches community function from the other side, i.e.
starting with a process and identifying the key players. Here, the critical step is in
defining incubation conditions where only the organisms capable of the function (s) of
interest are active.

Since its first coupling with molecular-phylogenetic techniques in the late 1990s,
isotopic labelling has provided important insights into the function of specific, un-
cultured micro-organisms. Continuing technological advances and further integration
with environmental-genomics and -proteomics approaches can only enhance the value
of this area of microbial ecology.

SGM symposium 65

Function of uncultured micro-organisms

REFERENCES

Adamczyk, J., Hesselsoe, M., Iversen, N., Horn, M., Lehner, A., Nielsen, P. H., Schloter,
M., Roslev, P. & Wagner, M. (2003). The isotope array, a new tool that employs
substrate-mediated labeling of rRNA for determination of microbial community
structure and function. Appl Environ Microbiol 69, 6875-6887.

Boschker, H. T. S., Nold, S. C, Wellsbury, P., Bos, D., de Graaf, W. # Pel, R., Parkes, R. J. &
Cappenberg, T. E. (1998). Direct linking of microbial populations to specific
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Curtis, T. P., Sloan, W. T. & Scannell, J. W. (2002). Estimating prokaryotic diversity and
its limits. Proc Natl Acad Sci USA 99, 1 0494-1 0499.

Daims, H., Nielsen, J. L, Nielsen, P. H., Schleifer, K.-H. & Wagner, M. (2001). In situ
characterization of Nitrospira-Wke nitrite-oxidizing bacteria active in wastewater
treatment plants. Appl Environ Microbiol 67, 5273-5284.

Daims, H., Lucker, S. & Wagner, M. (2005). daime, a novel image analysis program for
microbial ecology and biofilm research. Environ Microbiol (in press).

DeLong, E. F. (2002). Microbial population genomics and ecology. Curr Opin Microbiol
5, 520-524.

DeLong, E. F. (2005). Microbial community genomics in the ocean. Nat Rev Microbiol
3, 459-469.

Dumont, M. G. & Murrell, J. C. (2005). Stable isotope probing - linking microbial identity
to function. Nat Rev Microbiol 3, 499-504.

Guschin, D. Y, Mobarry, B. K., Proudnikov, D., Stahl, D. A., Rittmann, B. E. & Mirza-
bekov, A. D. (1997). Oligonucleotide microchips as genosensors for determinative
and environmental studies in microbiology. Appl Environ Microbiol 63, 2397-2402.

Hesselsoe, M., Nielsen, J. L, Roslev, P. & Nielsen, P. H. (2005). Isotope labeling and
microautoradiography of active heterotrophic bacteria on the basis of assimilation of
1 4 C 2 . Appl Environ Microbiol 7 1 , 646-6 5 5 .

Lee, N., Nielsen, P. H., Andreasen, K. H., Juretschko, S., Nielsen, J. L, Schleifer, K.-H. &
Wagner, M. (1999). Combination of fluorescent in situ hybridization and micro-
autoradiography -a new tool for structure-function analyses in microbial ecology.
Appl Environ Microbiol 65, 1289-1297.

Loy, A., Lehner, A., Lee, N., Adamczyk, J., Meier, H., Ernst, J., Schleifer, K.-H. &
Wagner, M. (2002). Oligonucleotide microarray for 16S rRNA gene-based detection
of all recognized lineages of sulfate-reducing prokaryotes in the environment. Appl
Environ Microbiol 68, 5064-5081 .

Loy, A., Schulz, C, Lucker, S., Schopfer-Wendels, A., Stoecker, K., Baranyi, C, Lehner,
A. & Wagner, M. (2005). 16S rRNA gene-based oligonucleotide microarray for
environmental monitoring of the betaproteobacterial order " Rhodocydales" . Appl
Environ Microbiol 71,1 373-1 386.

Lueders, T., Wagner, B., Claus, P. & Friedrich, M. W. (2004). Stable isotope probing of
rRNA and DNA reveals a dynamic methylotroph community and trophic interactions
with fungi and protozoa in oxic rice field soil. Environ Microbiol 6, 60-72.

Manefield, M., Whiteley, A. S., Griffiths, R. I. & Bailey, M. J. (2002). RNA stable isotope
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Environ Microbiol 68, 5367-5373.

Manefield, M., Griffiths, R. I., Leigh, M. B., Fisher, R. & Whiteley, A. S. (2005).
Functional and compositional comparison of two activated sludge communities
remediating coking effluent. Environ Microbiol 7, 71 5-722.

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10 M. Wagner and M. W. Taylor

Nielsen, J. L, Juretschko, S., Wagner, M. & Nielsen, P. H. (2002). Abundance and

phylogenetic affiliation of iron reducers in activated sludge as assessed by fluores-
cence in situ hybridization and microautoradiography. Appl Environ Microbiol 68,

4629-4636.
Nielsen, J. L, Wagner, M. & Nielsen, P. H. (2003a). Use of microautoradiography to study

in situ physiology of bacteria in biofilms. Rev Environ Sci Biotechnol 2, 261-268.
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of cell-specific substrate uptake by probe-defined bacteria under in situ conditions by

microautoradiography and fluorescence in situ hybridization. Environ Microbiol 5,

202-211.
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ecology and evolution: a ribosomal RNA approach. Annu Rev Microbiol 40, 337-

365.
Orphan, V. J., House, C. H., Hinrichs, K.-U., McKeegan, K. D. & DeLong, E. F. (2001).

Methane-consuming archaea revealed by directly coupled isotopic and phylogenetic

analysis. Science 293, 484^487.
Orphan, V. J., House, C. H., Hinrichs, K. U., McKeegan, K. D. & DeLong, E. F. (2002).

Multiple archaeal groups mediate methane oxidation in anoxic cold seep sediments.

Proc Natl Acad Sci U 5 A 99, 7663-7668.
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probe technique for the determination of radioisotope uptake by specific microbial

cell types in situ. Appl Environ Microbiol 65, 1 746-1 752.
Rappe, M. S. & Giovannoni, S. J. (2003). The uncultured microbial majority. Annu Rev

Microbiol 57, 369-394.
Schloss, P. D. & Handelsman, J. (2004). Status of the microbial census. Microbiol Mol Biol

Rev 68, 686-691.
Skovhus, T. L, Ramsing, N. B., Holmstrom, C, Kjelleberg, S. & Dahllof, I. (2004). Real-
time quantitative PCR for assessment of abundance of Pseudoalteromonas species in

marine samples. Appl Environ Microbiol 70, 2373-2382.
Venter, J. C, Remington, K., Heidelberg, J. F. & 20 other authors (2004). Environmental

genome shotgun sequencing of the Sargasso Sea. Science 304, 66-74.
Wagner, M. (2004). Deciphering the function of uncultured microorganisms. ASM News

70, 63-70.
Wagner, M. (2005). The community level: physiology and interactions of prokaryotes in the

wilderness. Environ Microbiol 7, 483-485.
Wagner, M., Amann, R., Lemmer, H. & Schleifer, K.-H. (1993). Probing activated sludge

with oligonucleotides specific for proteobacteria: inadequacy of culture-dependent

methods for describing microbial community structure. Appl Environ Microbiol 59,

1520-1525.
Wagner, M., Horn, M. & Daims, H. (2003). Fluorescence in situ hybridisation for the

identification and characterisation of prokaryotes. CurrOpin Microbiol 6, 302-309.
Wilmes, P. & Bond, P. L. (2004). The application of two-dimensional polyacrylamide gel

electrophoresis and downstream analyses to a mixed community of prokaryotic

microorganisms. Environ Microbiol 6, 91 1-920.
Wilson, K. H., Wilson, W. J., Radosevich, J. L., DeSantis, T. Z., Viswanathan, V. S.,

Kuczmarski, T. A. & Andersen, G. L. (2002). High-density microarray of small-

subunit ribosomal DNA probes. Appl Environ Microbiol 68, 2535-2541 .

SGM symposium 65

Biof ilms and metal geochemistry;
the relevance of micro-organism-
induced geochemical
transformations

Lesley A. Warren

School of Geography and Earth Sciences, McMaster University, 1 280 Main St. West, Hamilton,
Ontario, Canada L8S4K1

INTRODUCTION

This chapter is intended to provide a brief overview of the key concepts underlying the
emerging area of environmental microbial metal geochemistry, rather than an exhaus-
tive synthesis. The reader is referred to the following more comprehensive reviews
on the biogeochemistry of metals (Warren & Haack, 2001), metal-mineral reactions
(Brown & Parks, 2001), emerging molecular-level geochemical techniques (O'Day,
1999; Brown & Sturchio, 2002) and a recent synthesis of how genetic expression in the
environment can underpin geochemical reactions (Croal et al., 2004). The relevance of
micro-organisms to metal behaviour arises from the overlap of the biosphere with the
geosphere and the transformations that occur because of their interactions. Micro-
organisms have evolved in intimate association with the rocks, soils and waters (i.e.
geosphere) in which they find themselves. In order to grow and survive, they have adap-
ted to these environments and use the inorganic components to drive their metabolic
machinery; the myriad functional pathways by which they do so ensure that they
influence a number of key elemental cycles in the process. As a consequence, many
important geochemical processes are ultimately shaped by life, rather than strict
geochemical equilibria, a fact that is increasingly recognized as strict geochemical
principles fail to constrain observed environmental behaviour.

Trace-metal behaviour in the environment is of increasing global concern as water and
soil contamination with these toxic substances continues and the detrimental effects on
ecosystems and human health emerge. While investigation into metal behaviour has
spanned many disciplines reflecting discipline-specific foci and approaches to the topic

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

12 L.A.Warren

(e.g. biology, chemistry, geochemistry, toxicity, physics), what coalesces from this broad,
substantial literature is that the controls on aqueous metal behaviour supersede
individual discipline boundaries; rather, it is often dynamic, complex, biological (prin-
cipally micro-organism driven) and geochemical linkages that drive geochemistry
(Reysenbach & Shock, 2002; Newman & Banfield, 2002; Warren & Kauffman, 2003;
Warren, 2004).

Probing microbial-mineral-metal interactions has provided substantive evidence of
microbial shaping of metal fate (e.g. Ehrlich, 2002; Holden & Adams, 2003 ; Islam et al.,
2004), and has clearly alluded to the need to understand how microbial activity and
geochemistry interact to ultimately determine metal impact. It is increasingly evident
that microbial activity can substantially impact aqueous geochemical behaviour in
ways not predicted by classic geochemical models, a serious issue for contaminated
environments where bacteria flourish such as acid mine drainage (AMD) or contamin-
ated subsurface environments where migration of contaminants to groundwater
supplies represents a serious health hazard.

AQUEOUS METAL GEOCHEMISTRY: FUNDAMENTALS

At the most fundamental level, the single most important predictor of metal behaviour,
i.e. mobility within the geosphere, bioavailability, bioaccumulation and toxicity, is the
solution or dissolved (operationally defined typically as less than 0-45 Jim or, more
rigorously, <0-l ^m) concentration of the element (e.g. Martell et al., 1988; Campbell
& Tessier, 1989; Hare, 1992; Campbell, 1995; Unz & Shuttleworth, 1996; Warren et al.,
1998; Hassler et al., 2004). Typically, the greater the solution concentration of a metal,
the more likely are the negative impacts associated with that contaminant, reflecting
greater mobility, bioavailability and toxicity. Thus, trace-element geochemistry has
focused on understanding the processes by which metals partition between the solution
and solid compartments within environmental systems and the controls on the pro-
cesses involved (Warren & Haack, 2001; and references therein). While precipitation
of metals within metal-bearing minerals such as sulfides, carbonates or hydroxides
can play a role, principally in sedimentary systems where concentrations of metals can
reach saturation with respect to a given mineral phase, it is now well established that
metal solid-solution partitioning is often dominantly controlled by interfacial reac-
tions occurring between charged functional groups at solid surfaces and solution metal
species, collectively referred to as sorption reactions (Jenne, 1968; Dzombak & Morel,
1990; Stumm & Morgan, 1996; Brown & Parks, 2001; Brown & Sturchio, 2002). In the
majority of cases, most metal-solid interactions can be considered dynamically
reversible, such that changes in certain geochemical conditions can lead to release of
metals back into solution. A further, defining caveat to the framework of metal-solid
interactions is that neither metal (s) nor solid phases can be considered collectively,

SGM symposium 65

Biofilms and metal geochemistry 13

as the behaviour observed will be predicated upon which specific metal(s) and solid
phase(s) are involved, reflecting relative affinities, differing mechanisms of seques-
tration and controls on reactivity

Aqueous geochemistry has served as the foundation discipline for much of the environ-
mental metal literature spanning freshwater and marine metal behaviour, subsurface
metal migration and the bioavailability and toxicity of metals to biota (e.g. Lee et al.,
2004; Martino et al., 2004; Basta et al., 2005). The defining controls of pH, redox status
and ionic strength (Allard et al., 1987; Johnson, 1990; Brown & Parks, 2001; Hassler
et al., 2004) on metal partitioning reflect their influence on: (i) formation/dissolution
of key solid phases for sequestering trace metals (e.g. carbonates, sulfides and oxy-
hydroxide minerals, organic matter), as well as metal speciation and associated
behaviour; and (ii) sorption reactions.

The nature and extent of a mineral's importance for metal sequestration in any given
system will reflect its respective relative abundance and spatial distribution in the
geosphere. The heterogeneous distribution of key mineral phases reflects the differing
conditions and controls on their formation and dissolution and also provides some
delineation of which minerals are likely to occur and play a role in metal behaviour
across differing geochemical environments. Minerals can sequester metals through
sorption at their surfaces as well as through direct precipitation or solid-solution
formation within a mineral (Stumm & Morgan, 1996; Brown & Parks, 2001).
The mechanism by which a metal becomes associated with a solid will determine its
relative sensitivity to remobilization and the conditions under which release may occur.
Sorption reactions can be considered reversible: metal uptake to solids is not a perm-
anent phenomenon, as small fluctuations in either pH or solution metal(s) concen-
tration can often lead to release of metals from a solid surface (Warren & Haack, 2001).
Of key importance to note is that both the relative affinities of different solid phases for
different metal ions, as well as the controls on their reactivity, differ. Thus, predicting
metal solid-solution behaviour requires information on the types and abundances of
solid phases present, the specific metal(s) and respective concentrations involved, as
well as system geochemical conditions.

Important solid phases

In general, there are four major solid phases that can be considered most relevant
for metal sequestration in aqueous systems: (i) carbonates, (ii) oxyhydroxide minerals
(Fe and Mn); (iii) organic matter (live and dead) and (iv) sulfidic minerals (Tessier et al.,
1979). Many studies have shown that these typically comprise the majority of solid
phases involved in metal uptake (Tessier 6c Campbell, 1988; Warren & Zimmerman,
1994; Brown et al., 1999; Filgueiras et al., 2002). Another final sedimentary pool is often

SGM symposium 65

14 L.A.Warren

referred to as the refractory or mineralized component (Tessier et al., 1979), reflecting
metals held within crystalline mineral lattices. However, metals associated with this
refractory pool are not likely to be released under normal geochemical fluctuations and
thus its utility is usually for determining the absolute total mass of metal associated
with a given sediment. Changes in pH, redox or ionic strength can affect metal uptake
to the previous four fractions, and thus much research has focused on understanding
metal associations with these specific sedimentary pools, which commonly occur as
heterogeneous mixtures in environmental sediments and soils.

Carbonate minerals [e.g. limestone, CaC0 3 (s), or dolostone, Ca,MgC0 3 (s)] are typi-
cally found in circumneutral to alkaline pH environments, where their concentrations
can effect significant metal uptake, especially for certain elements which show high
affinity for carbonates, such as Cd and U. Carbonate minerals are susceptible to acid
dissolution, and thus decreasing pH can lead to dissolution and subsequent release of
carbonate-associated metals back into solution.

In contrast, sulfidic minerals such as pyrite [FeS(s)] tend to be associated with anoxic
sedimentary systems where the necessary reducing conditions are present for their
formation and/or preservation. Sulfidic minerals are commonly formed as a con-
sequence of sulfate reduction in anoxic sediments, which leads to a build-up of reduced
S in the sediment pore waters, which, when sufficient concentrations build up to satu-
rate with respect to various metal sulfide mineral phases, can then precipitate as a metal
sulfide [e.g. CdS(s), CuS(s), FeS (s) ; Warren et al., 1998]. Once formed, sulfide minerals
tend to be fairly robust, even in the presence of oxygen, to oxidative dissolution, except
where direct microbial catalysis is involved. Sulfidic minerals are also a key component
in mine-associated waste or tailings rock, driving the production of AMD.

The Fe and Mn oxyhydroxides [e.g. Fe(III) oxides, FeOOH(s) and Mn(III,IV) oxides,
MnOOH(s)] are highly metal-reactive and are typically widespread throughout most
environmental systems, making them a dominant solid phase controlling metal
behaviour in the environment. Oxyhydroxides of Fe and Mn are redox-sensitive, with
the more oxidized form of both elements, Fe(III) and Mn(III,IV), forming a solid phase,
while the reduced forms, Fe(II) and Mn(II), are soluble. Thus changes in redox
conditions will profoundly affect the concentrations of these solids and associated
metal behaviour. Although it should be noted that reductive dissolution of both
oxyhydroxides proceeds much more quickly with microbial catalysis than without
and that, while abiotic oxidation of Fe(II) to Fe(III) [and associated hydrolysis and
formation of FeOOH(s) solid particles] occurs spontaneously above pH 3 in the
presence of oxygen, abiotic Mn(II) oxidation is extremely slow below pH 9 (e.g. months
to years; Morel & Hering, 1993; Stumm & Morgan, 1996).

SGM symposium 65

Biofilms and metal geochemistry 15

Natural organic matter (NOM), both living (e.g. micro-organisms) and dead (both
labile and refractory, e.g. humic and fulvic acids), is an efficient, but often dynamically
reversible, metal sequestration pool (Gustafsson et aL, 2003; Smiejan et aL, 2003;
Jacob & Otte, 2004; Boullemant et aL, 2004). Sequestration by NOM can be especially
important in systems of high NOM concentration such as wetlands and eutrophic
lakes, typically through sorption reactions (discussed subsequently) associated with
their highly electrically charged surfaces. Decomposition of organic matter often leads
to release of associated metals into solution and thus the ability of the NOM pool in
any given system to hold onto metals will reflect the relative magnitude of decom-
positional processes.

Sorption reactions

The profound importance of interfacial reactions in controlling the ultimate partition-
ing of metals between the solid and solution phases is well recognized (Brown & Parks,
2001; Warren & Haack, 2001; and references therein). These interfacial or sorption
reactions involve attraction or bonding between a charged solution metal ion and a
charged functional group on the surface of the particle (Fig. la) and are viewed to
proceed analogously to complexation reactions involving charged dissolved species in
solution (Buffle, 1988). Mineral surface reactivity is described by physical character-
istics such as size (surface area to volume), crystallinity and nature and density of react-
ive binding sites (Brown et aL, 1999; Martinez & McBride, 1998). Organic matter
reactivity reflects both these same geochemical particle characteristics and, for live cells,
metabolic pathway and level of activity, which can shape external and internal micro-
geochemical environments, influence mineral formation and dissolution and alter
surface charge characteristics and thus potential reactivity (e.g. Ferris et aL, 1989, 1999;
Nelson et aL, 1995; Parmar et aL, 2000; Philip et aL, 2000; Templeton et aL, 2003a, b;
Brown etaL, 2004).

This particle surface charge is imparted by surface acid functional groups that variously
deprotonate or protonate depending on their pK a values. Commonly, mineral surface
functional groups are hydroxyl groups, which often display a spectrum of pK a values
reflecting the specific hydroxyl group co-ordinational environment on the particle (e.g.
edge or defect sites within the crystal structure). Micro-organisms also carry a surface
charge associated with their cell wall, exopolysaccharide (EPS) or sheath, arising from
constituent functional groups (e.g. carboxylic, phosphoryl, amine as well as hydroxyl;
Cox et aL, 1999). Given the lower pK a values particularly associated with carboxylic
groups (pK a ~4), micro-organisms almost always carry a net negative surface charge in
most aqueous solutions. Microbial ability to sorb cationic metal ions makes them a
highly relevant sorbent, especially in low-pH environments where important mineral

SGM symposium 65

16 L.A.Warren

(a)

Anionic sorption

/"~NMC£

Net particle charge

Surface functional gr oup H V" H

Cationic sorption

Particle surface

Solution

Interracial
region

(b)

£U
-Q

O
t/1

NOM /

' Fe oxide

Fig. 1. Interfacial reactions involving particle-surface functional groups and solution ions play
a defining role in controlling metal partitioning between solid and solution phases, ultimately
influencing resulting metal behaviour, (a) Particle surfaces in aqueous systems carry a net overall
charge reflecting the acid-base characteristics of their functional surface groups, dependent on
system pH conditions. Generally, as pH decreases, greater anionic sorption occurs as particle surface
charge becomes net positive, reflecting greater uptake of protons. In contrast, as pH increases, a
greater net negative charge exists on particle surfaces so that cationic solution species such as
divalent metal ions (e.g. Cd 2+ ) sorb more effectively to the negatively charged particle. The pH at
which individual particles (e.g. different minerals, organic matter, dead or alive) become negatively
or positively charged is dependent on the structural composition of the functional groups involved.
Typically, most mineral functional groups are hydroxyl groups (as shown), although organic matter
can have several types of functional groups with widely differing p/C a values (e.g. carboxylic versus
amine groups; Coxef a/., 1999). (b) Generalized sorption edges shown for three important metal
sorbents in natural systems. Most cationic elements show a characteristic exponential increase in
sorption over a relatively small pH range for a given solid, reflecting pH-p/C a relationships for that solid.
Note that the pH at which the 'edge' effect occurs for a given solid will vary depending on the metal
involved. Mn oxyhydroxides and organic matter are typically negatively charged at much lower pH
values than Fe oxides, reflecting the differences in the acid-base characteristics of their respective
surface functional groups, and thus can sorb cationic species at much lower pH values than can
Fe oxyhydroxides, which, in contrast, are often effective anionic sorbents at low pH.

SGM symposium 65

Biofilms and metal geochemistry 17

sorbents (e.g. Fe oxyhydroxides) are less effective. Further, cells need be neither viable
nor intact to act as metal sequesters (Ferris et al., 1989), as the functional groups
associated with their cell-wall structures can still sorb metals (often with higher
sorptive capacities) when the cells die or lyse (Urrutia, 1997).

Most particles in solution carry a charge and, as that surface becomes more net
negatively or positively charged (many mineral surfaces are amphoteric), their relative
abilities to sorb cationic or anionic species, respectively, increases. Thus, the over-
whelming importance of pH in regulating trace-metal sorption across systems has been
widely accepted (Stumm, 1992; Stumm & Morgan, 1996). As pH increases, there is a
greater net deprotonation of particle surfaces and associated sorption of cationic metal
species, partitioning them to the solid, and less mobile, bioavailable and toxic, phase. In
contrast, as pH decreases, there is greater sorption potential for anionic species
reflecting the greater positive charge associated with particle surfaces (the exact pH at
which this occurs is solid dependent). Further, adding to the complexity in dealing with
heterogeneous solid and metal mixtures as is observed in the environment, 'edge' effects
are observed for metal sorption to a given solid (Fig. lb), whereby sorption increases
substantially over a relatively narrow pH range as the pH of the system moves above the
pK a for dissociation of the specific solid's surface acid-base sites (Dzombak & Morel,
1990). However, these pH edges are both solid dependent, reflecting the different pH
values at which a surface becomes net negatively or positively charged (Fig. lb), and
metal dependent, reflecting differential affinities of specific elements (often correlated
to a metal's ability to hydrolyse). As is evident from Fig. 1, organic matter, e.g. micro-
organisms, and Mn oxyhydroxides are able to sorb cationic species at much lower pH
values than Fe oxyhydroxides (reflecting their different pH values, or pH at which
their surfaces shift between net negative and net positive overall charge; Fig. 1). This
makes them of key importance in low-pH systems, whilst Fe oxyhydroxides would be
more important for anionic species under lower pH conditions. Typically, these
sorption edges are also element specific for a given solid phase, indicating that selective
affinity occurs. This negates the possibility of a generalized model for metal sorption
independent of either the metal or solid involved.

In brief, the role of ionic strength as a control in metal solid-solution partitioning
derives from its influence on the potential electrostatic attractions that occur in
solution, as well as on the nature of the interfacial region of the solid surface (Puis
et al., 1991; Lores & Pennock, 1998). The most labile fraction of solid-associated
metals is often referred to as 'exchangeable', referring to those metals that are
electrostatically attracted to particle surfaces, but not necessarily bonded to a specific
site on the surface or to a specific sedimentary component within a mixed sediment

SGM symposium 65

18 L.A.Warren

pool. This fraction is the most easily released back into solution with any changes
in solution conditions. In freshwater systems, increasing ionic strength increases
both solution interactions, thereby decreasing element activity and compressing the
electrical double layer surrounding the particle, and the associated sorptive uptake of
metals (Drever, 1997).

The role of redox status

Changes in redox status can affect both solids and metals that are redox sensitive,
influence partitioning between the solid and solution pools and thus change metal
behaviour. Oxyhydroxide minerals are particularly sensitive to changes in redox status,
such that reducing conditions can lead to their reductive dissolution and the associated
release of any sorbed metals into solution. Oxidizing conditions would favour
(especially for Fe) solid formation and associated sequestration of metals. Reductive
dissolution of Fe and Mn oxyhydroxides is commonly driven by micro-organism Fe and
Mn reduction coupled to organic matter degradation (see equation below; note that
'CH 7 0' represents organic matter), where reduction of the oxidized Fe and Mn in the
oxyhydroxide mineral leads to dissolution of the solid and release of any associated
metals from the oxyhydroxide into solution:

4Fe(III)OOH(s) + CH 2 + 8H + ^4Fe(II) + C0 2 +7H 2

Typically, reductive dissolution of both oxyhydroxides, as well as oxidative dissolution
of sulfidic minerals or organic matter, effectively proceeds only with microbial cata-
lysis: abiogenic reaction rates are extremely slow (e.g. Morel & Hering, 1993; Stumm &
Morgan, 1996; O'Day et al., 2004). Further, some redox-active (and often therefore
bioactive) metals such as U, Cr and As show profoundly different behaviour depending
on their redox status [e.g. U(VI) or U(IV), Cr(VI) or Cr(III), As(V) or As(III)]. For
instance, Cr speciation is controlled by a shifting array of processes that include redox
transformations, precipitation/dissolution and sorption/desorption reactions, all of
which are intimately tied to both oxidation state and the geochemical status of a system
(Bartlett, 1991; Richard &; Bourg, 1991; Baruthio, 1992). In its VI oxidation state, Cr
commonly occurs as the ligand CrOj - , which is highly mobile, toxic and soluble (Felter
& Dourson, 1997) and is generally controlled by sorption-desorption reactions at
lower concentrations. Cr(VI), as chromate, is sorbed to minerals under lower pH
conditions, when surface sites are positively charged. For example, Zachara et al. (1989)
have shown that the degree of chromate sorption is greater in acidic soils and in
subsurface media containing iron oxyhydroxides and clays. The reduced form of Cr,
Cr(III), is insoluble in aqueous environments. Precipitation of Cr(OH) 3 or (Cr,Fe)-
(OH) 3 compounds commonly controls its behaviour.

SGM symposium 65

Biofilms and metal geochemistry 19

THE OVERLAP BETWEEN MICROBIOLOGY AND METAL
GEOCHEMISTRY

The role of emerging technologies

It is only a small step across discipline divides to see the very real and often significant
overlap between microbial activity in the environment and metal geochemistry In large
part, the recognition that microbial activity can affect metal geochemistry has grown
with our ability to examine the linkages at the appropriate scale, i.e. the micron level.
Our ability to effectively probe and 'image' at the appropriate scale of resolution, using
such techniques as X-ray absorption spectroscopy (XAS; Brown & Sturchio, 2002),
atomic force microscopy (AFM) and scanning transmission X-ray microspectroscopy
(STXM; O'Day, 1999), has increased our understanding of the mechanisms involved in
solid-metal interactions. In addition, perhaps the most powerful tool which revolution-
izes our ability to examine micro-organisms in an environmental context has been the
development of culture-independent molecular-biological tools for phylogenetic
characterization (e.g. Pace, 1997; Kaeberlein et al., 2002), which have permitted evalu-
ation of community genetic diversity.

Relevance of microbial functional metabolism

Micro-organisms can influence metal behaviour through a number of processes,
reflecting the direct link between types and rates of metabolic activity and geochemical
conditions, processes and reaction rates (Fig. 2). The occurrence of micro-organisms in
almost every environment investigated (e.g. Bennett et al., 2000; Bond et al., 2000;
Chapelle et al., 2002; Takai et al., 2004) hints at their widespread influence on geo-
chemical processes. Unlike higher eukaryotic organisms, micro-organisms are not
restricted by geographical barriers (Finlay, 2002). Micro-organisms seek energy and
carbon to survive and grow. In the environment, they select for geochemical conditions
which are favourable and commonly catalyse geochemical processes for their growth
(Nealson & Stahl, 1997; Nealson, 2003). While thermodynamics sets energetic
constraints on microbial activity, i.e. organisms must live within the realm of reactions
that are thermodynamically favourable, it is increasingly clear that the extent of this
domain of possible reactions in the geosphere is not yet completely described (e.g.
Spear etal., 2005).

Beyond the simple aerobic versus anaerobic demarcation in metabolism, there are
myriad potential metabolic pathways (i.e. electron donors and acceptors) with
relevance for metal behaviour through which micro-organisms can selectively shape
geochemical processes dependent on the particular redox couples that they catalyse.
Such processes as nitrification and denitrification, methanogenesis, methanotrophy,
sulfur/iron/manganese oxidation and reduction (Holt & Leadbetter, 1992; Nealson &

SGM symposium 65

20 L. A. Warren

O
a*

Mechanism

Indirect

Fe/Mn redox

reactions

»*

Fe/Mn oxides

Dissolution
\

Formation

Geochemical

microenvironments

Su rf ace
reactions
/
/
I

Release of M

Sorption/immobilization of M

Increased mobility,
bioavailability, toxicity

Decreased metal impacts through
decreased solution partitioning

Fig. 2. Microbial influence on metal geochemistry can occur through both indirect and direct
mechanisms. In both instances, microscale influence by microbial metabolic activity can scale to bulk-
system impacts for metal behaviour, depending on the level of metabolic activity. Indirect mechanisms
refer to general changes in the local geochemical environment associated with metabolic activity that
then directionally affect metal solid-solution partitioning. Direct mechanisms refer to those specifically
catalysed by microbes, e.g. Fe oxidation will lead to the formation of Fe oxyhydroxides and will likely
lead to metal sequestration, while Fe reduction will dissolve Fe oxyhydroxides, liberating any
associated metals into solution, increasing their mobility, bioavailability and likely impact.

Stahl, 1997) and hydrogen-based metabolism (Chapelle et ai, 2002; Spear et al., 2005)
can all substantively affect metal speciation and thus behaviour (Fig. 3).

In addition to the ecological factors of sufficient nutrients, water and protection from
predation, probably one of the most important parameters defining metabolic habitats
is oxygen concentration and/or flux. This is also a defining geochemical parameter
controlling the oxidative state of reactive geochemical components and thus segre-
gating differing geochemical processes within the environment. Thus, it is clear that
aerobic and anaerobic organisms will select for differing environments on the basis of

SGM symposium 65

Biofilms and metal geochemistry 21

Metal behaviour -
fate, toxicity, bioavailability etc.

Metal toxicity feedback
on biological activity f

Solid-solution partitioning

Dynamic process determined by

identity and concentration of
both solid and solution players
and reflecting system
geochemical conditions

Biological activity

Dependent on extent and
type of microbial activity;
influences geochemical
properties through;

(i) PASSIVE: influencing
micro-geochemical
conditions through
metabolism(s)

(ii) ACTIVE: specific

metabolic control -
solids or metals
i.e. mineral dissolution,
metal release

Biological—

geochemical
linkages
interdependent:
feedback occurs

Solid-sorption interface! reactions
L^ 4 pH leads to:

* Solid-associated M z+

Redox status

Solids: redox sensitive

L>. Formation/dissolution

Metals: some redox-active metals

L* So I i d aff i n i ty red ox-state
dependent

CONTROL

Biological

Geochemical

Will reflect the nature and magnitude
of micro-organism activity

Fig. 3. Understanding the controls of metal geochemistry is increasingly interpreted as a dynamic
array reflecting the interactions between microbial activity and the important controls on metal
solid-solution partitioning. Feedback mechanisms of local geochemical conditions on existence,
activity and growth of micro-organisms also occur. The relative dominance of microbial or classic
geochemical controls on metal behaviour will reflect the extent of microbial activity in a given
system or at a given time.

oxygen concentration, which similarly stratifies geochemically important solids and/or
forms of elements important in metal behaviour. Thus, metal impacts will be selectively
dependent on both the types of metabolism and levels of activity occurring in any
micro-environment. Regardless of the specific metabolic guild involved, the common
and widespread occurrence of microbial control on important geochemical processes is
increasingly demonstrated (e.g. biogenic Fe oxide formation: Chafetz et al., 1998;
Newman & Banfield, 2002; Haack & Warren, 2003).

SGM symposium 65

22 L. A. Warren

As field investigations (e.g. Smiejan et al., 2003; Brown et al., 2004; Labrenz & Banfield,
2004; Meylan et al., 2004) and associated laboratory experimentation to probe mech-
anisms and effects (e.g. Warren & Ferris, 1998; Parmar et al., 2000; Moreau et al., 2004;
Tani et al., 2004a, b; Templeton et al., 2003a, b) demonstrate, micro-organisms influ-
ence trace-metal behaviour through both passive and active mechanisms (Fig. 2).
Passive or indirect mechanisms are viewed as those associated with general metabolic
activity, i.e. where the particular metabolic pathways and levels of activity associated
with a given microbial community induce changes in geochemical conditions within the
immediate micro-geochemical environment. These can then influence metal behaviour,
e.g. increasing pH would generally favour sorption of more metals to the solid com-
partment. Active or direct mechanisms are defined as those that reflect direct catalysis
by micro-organisms of geochemical processes that control metal behaviour, e.g.
reduction of Fe resulting in the dissolution of Fe oxide minerals and the associated
release of metals into solution. Increasing evidence indicates the ability of many micro-
organisms to directly metabolize many bioactive (redox-active) elements such as U, Cr
and As (e.g. Lores & Pennock, 1998; Lovley & Anderson, 2000; Liu et al., 2002), which
alters the behaviour of these metals in the process.

Micro-organisms can also profoundly influence metal behaviour through associated
impacts on mineral formation and dissolution (Konhauser et al., 1993; Bennett et al.,
2000; Haack & Warren, 2003; Templeton et al., 2003a; Moreau et al., 2004). However,
it is also clear that the geochemical conditions in which micro-organisms find them-
selves can influence their viability, activity and growth and, commonly, interactions
between microbial behaviour and geochemistry exert feedback on each other (Warren
& Haack, 2001; Warren & Kauffman, 2003). Thus, there is the potential for geo-
chemical feedback on biological processes that will in turn influence the metal
behaviour observed, which clearly indicates that the linkages between microbial activity
and geochemical behaviour are complex and dynamic (Fig. 3).

Biofilms: existence, structure and function

It is increasingly clear that many micro-organisms self-organize into diverse biofilm
communities (Davies et al., 1998; Branda et al., 2005), which can produce complex and
dynamic internal geochemical conditions based on the particular array of metabolic
processes occurring within the biofilm (Little et al., 1997). Biofilms form at interfaces
between solid and solution phases, overlapping with the same dynamic reactive zone
that controls metal partitioning. Biofilms can also form on particles suspended in
solution (Leppard et al., 2003, 2004; Roberts et al., 2004) or at the bed-sediment surface
(Fig. 4a; Vigneault et al., 2001; Haack & Warren, 2003). In either scenario, these
biologically controlled layers can selectively control metal-solid interactions depending
on the particular geochemical conditions created by the biofilm's specific metabolic

SGM symposium 65

Biofilms and metal geochemistry 23

array (Fig. 4b). Biofilms must provide a selective advantage for the organisms that form
them. Typically, biofilms are thought to provide structural protection, through the
associated EPS matrix, from predation and/or impacts of potentially toxic con-
taminants (Lawrence et ai, 1995, 1998; Neu & Lawrence, 1997; Davies et al., 1998).

Mixed-community biofilms exhibit stratified metabolic functions reflecting the ener-
getic array of reactants and products which link these microbial communities together.
Further, minerals are commonly associated with biofilms (e.g. Douglas 6c Beveridge,
1998; Ferris et ai, 1999; Haack 6c Warren, 2003; Templeton et ai, 2003a; Moreau et ai,
2004), either through passive biomineralization, where the EPS and/or cells of the bio-
film act as nucleation templates for minerals to form, or actively, where specific minerals
are catalysed through specific functions of the biofilm. In either scenario, minerals are
commonly associated with biofilms as either bioreactants and/or bioproducts of micro-
bial metabolism.

Biofilm metabolic links to metals

Since biofilms represent a concentrated cell mass with significant metabolic activity,
their occurrence in the environment can lead to a significant biological overlay on
geochemically reactive processes relevant to metal behaviour that will be governed by
the ecological controls on biofilm formation and function. While it is generally held
that most systems are at or near equilibrium with respect to acid-base reactions (e.g.
pH), it is very clear that, without microbial catalysis of important redox geochemical
reactions, many would not occur over relevant timescales due to the extremely slow
kinetics of many of these reactions in the geosphere. Another clearly important factor
to consider is that geochemical processes that proceed abiogenically (geochemically
controlled) often do so under different environmental conditions and controls and with
potentially different geochemical outcomes from biogenically micro-organism-
catalysed processes (e.g. Warren 6c Ferris, 1998; Haack 6c Warren, 2003; Tani et ai,
2004a, b).

Environmental biofilm metal geochemistry

Biofilms are common in metal-contaminated and/or low-pH systems (Southam 6c
Beveridge, 1992; Ledin & Pedersen, 1996; Nordstrom & Southam, 1997; Edwards et
ai, 1999, 2000; Hunt et ai, 2001; Leveille et ai, 2001; Vigneault et ai, 2001; Baker et ai,
2004; Labrenz 6c Banfield, 2004), indicating that these extreme conditions provide
favourable conditions for adapted microbial communities to flourish. They also
indicate that there is significant potential for biofilms in these environments to impact
metal geochemistry. As yet, there are not many studies that specifically evaluate the
linkage between metal behaviour and biofilm metabolic functional activity However,
acid mine drainage (AMD) has long been recognized as one of the earliest environ-

SGM symposium 65

24 L. A. Warren

SOLID:

suspended AW

particles

-►-., SOLUTION

Metal partitioning? ^
Biofilm r

SOLID: bed

sediments

(a)

SOLUTION

EMR

Metal
L_

(b)

Sequestration

Solubilization

BIOFILM

Photoautotrophs

H + +C0 2 + P + N

**.

(CH 2 0) n

Mn oxi d-fze rs Qxic-anoxic interfa ce M n 2+ ■

Fe oxidizers

Sibxidizers H 2 S

Fe 2 '

-► S<

Fe/Mn reducers
S reducers/

> s 2 or

Mn 2+ ^-
Fe 2+ <«-

]" ,, "j--Bjomineral formation

Mn oxides**..

Fe oxides
Fe 3+ S0 4 minerals

"Pr 2 5-*tn7Z.S° < S 2 0§"

disproportionators """* <;

Mineral dissoluticfh

Fig. 4. Biofilms form at solid-solution interfaces, either at bed-sediment solution interfaces or at
suspended particle surfaces (a), thus overlapping with the reactive interface for metal solid-solution
partitioning. These biologically controlled interracial structures can profoundly influence the nature,
extent and types of geochemical processes that occur, particularly with respect to those that influence
metal partitioning in a manner dependent on the particular array of metabolic functions expressed
within a given biofilm (b). As shown in (b), in this case a biofilm with phototrophs occurring at the
surface driving 2 fluctuations within the surface biofilm, a generalized microbial biofilm is typically
a stratified community that will show depth-dependent metabolic pathways typically as a function
of oxygen status, which will also strongly influence the types of geochemical processes and the
nature of both solid and solution elements that can occur at any given depth within the biofilm

SGM symposium 65

Biofilms and metal geochemistry 25

ments for evidence of microbially controlled geochemical processes (e.g. Ledin &
Pedersen, 1996; Nordstrom & Southam, 1997), providing some evidence of the shift in
geochemical behaviour with micro-organism activity. AMD is strongly driven by
microbially induced weathering and oxidation of sulfidic (e.g. pyrite, pyrrhotite)
minerals exposed during the mining of base metals and coal (Fig. 5). The oxidation of
sulfide-containing waste rock is significantly catalysed by iron- and sulfur-oxidizing
bacteria that are attached or in close proximity to the sulfide minerals. This oxidation
produces large amounts of acid. Recent investigation of AMD-associated biofilm metal
dynamics (Haack & Warren, 2003) indicates that these biological structures are
efficient metal sequesterers (Fig. 4b), through bacterially catalysed biomineralization
reactions that would not be predicted to occur abiogenically

While studies are emerging that provide direct evidence of both biomineralization by
biofilms (e.g. Haack & Warren, 2003; Templeton et al., 2003b; Labrenz & Banfield,
2004; Moreau et al., 2004) and metal uptake by biofilms (e.g. Templeton et al., 2001,
2003a; Vigneault et al., 2001; Haack & Warren, 2003), as yet relatively few have
integrated metal geochemistry dynamics with microbial community dynamics.
However, those studies that have been done demonstrate that the following two major
processes influence biofilm metal uptake. Firstly, the process of biofilm-associated
mineral formation, whether passive or active, often plays a key role in metal seques-
tration (e.g. Nelson et al., 1995, 1999; Ferris et al., 1999). However, subsequent
dissolution of biominerals, which tend to be small and amorphous, can lead to
dynamic temporal fluctuations in biofilm metal content as mineral formation and
dissolution are cyclically catalysed (Haack & Warren, 2003). Secondly, the organic
fraction of the biofilm can also play a substantive role in metal sequestration through
sorption to EPS as well as to cell walls (live and/or dead cells) and intracellular metal
uptake, accumulation and storage (Beveridge, 1989; Schorer 6c Eisele, 1997; Liinsdorf
et al., 1997; Vigneault et al., 2001; Haack & Warren, 2003; Brown et al., 2004; Meylan
etal.,2004).

('CH 2 0' represents organic carbon molecules; EMR, electromagnetic radiation). In an interlinked
fashion, the micro-organisms dependent on the geochemical components for their growth and
survival and the geochemical components influenced by the microbial catalysis of reactions driving
their consumption or production, a biofilm is a geochemical reactor that is strongly driven by biology
and overlaps directly with the reactive metal zone between solids and solution. The nature of the
metabolic processes that occur within each functional metabolic stratum can profoundly influence
key geochemical processes controlling metal behaviour, particularly those reactions involving the
formation and dissolution of metal reactive minerals such as Fe and Mn oxyhydroxides (as shown in
b). Thus, biofilm metal behaviour is likely to vary across systems, reflecting the specific interplay of
geochemistry and microbiology, as well as for a given biofilm structure in both a depth-dependent
and a temporal manner. If the biofilm has a phototrophic component (as exemplified in b), this will
show shifts in respiration dominance (and therefore 2 saturation) over diel timescales.

SGM symposium 65

26 L. A. Warren

FeS 2 + 3*50 2 + H 2
2Fe 2+ + 0-5O 2 + 2H +

FeS 2 +14Fe 3+ + 8H 2

Fe 2+ + 2SO|-+2H +

Biocatalysed
reactions

2Fe 3+ + H 2

(2)

DRIVE

15Fe 2+ + 2S0 2 " + 16H+ (3)

Abiotic
reaction

H 2 0, 2
micro-organisms

AMD creation:

highly acidic and

metal-laden

discharge

Fig. 5. The weathering of tailings, metal-sulfide waste rock material, is catalysed through bacterial
action (the diagram indicates that reactions 1 and 2 are biogenic and drive reaction 3, an abiogenic
process that liberates significant acid). Abiogenic oxidation/weathering of these reduced sulfidic
minerals, even when exposed to 2 , are exceedingly slow. The bacterium Thiobacillus ferrooxidans
catalyses oxidation of FeS (x) minerals through direct oxidation of FeS (x) (pyrite, FeS 2 , used here) and
minerals (equation 1 ) and through the oxidation of Fe 2+ to Fe 3+ (equation 2, increasing rates of Fe 2+
oxidation up to five orders of magnitude; Singer & Stumm, 1 970). This biogenically produced Fe 3+
then abiogenically catalyses further (more rapid) abiogenic weathering of FeS (x) (equation 3). It is
reaction 3 that drives significant acid production associated with AMD. In the process, acid is liberated,
metals are typically solubilized and resulting AMD is both acidic and metal laden. The diagram
indicates that the microbe-driven processes of (1 ) and (2) (top left half of diagram) drive abiogenic
Fe 3+ -subsequent weathering of the FeS (x) minerals.

CONCLUSIONS

Trace-metal behaviour is controlled by a complex and shifting array of processes that
include redox transformations, precipitation/dissolution and sorption/desorption
reactions, all of which are intimately tied to the geochemical status of a system and can
be selectively influenced by microbial activity in a community-dependent manner,
reflecting the particular functional array involved. Our understanding of how biology
and geochemistry integrate in such a dynamic array of potential interactions is, at best,
rudimentary However, it is clear that, in many environmental instances, an overlap
occurs between microbial activity and metal behaviour, reflecting the physical over-
lap of microbial biofilm habitat and reactive metal zones in the environment. Micro-

SGM symposium 65

Biofilms and metal geochemistry 27

organisms influence their immediate micro-geochemical environment as well as directly
catalyse geochemical transformations required for their survival and growth. In both
instances they can profoundly influence the metal behaviour observed. However, micro-
organisms are also intimately affected by their environment and thus feedback is
exerted in both directions as dynamic and fluctuating geochemical conditions influence
metabolic pathways, rates and times of expression. Thus, it is evident that under-
standing and predicting geochemical behaviour requires an understanding of the often
complex and dynamic linkages between ecological controls on where micro-organisms
exist, how they make their living and their level of metabolic activity, and the geo-
chemical processes they use to sustain themselves. Further, microbial plasticity ensures
that organisms not only exploit available redox couples, but they also respond to
microscopic variations in geochemical conditions, whilst simultaneously creating
heterogeneity from the output of their own metabolic activity (Finlay, 2002). Thus,
system-dependent and site-dependent microscale spatial and temporal heterogeneity
needs to be investigated in order to constrain at the geosphere level the complex
interactions of micro-organisms and geochemistry that result in the ultimate behaviour
observed.

The challenge for this emerging field is to define the relevant controls in an integrated
ecological and geochemical space and to develop sensitive techniques for their detection
and accurate quantification. Emerging techniques are providing new insight into 'which
organisms are present' in different environments, a fundamental question that was,
until recently, difficult to answer. The extension of this question to 'what are they doing'
is the current focus and the future of this area, as we seek to address mechanistically
how microbial metabolism or activity links directly to geochemical reactions and
shapes outcomes of relevant processes (Croal et al., 2004; Warren, 2004). As recog-
nition of the need to examine the types and levels of metabolic activity occurring in the
environment (e.g. proteomics, metallomics) has grown in appreciation of the challenge
of quantifying microbial links to geochemical reactions, increasing methodological
development has occurred. It is still a challenge to characterize metabolic pathways
with microbial strains only identified on the basis of 16S rRNA oligonucleotide
sequences, as so few environmental micro-organisms are currently in culture. This
means that, while an exponentially growing database of environmental strains is
genetically identified, the vast majority of these novel organisms remain undescribed in
terms of their metabolic pathways. Because of this hurdle, other techniques are focused
on evaluation of gene expression in the environment (dependent on their character-
ization) or on using isotopes to track community diversity, structure and function. The
field of molecular ecology is providing exciting new studies that track both molecular
diversity and function, with apparently promising selectivity, through both labelling
and probing of environmental processes using stable isotopes (e.g. Engel et al., 2004;

SGM symposium 65

28 L. A. Warren

Griffiths et al., 2004; Lueders et ai, 2004; Manefield et al., 2004; Pearson et ai, 2004;
Andersen et ai, 2005; Lu et al., 2005). Until such time as gene expression (i.e. functional
activity) for any specific organism in the environment is easily identifiable and
quantifiable, the use of isotopes as an additional tool to probe actively occurring
geochemical transformations linked to microbial activity appears highly promising.

ACKNOWLEDGEMENTS

This chapter presents a synthesis of ideas that have grown, coalesced and crystallized predomi-
nantly from the many discussions and interactions I have had with my students over the last
5 years. I would like especially to thank Dr Elizabeth Haack, Luc Bernier, Derek Amoves, Corrie
Kennedy, Tara Nelson and Lisa Melymuk. I would also like to acknowledge substantial funding
support from NSERC, CF1, OIT, McMaster University and Noranda/Falconbridge Ltd.

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Biofilms and metal geochemistry 33

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34 L. A. Warren

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SGM symposium 65

Minerals, mats, pearls and veils:
themes and variations in giant
sulfur bacteria

Neil D. Gray and Ian M. Head

School of Civil Engineering and Geosciences, Institute for Research on the Environment and
Sustainability and Centre for Molecular Ecology, University of Newcastle, Newcastle upon Tyne,
NE17RU, UK

MINERALS, MATS, PEARLS AND VEILS: A PANOPLY OF GIANT
SULFUR BACTERIA

The biogeochemical cycling of sulfur has been at the heart of microbial ecology since
the mid-19th century This is due, at least in part, to the striking forms of many of the
organisms involved in the transformation of reduced sulfur species. Giant sulfur
bacteria were among the earliest micro-organisms to capture the interest of micro-
biologists exploring the links between geochemical cycling of the elements and the
microbiota responsible. Consequently, giant sulfur bacteria were among the first
bacteria described. Organisms resembling Beggiatoa ('Oscillatoria alba') were des-
cribed as early as 1803 (Vaucher, 1803), but were included in the genus Beggiatoa some
time later (Trevisan, 1842). Thiothrix (Rabenhorst, 1865; Winogradsky, 1888), Achrom-
atium (Schewiakoff, 1893) and Tbioploca (Lauterborn, 1907) were all described by the
early 20th century and Winogradsky (1887, 1888) had already formulated the principles
of lithotrophic growth based on sulfide oxidation, from his work on Beggiatoa species.
Surprisingly for such conspicuous organisms, novel giant sulfur bacteria are still being
described (Guerrero et ai, 1999; Schulz et ai, 1999).

Achromatium

Bacteria of the genus Achromatium are remarkable. Cells of up to 125 ^im in length
have been reported (Babenzien et al., 2005; Head et al., 2000a) and, in addition to
characteristic sulfur globules that become visible on treatment of the cells with dilute
acid, their large oval cells are typically filled with enormous inclusions of calcium
carbonate (Fig. 1). As with many giant sulfur bacteria, they have as yet eluded

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

36 N. D. Gray and I. M. Head

Fig. 1. DAPI (4,6-diamidino-2-phenylindole)-stained preparation of a crudely purified suspension of
Achromatium cells. The diatom cells illustrate the large size of the Achromatium cells.

cultivation in the laboratory. Nevertheless, the ability to physically enrich cells from
environmental samples (De Boer et aL, 1971; Head et aL, 2000a) has permitted the
inference of some physiological properties (Gray et aL, 1997, 1999a, b, 2000, 2004).
Achromatium species are capable of oxidizing reduced sulfur species completely to
sulfate (Gray et al., 1997). This is a feature of bacteria capable of conserving energy
from sulfur oxidation, in contrast to chemoheterotrophic sulfur bacteria that appear to
utilize sulfide oxidation as a means to detoxify sulfide or neutralize reactive oxygen
species produced during aerobic metabolism (Burton & Morita, 1964), although
considerable ambiguity and debate still surround the precise role of sulfide oxidation in
these bacteria (Strohl, 2005). Recent evidence also suggests that, although oxygen
is probably the preferred electron acceptor for Achromatium species, nitrate is also
probably used as an alternative electron acceptor (Gray et aL, 2004). The use of nitrate
as a terminal electron acceptor is a trait shared with several other giant sulfur bacteria.
Although the name of only one species of Achromatium has been validly published
{Achromatium oxaliferum; Schewiakoff, 1893), it is clear from comparative analysis of
16S rRNA gene sequences recovered from Achromatium cells physically enriched from
sediment samples that several distinct species exist and, typically, several species coexist
in a single sediment (Babenzien et aL, 2005; Gray et aL, 1999b; Head et aL, 1996). All
Achromatium species investigated to date form a distinct monophyletic group within

SGM symposium 65

Ecology of giant sulfur bacteria 37

Large marine Beggiatoa, Thiomargarita and Thioploo

d Jhioploca ingrica and other freshwater Thbploca
Thin, cultured marine Beggiatoa

Thin, freshwater Beggiatoa

' Gammaproteobacteha '

Sulfur-oxidizing symbionts

Sulfur-oxidizing symbionts
^^fl Thtomkmspira

Thiothrix

Achromatium

Aliochromatium/Tb ermo en romatium
Sulfur-oxidizing symbionts

Addithiobatillus

Thiovulum rnajus

SutfuKttidiang symbionts ' Eps il onproteobact eria'

Fig. 2. Phylogenetic tree of representatives of the sulfur-oxidizing bacteria from the

' Gammaproteobacteria' and ' Epsilonproteobacteria' , including the giant sulfur bacteria from the

genera Achromatium, Beggiatoa, Thiomargarita, Thbploca, Thiothrix and Thiovulum.

the 'Gammaproteobacteria' (Fig. 2) and, in most analyses, appear to be related to
anoxygenic, photosynthetic bacteria of the family Cbromatiaceae. However, the fifth
release of the taxonomic outline of prokaryotes from the second edition of Bergeys
Manual of Systematic Bacteriology lists Achromatium in the order 'Thiotricbales' as
genus II within the family 'Tbiotrichaceae' (Garrity et al., 2005).

Beggiatoa

In contrast to Achromatium species, which are found as large populations of free-living
unicells, Beggiatoa species often form extensive, conspicuous mats on the surface of
sediments or associated with sulfur springs, hydrothermal vents or cold-seep environ-
ments (Strohl, 2005; Teske &C Nelson, 2004). The mats can be extremely thin and
compact (hundreds of micrometres) or much more diffuse and thick (centimetres),
depending on the prevailing hydrodynamic conditions, which dictate whether substrate
transport to the cells is diffusion-limited (Gunderson et al., 1992; Jorgensen &
Revsbech, 1983). Indeed, there have been reports of B^gg/^^o^-dominated mats as thick
as 30-60 cm in some marine hydrothermal systems associated with populations of tube

SGM symposium 65

38 N. D. Gray and I. M. Head

worms (McHatton et al., 1996; Nelson et al., 1989). Several species have been obtained
in pure culture, but the largest mat-forming Beggiatoa species discovered have yet to be
grown in pure culture. Beggiatoa species are physiologically varied: some strains have
been shown to be obligate chemolithoautotrophs, whilst others are facultative chemo-
lithoautotrophs or mixotrophs (Hagen & Nelson, 1996; Nelson & Jannasch, 1983).
Strains capable of autotrophic growth have been obtained from marine environments,
whereas evidence to date suggests that freshwater isolates are primarily heterotrophs
and evidence for lithoautotrophic growth of freshwater Beggiatoa isolates has been
contentious (Strohl, 2005). Nevertheless, lithoautotrophic growth has been shown for
more recently isolated freshwater Beggiatoa strains (Grabovich et ai, 2001; Patritskaya
et al., 2001) and evidence for biochemical control of the switch from heterotrophic
growth to mixotrophic or lithoautotrophic growth has been provided (Eprintsev et al.,
2003, 2004; Stepanova et al., 2002). The heterotrophic freshwater strains produce
characteristically thin filaments, as do cultivated marine strains capable of autotrophic
growth. Neither of these thin-filament types produces large vacuoles and both have
been cultivated in the laboratory. Beggiatoa cells with wide filaments (up to 200 ^im in
diameter) have yet to be grown in laboratory culture and are characterized by large
vacuoles that are believed to play a role in storing nitrate, which is used as an oxidant in
these enormous bacteria. However, recently studied vacuolated filaments of giant
bacteria that resemble Thiotbrix species morphologically, but are related most closely
to Beggiatoa species, do not appear to store nitrate, suggesting that the role of large
intracellular vacuoles may be more complex than simply providing a repository for
large amounts of oxidant (Kalanetra et al., 2004). Most large-vacuolated Beggiatoa
species accumulate high concentrations of nitrate (Kalanetra et al., 2004; McHatton
et al., 1996; Mufimann et al., 2003), which is used as an oxidant either by reduction to
ammonia or through denitrification, thus providing a link between the biogeochemical
cycles of nitrogen and sulfur (Teske &C Nelson, 2004). Like Achromatium species,
Beggiatoa species belong to the "Gammaproteohacteria\ where they form a mono-
phyletic group with Thioploca and Thiomargarita species (Fig. 2; Kalanetra et al., 2004;
Kojima et al., 2003 ; Mu&nann et al., 2003 ; Teske et al., 1996) . The uncultured vacuolate
bacteria {Beggiatoa, Thioploca and Thiomargarita species) themselves form a distinct
phylogenetic group within this, with freshwater and marine thin-filament Beggiatoa
species each occupying distinct lineages (Teske 6c Nelson, 2004). It is also clear that
there is considerable diversity within communities of Beggiatoa, which often exhibit
filaments with different diameters (Nelson et al., 1989).

Thioploca

In many respects, Thioploca species resemble Beggiatoa species, especially the large-
vacuolated Beggiatoa found in marine environments. Members of the genus Thioploca

SGM symposium 65

Ecology of giant sulfur bacteria 39

are differentiated from those of the genus Beggiatoa on the basis that trichomes of
Thioploca species are usually enclosed in a thick polysaccharide sheath. Whilst some
Beggiatoa strains have been isolated in pure culture, no Thioploca strains have been
grown axenically in the laboratory. Like Beggiatoa strains, they form multicellular
filaments that can be several centimetres long. Several morphotypes can be differ-
entiated on the basis of filament diameter (Schulz et al., 1996) and these correspond to
different species, defined on the basis of 16S rRNA gene sequence identity (Teske et al.,
1996). Moreover, most Thioploca species described to date, in common with large
Beggiatoa species, harbour vacuoles that are presumed to store nitrate, which is used as
an oxidant by both Beggiatoa and Thioploca species (McHatton et al., 1996). These
occur even in Thioploca species with relatively thin (2-5 ^m diameter) filaments
(Zemskaya et al., 2001), whereas thin Beggiatoa species apparently do not have large
vacuoles (Teske & Nelson, 2004). rRNA gene sequence-based analyses of Thioploca
isolates recovered from naturally occurring mats have shown that they are related
closely to, and form a monophyletic group with, the large-vacuolated Beggiatoa species
and members of the genus Thiomargarita, which also exhibit a large internal vacuole
implicated in nitrate storage (Mufimann et al., 2003; Teske et al., 1996, 1999). Given the
similarity with Beggiatoa and the fact that, under certain circumstances, Thioploca
trichomes leave their sheaths and are then morphologically indistinguishable from
Beggiatoa filaments, it is likely that the classification of these two very similar genera
will be reconciled as more phylogenetic and physiological information accumulates.
Extensive mats of Thioploca cells have been discovered worldwide in both marine and
freshwater environments (Teske & Nelson, 2004) and Jorgensen & Gallardo (1999)
have suggested that the discontinuous occurrence of Thioploca mats that extend for a
distance of over 3000 km along the west coast of South America probably represent the
largest communities of sulfur bacteria on the planet.

Mats of Thioploca cells differ in structure from those composed of Beggiatoa cells.
This is principally a consequence of the presence of multiple Thioploca trichomes
within an extensive sheath that can penetrate several centimetres below the sediment
surface (Schulz et al., 1996). This innovation, coupled with the intracellular storage
of large amounts of nitrate, has suggested a novel ecological strategy whereby Thio-
ploca species exploit pools of oxidant and reductant that are separated in space.
Thioploca trichomes exhibit a tactic response towards nitrate and, when nitrate
concentrations are high in the overlying water, the trichomes extend into the water
column, where it is believed that they accumulate nitrate within their extensive system
of vacuoles. They then retreat back into their sheaths, carrying their store of oxidant to
greater depth in the sediment, where electron donor in the form of sulfide is freely
available either from sulfate reduction or geothermal sources (Huettel et al., 1996).

SGM symposium 65

40 N. D. Gray and I. M. Head

Fig. 3. Mats and streamers of filamentous sulfur bacteria from springs on Sulfur Mountain, CA, USA.

Thiothrix

Conspicuous streamers of Thiothrix cells are often noticeable in flowing waters from
sulfur springs (Fig. 3). In this respect, they differ from the sulfur bacteria considered
above. Beggiatoa species, in particular, rely on molecular diffusion through a stagnant
boundary layer to provide electron donor and acceptor to the cells (Schulz & Jorgensen,
2001). By virtue of specialized holdfast cells, members of the genus Thiothrix can
maintain themselves even in vigorously flowing waters that would wash away
filamentous sulfur bacteria lacking this adaptation. The rapid fluid flow decreases the

SGM symposium 65

Ecology of giant sulfur bacteria 41

Fig. 4. Rosettes of Thiothrix sp. from a sulfur spring, Gilsland Spa, Northumberland, UK.

unmixed boundary layer, freeing Thiothrix cells from the limitations of molecular
diffusion. It has long been considered that defining features of members of the genus
Thiothrix are their ability to deposit intracellular sulfur and their growth in character-
istic rosette structures that emanate long filaments (Fig. 4; Larkin, 1989). This has,
however, been called into question by a number of observations. The genus Thiothrix
includes organisms with highly variable morphology, such as Thiothrix eikelboomii
and Thiothrix defluvii, that do not necessarily produce rosettes (Howarth et aL, 1999;
Unz & Head, 2005). Furthermore, there is evidence that Leucothrix species, which also

SGM symposium 65

42 N.D.Gray and I. M. Head

produce rosettes and were until recently considered as heterotrophs, may deposit intra-
cellular sulfur when growing lithoheterotrophically (Grabovich et al., 2002). In addi-
tion, recently characterized bacteria that are morphologically reminiscent of Thiothrix
species have been shown to be related more closely to vacuolated Beggiatoa and
Thioploca strains (Kalanetra et al., 2004). Spectacular accumulations of Thiothrix cells
have been reported from sulfur springs (Brigmon et al., 2003; Larkin & Strohl, 1983;
McGlannan & Makemson, 1990) and in a range of cave environments (Brigmon et al.,
1994; Engel et al., 2004), associated with methane seeps (Pimenov et al., 2000) and,
more recently, they have been reported as an important component of novel prokaryotic
communities with a string-of-pearls morphology (Moissl et al., 2002).

Although not related closely to the genera Beggiatoa or Thioploca, Thiothrix species
also belong to the ' ' Gammaproteobacteria' and this group is replete with morpho-
logically distinctive sulfur bacteria (Fig. 2).

Thiomargarita

Thiomargarita cells are enormous. The bacterium produces large, spherical cells, and
cells in excess of 700 ^m in diameter have been reported (Schulz et al., 1999). The large,
spherical cells reside in a mucus sheath that links the otherwise unconnected cells in a
chain. Despite the huge size of Thiomargarita cells, the bacterium was discovered only
recently (Schulz & Jorgensen, 2005; Schulz et al., 1999). This may reflect the habitat
occupied by the bacterium. Thiomargarita isolates were first reported from seafloor
sediments at a depth of approximately 100 m, off the coast of south-western Africa.
Like Thioploca-dom'mated sediments off the western seaboard of South America, the
overlying water has low oxygen tension and relatively high nitrate concentrations,
which are a consequence of the very high productivity in these regions that are
characterized by upwelling of nutrient-rich deep waters. Like the large, marine
Beggiatoa and Thioploca species, Thiomargarita cells have large, central vacuoles that
constitute approximately 98 % of the biovolume of the cells, and accumulate nitrate to
levels of hundreds of millimolar (Schulz et al., 1999). Thiomargarita species do not
exhibit the gliding motility that is characteristic of Beggiatoa and Thioploca species
and rely instead on being transported passively into the water column by processes such
as gas venting, storms and wave action. This permits them to access their sources of
oxidant (oxygen and nitrate), which can be used to oxidize stored elemental sulfur or
sulfide when the cells are deposited back into the organic-rich sulfidic sediment that
they normally inhabit. The kinds of event that transport Thiomargarita cells passively
into the water column are innately sporadic and this may explain their huge size, which
allows them to store large quantities of oxidant that may be required for long-term
survival between resuspension events. In addition, they are tolerant of high oxygen
concentrations, whereas Beggiatoa and Thioploca species are microaerophilic and

SGM symposium 65

Ecology of giant sulfur bacteria 43

have a phobic response to high oxygen concentrations (Huettel et al., 1996). Empirical
observations suggest that Thiomargarita cells may survive in the absence of nitrate for
over 700 days (Schulz & Jorgensen, 2001).

This large-vacuolated bacterium is phylogenetically related most closely to the vacuo-
lated Beggiatoa and Thioploca species and, based on comparative 16S rRNA gene
sequence analysis, forms a monophyletic group with them in the 'Gammaproteo-
bacteria' (Fig. 2; Schulz etal., 1999).

Thiovulum

In 1913, Hinze described two species of Thiovulum (Hinze, 1913; Starr & Skerman,
1965): Thiovulum majus (the type species) and 'Thiovulum minus'. These were identi-
fied in marine environments and all isolated Thiovulum species to date come from
marine habitats. Although perhaps not in the same realm as the microbial giants
described above [Thiovulum cells range from 5 to 25 \im in diameter; Fenchel, 1994),
conspicuous 'veils' of Thiovulum cells are produced, associated with gradients of
sulfide and oxygen (Jorgensen & Revsbech, 1983). Perhaps their most remarkable
feature is their rapid motility, which has been measured at over 600 ^m (24-60 cell
lengths) s . This is equivalent to a 2 m fish swimming at up to >400 km h _1 (60 body
lengths s _1 ). Their rapid swimming is facilitated by extensive flagellation over the
surface of the cells and they use this either to propel the multicellular floating veils that
they form into the optimum position in opposed gradients of sulfide and oxygen
(Jorgensen & Revsbech, 1983) or to direct water flow around tethered veils to reduce
diffusional limitation of substrate transport to the cells (Fenchel & Glud, 1998). It is
now clear that a number of other bacteria have adopted similar strategies to enhance
transport of substrates to the cells (Thar & Kuril, 2002).

Apart from aspects of their chemosensory behaviour, motility and ultrastructure, there
are few studies on the physiology of Thiovulum species and only one partial 16S rRNA
gene sequence from authenticated Thiovulum cells has been deposited in the public
databases. The one Thiovulum species that has been characterized phylogenetically
to date belongs to the 'Epsilonproteobacteria' (Fig. 2; Lane et al., 1992; Romaniuk et
al., 1987), although it seems likely that, with more extensive sampling, the picture of
the taxonomy of the genus Thiovulum will change.

GIANT SULFUR BACTERIA: ENERGY AND GEOCHEMICAL
SIGNIFICANCE

Bacteria involved in the oxidative side of the sulfur cycle couple the oxidation of
reduced sulfur species to the reduction of dissolved oxygen, nitrate and possibly
oxidized metals. Many generate energy from these processes and reduce C0 2 for

SGM symposium 65

44 N.D.Gray and I. M. Head

biosynthesis. Some also have the capacity to fix dinitrogen gas and Beggiatoa species
have been best studied in this respect (Nelson et ai, 1982; Polman & Larkin, 1988).
Sulfur bacteria therefore play a pivotal role in the biogeochemical cycling of sulfur,
carbon, nitrogen and possibly metals. Their activities also affect the precipitation and
dissolution of minerals and can have consequences for higher trophic levels in a variety
of ecosystems.

The reduced sulfur that provides the energy to drive the biogeochemical activities of
sulfur bacteria can come from a variety of sources. Reduced sulfur can be liberated
from decaying organic matter, be produced from dissimilatory sulfate reduction or
released as a result of volcanic activity or from other geological sources. Environments
characterized by different sources of reduced sulfur present different geochemical
constraints on electron-donor and -acceptor availability, which have led to the numer-
ous ecological adaptations observed in sulfur-oxidizing bacteria. One major ecological
adaptation has been the evolution of sulfur bacteria with a large cell size.

The size of giant sulfur bacteria not only represents a remarkable evolutionary adap-
tation to the lifestyles that they have adopted, but also permits them to be manipulated
experimentally under in situ conditions and used as models to establish links between
environment, physiology and evolutionary diversification that may be extrapolated
more widely Some of the adaptations in the energetics and biochemistry of giant sulfur
bacteria have consequences for ecosystem-scale processes. These are discussed in
the context of the geochemical features of habitats dominated by conspicuous
sulfur bacteria.

Energetic considerations of sulfur oxidation

The transformation of S"~to SOj - by bacteria requires the transfer of eight electrons to
an appropriate electron acceptor. What is known or assumed about electron transport
in giant sulfur bacteria has been gleaned both directly (Cannon et ai, 1979; Schmidt &
DiSpirito, 1990; Strohl et ai, 1986) and by analogy with mechanisms observed in other
sulfur bacteria. What therefore follows, by necessity, does not represent a compre-
hensive picture of sulfur-oxidation pathways, but focuses on those areas where direct or
circumstantial evidence has been obtained from studies of giant sulfur bacteria.
Readers seeking a comprehensive review of sulfur-oxidation pathways are directed to
some excellent reviews (Friedrich et ai, 2001; Kappler &C Dahl, 2001; Kelly, 1999; Kelly
et ai, 1997). What these studies indicate is that there is no universal pathway of sulfur
oxidation and the giant sulfur bacteria, like all sulfur oxidizers, display considerable
variation in pathways of dissimilatory sulfur metabolism, even between closely related
species. For instance, the oxidation of hydrogen sulfide to sulfur, the first step in sulfur
oxidation, is linked either to the reduction of a c-type cytochrome, mediated by

SGM symposium 65

Ecology of giant sulfur bacteria 45

flavocytochrome c-sulfide dehydrogenase, or reduction of a quinone, mediated by
sulfide-quinone reductase, in phototrophic and lithotrophic bacteria (Friedrich et al.,
2001). However, very little information is available on how these processes are mediated
in the giant sulfur bacteria. Some studies have shown that the conversion of sulfide to
sulfur in freshwater Beggiatoa alba cells was not prevented by dibromothymoquinone,
an inhibitor of ubiquinone reduction, whereas inhibitors of flavoproteins (e.g. thenoyl-
trifluoroacetone) did suppress sulfide oxidation (Schmidt et al., 1987). This suggests
that electrons from sulfide enter the electron-transport chain via a flavocytochrome in
Beggiatoa alba (Schmidt et al., 1987). How typical this is of other giant sulfur bacteria is
not clear and Beggiatoa alba is not considered to be capable of chemolithotrophic
growth. This first step in the oxidation process is critical, because the sulfur produced
can then be either oxidized further or stored by precipitation as intracellular inclusions.
Although a characteristic of all giant sulfur bacteria is the deposition of intracellular
sulfur, a central problem still to be resolved in sulfur oxidation by chemolithotrophs is
the means by which zero-valent sulfur is transformed to SO'f and whether energy is
conserved from this step (Friedrich et al., 2001; Kelly et al., 1997). It has been proposed
that the process may involve oxidation of S° to sulfite by a sulfur oxygenase; sulfite
generated may then react with elemental sulfur to generate thiosulfate, which is
observed as a product in vitro (Kelly, 1999). The requirement for molecular oxygen by a
sulfur oxygenase means that this mechanism would not allow energy conservation and
would not be feasible during anaerobic sulfur oxidation by nitrate-reducing Tbioploca,
Beggiatoa or Tbiomargarita species, for example. Electrons from Tbiobacillus denitrifi-
cans growing with nitrate as an electron acceptor in the absence of oxygen have been
shown to be linked to a respiratory chain (Kelly, 1999) and there is evidence that, even in
aerobic sulfur oxidizers, energy is conserved during elemental sulfur oxidation and does
not involve a non-energy-conserving oxygenase. It is clear that several mechanisms may
be involved in elemental sulfur oxidation and the details may be different in different
bacterial taxa (Kelly, 1999). The final step in the sulfide-oxidation process is the
transformation of sulfite to sulfate and it is this section of the sulfur-oxidation pathway
for which most data are available in the giant sulfur bacteria.

Sulfite oxidation to sulfate is known to occur by three different pathways: one direct
and two indirect (Fig. 5; Kappler & Dahl, 2001). The first of the indirect pathways is in
essence a reversal of sulfate reduction: sulfite is oxidized to sulfate via the formation of
adenosine 5'-phosphosulfate (APS), mediated by APS reductase, and subsequent
decomposition by reaction with pyrophosphate to form ATP and sulfate (catalysed
by ATP sulfurylase; Fig. 5, pathway Ha). Alternatively, APS and orthophosphate may
be converted to ADP by ADP sulfurylase (adenylylsulfate : phosphate andenylyl-
transf erase), with the ADP being converted to ATP by adenylate kinase (Fig. 5, pathway
lib). Both indirect pathways therefore produce a molecule of ATP by substrate-level

SGM symposium 65

46 N. D. Gray and I. M. Head

Sulfide

Sulfur

4e-

03
Q.

Sulfite : acceptor

oxidoreductase

2e-

Sulfite

AMP ||

y/ APS re

reductase

PR
ATP

sulfurylase
ATP

APS

Sulfate

lib

Adenylylsulfate:

phosphate
ad enylyltransf erase

Adenylate A ^-ADP
kinase y
ATP + AMP

Fig. 5. Some of the alternative pathways of sulfur oxidation thought to be predominant in giant
sulfur bacteria. The pathways for oxidation of sulfide through sulfur to sulfite and the involvement of
intermediates, such as thiosulfate and tetrathionate, have not been elucidated in giant sulfur bacteria
and a minimal representation of oxidation of sulfide to sulfite is provided. This shows the number of
electrons generated at each step with potential to contribute to oxidative phosphorylation, but does
not represent the full range of reactions that may contribute to sulfide and sulfur oxidation.

phosphorylation. The direct mechanism of sulfite oxidation involves the enzyme
sulfite : acceptor oxidoreductase, which oxidizes sulfite to sulfate without the need for
AMP (Fig. 5, pathway I). This pathway only generates ATP by oxidative phosphory-
lation. Many sulfur bacteria carry both enzyme systems. However, whilst sulfite :
acceptor oxidoreductase is widespread among sulfur-oxidizing lithoautotrophs, the
APS pathway is less prevalent (Kappler & Dahl, 2001). Nevertheless, some Beggiatoa
strains exhibit both a direct and an indirect pathway (Hagen & Nelson, 1997) and the
APS reductase pathway occurs in sulfur-oxidizing symbionts of invertebrates and APS
reductase genes have been detected in Achromatium species (Head et al., 2000b; Nelson
& Hagen, 1995). In contrast, some Beggiatoa isolates do not exhibit APS reductase
activity (Grabovich et al., 1998, 2001; Hagen 6c Nelson, 1997).

SGM symposium 65

Ecology of giant sulfur bacteria 47

However, there is potential for substrate-level phosphorylation in some sulfur-oxidizing
pathways and, indeed, in some Thiothrix species, substrate-level phosphorylation is
believed to be the sole mechanism of energy conservation (Grabovich et ai, 1999;
Odintsova & Dubinina, 1993).

The thermodynamics and, hence, the energy yield of these oxidative processes are
dependent upon the particular electron acceptor used, the chemistry of the reduced
sulfur compound being oxidized and the reduced and oxidized products of the reaction.
Thermodynamic and kinetic considerations of sulfur oxidation have considerable bear-
ing on the ecology and physiology of the sulfur bacteria. For instance, the energy yield
from the oxidation of 1 mol HS - by 2 mol 2 to sulfate ostensibly yields -732-6 kj
(mol sulfur) -1 (Kelly, 1999). Whilst this may be important for giant sulfur bacteria,
many that are adapted to low oxygen concentrations may use nitrate as an alternative
electron acceptor. There is good evidence that the large-vacuolated Beggiatoa and
Thioploca species reduce nitrate to ammonium rather than dinitrogen gas, which only
accounted for approximately 15 % of the nitrate reduced (Otte et ai, 1999), and there
has been some debate regarding the validity of some earlier suggestions that Beggiatoa
species are capable of denitrification (Fossing et ai, 1995; McHatton et al., 1996;
Sweerts et al., 1990). Based on the free energy of formation values used by Kelly (1999),
it is possible to calculate the energy yield from the oxidation of sulfide coupled to
nitrate reduction to N 2 according to the equation

5HS-+8N03+3H + ^5SOf +4N 2 +4H 2

This yields approximately the same amount of energy [-744-76 kj (mol sulfur) -1 ] as
sulfide oxidation with oxygen [-732-6 kj (mol sulfur) -1 ]. Note that the energy yield
from reduction of nitrate to dinitrogen gas or ammonium is considerably greater than
that from reduction of nitrate to nitrite. It should also be noted that these values were
calculated under standard conditions and are useful for comparative purposes, but
probably do not reflect the energy yield under in situ conditions.

In contrast, the reaction

HS - +N03 + H + +H 2 O^NHj+SOj -

yields -463-8 kj (mol sulfur) -1 . On the face of it, one might conclude that these free-
energy values make the dissimilatory reduction of nitrate to ammonium (nitrate
ammonification or DNRA) less favourable, but several giant sulfur bacteria adopt this
way of life and reduction of NOT to NH4 is quantitatively important in many organic-
rich coastal marine sediments (Sorensen, 1978). In these sediments, whilst reduction of
nitrate to N 2 and consumption of 2 are restricted to the upper few millimetres
of sediment, the reduction of NO3 to NH|is significant at greater depth. The enzymol-
ogy of the two pathways of nitrate reduction is, however, quite different, as shown by

SGM symposium 65

48 N. D. Gray and I. M. Head

acetylene treatment of sediment microcosms. Acetylene prevents denitrification, but
DNRA is unaffected (Bonin et ai, 1998). Interestingly, the relative balance of these two
pathways has a significant effect on the sedimentary nitrogen cycle. This is because
denitrification produces nitrogen gas, which is biologically and chemically inert; in
contrast, DNRA produces ammonium that, under suitable conditions, can be reoxi-
dized to nitrate. Thus, biologically available nitrogen is more likely to be retained in
sediments where the majority of nitrate is reduced to ammonium (Sorensen, 1978).
Whilst this is a useful outcome with respect to the bacteria that utilize nitrate, it does
not in itself explain why, given the higher energy yield, nitrate is not always reduced
preferentially to N ? . A number of studies have indicated that, in high-sulfide environ-
ments, DNRA coupled to sulfide oxidation is favoured because of the inhibition
of denitrification at high sulfide concentrations (e.g. Brunet & Garcia-Gil, 1996).
However, it is also noteworthy that, in reduced environments, because of the stoichio-
metry of the two reactions, DNRA is probably no less thermodynamically favourable
than denitrification. In reducing marine sediments, sulfide is unlikely to be limiting,
whereas nitrate is likely to be less abundant. Only 1 mol nitrate is required to oxidize
1 mol sulfide to sulfate when nitrate is reduced to ammonium (an eight-electron
transfer), whereas 1-6 mol nitrate is required to oxidize 1 mol sulfide to sulfate when the
nitrate is reduced to N 7 (a five-electron transfer). Despite a more favourable energy yield
for denitrification than for DNRA when the values are expressed as kj (mol sulfur
oxidized) -1 , the energy yield in kj (mol nitrate reduced) -1 is broadly the same for the
two reactions, i.e. -465-5 kj (mol nitrate) -1 for denitrification and -463-8 kj (mol
nitrate) -1 for DNRA. From this different perspective, if nitrate is limiting compared
with the availability of electron donor, it would appear that the two processes are
equally energetically favourable.

However, the energy yield of electron transport is not dictated solely by thermo-
dynamics. The biology of the system may also have significance regarding the ability to
realize the maximal thermodynamic yield. To generate ATP by oxidative phosphory-
lation, it is necessary to generate a proton-motive force. This depends on the ability of
each step in the electron-transport pathway to export protons efficiently across the cell
membrane. The efficiency is dependent on the mode of respiration employed. The
aerobic electron-transport chain, because of its topography, is more efficient at
converting the chemical energy of electron donors into ATP (Berks et al., 1995). This is
because cytochrome aa 3 is not only a proton pump, but it also delivers electrons to the
cytoplasmic side of the membrane, where they combine with oxygen and protons and
hence create a greater electrochemical gradient. In contrast, in denitrification, at least in
some organisms, the enzymes nitrite reductase, nitric oxide reductase and nitrous oxide
reductase are located on the periplasmic side of the membrane, the consumption of
protons occurs externally and the net gain in translocated protons is lower than might

SGM symposium 65

Ecology of giant sulfur bacteria 49

be suggested solely by the reduction potentials of the respective electron acceptors
(Berks et al., 1995). Thus, despite similar thermodynamics, different respiratory pro-
cesses can have very different energetic efficiency in terms of ATP formation. In the
process of DNRA, after the reduction of nitrate to nitrite, nitrite is reduced to
ammonium without the release of intermediates (Simon, 2002). In this process, nitrite is
reduced by a cytochrome c-nitrite reductase complex, which mediates the multi-
electron and multi-proton reduction of nitrite to ammonium. It is clear that this
process, when coupled to sulfide oxidation, produces ATP (Simon, 2002). In the small
number of cases studied to date, nitrite reduction to ammonium occurs on the outside
of the cell membrane where protons are consumed, reducing the effective proton-
motive force as in denitrification (Simon, 2002).

Geochemical significance of giant sulfur bacteria

The metabolic diversity of giant sulfur bacteria and the huge abundances that they can
achieve in appropriate environments mean that they play a significant role in a range of
biogeochemical processes. A further consequence of the activity of sulfur-oxidizing
bacteria is that they may have pronounced effects at higher trophic levels. The various
giant sulfur bacteria oxidize reduced sulfur species, deposit elemental sulfur, fix
inorganic carbon, utilize organic compounds as electron donors or carbon sources,
consume oxygen and reduce nitrate and nitrite to ammonium and perhaps, to a lesser
extent, gaseous products. Several are capable of fixing dinitrogen gas and some may
mobilize insoluble metals and other minerals. Giant sulfur bacteria can therefore
have important implications for the cycling of many elements.

Giant sulfur bacteria and the sulfur cycle. The oxidation of sulfide to elemental
sulfur and ultimately sulfate by giant sulfur bacteria clearly has considerable
consequences for sulfur cycling, especially where extensive mats of bacteria occur. The
sediments off the Pacific coast of South America, which harbour extensive mats of
Thioploca cells, exhibit high rates of sulfate reduction (up to 1500 nmol cm -3 day -1 );
however, sulfide rarely accumulates above a few tens of micromolar (Ferdelman et al.,
1997; Thamdrup 6c Canfield, 1996). The low concentrations of sulfide can be explained
to some extent by reaction of sulfide with reactive iron, which is considered an impor-
tant sink for sulfide in the Pacific coast sediments (Thamdrup 6c Canfield, 1996).
Nevertheless, it has been estimated that anywhere between 16 and 91 % of sulfide
reoxidation can be attributed to mats of Thioploca cells (Ferdelman et al., 1997; Otte
et al., 1999). This has the dual role of replenishing an important oxidant for anaerobic
carbon mineralization (sulfate) and reducing the steady-state concentration of toxic
sulfide. Sediments off the Namibian coast also exhibit low levels of dissolved sulfide
(Briichert et al., 2003). This is particularly unexpected, because of the relatively low
levels of reactive iron species in Namibian coastal sediments (Briichert et al., 2000;

SGM symposium 65

50 N. D. Gray and I. M. Head

Morse 6c Emeis, 1990). The effective removal of sulfide in this environment is therefore
due, in large part, to the effective reoxidation of sulfide by the mats of Tbiomargarita
cells. In this case, the giant sulfur bacteria could account for up to 55 % of the reoxi-
dation of sulfide and a positive correlation was found between the measured sulfide flux
from the sediment and the size of the population of Tbiomargarita cells (Brtichert et al.,
2003). Furthermore, considering that Tbiomargarita cells may contain up to 1-7 M
sulfur in their cytoplasm and that biomass densities as high as 176 g m -2 may occur off
the coast of south-west Africa (Brtichert et al., 2003), they represent a major repository
of elemental sulfur in these sediments. For example, the biovolume of Tbiomargarita
cells in the top few centimetres of Namibian coastal sediments was approximately
4 ^1 mm (Schulz et al., 1999). Given that only 2 % of this volume is cytoplasm, this
gives a value of 0-08 ,ul cytoplasm per cubic millimetre of sediment, which will contain
up to 4-35 \xg sulfur. Converting this to an areal basis, this equates to approximately
4-35 g sulfur m -2 contributed by the Tbiomargarita cells. Interestingly, measurements of
intracellular and extracellular elemental sulfur in Tbioploca-dom'mated sediments
off the coast of Chile showed that extracellular sulfur was abundant and had a different
depth distribution from intracellular sulfur, suggesting that chemical oxidation also
played an important role in sulfide oxidation in these sediments (Zopfi et al., 2001).
This is in agreement with the conclusions drawn based upon the relative availability of
reactive iron in Namibian and Chilean coastal sediments (Brtichert et al., 2003).

In freshwater sediments, the concentration of sulfate is orders of magnitude lower than
that in marine systems and it is rarely possible to measure dissolved sulfide in these
environments. Consequently, fewer studies of sulfur cycling have been conducted in
freshwater sediments. In many circumstances where there are extensive accumulations
of sulfur bacteria in freshwater environments, this may be due to locally high levels of
sulfide associated with decaying organic matter, rather than from sulfate reduction.
Nevertheless, even when no dissolved sulfide is detectable, large populations of giant
sulfur bacteria may be supported (Head et al., 1996, 2000b). These may have greater
significance for biogeochemical cycling in freshwater systems than the sulfate concen-
tration would suggest. Achromatium species are extremely effective at reoxidizing
reduced sulfur generated by sulfate reduction (Gray et al., 1997). This is the case even in
environments with extremely high levels of reactive iron, which would be expected to
effectively outcompete Achromatium species for the low levels of sulfide produced by
sulfate reduction (Head et al., 2000b). This not only has consequences for the sulfur
bacteria themselves, but also suggests that a greater fraction of organic carbon mineral-
ization in Achromatium-btznng sediments is channelled through sulfate reduction
than would be predicted from measurements of sulfate reduction in incubations
performed under anoxic conditions (Gray et al., 1997; Head et al., 2000a, b). It has been
suggested that, for sulfur bacteria to thrive under conditions of low sulfide and high

SGM symposium 65

Ecology of giant sulfur bacteria 51

reactive-iron concentrations, they must compete effectively with the rapid chemical
reaction between sulfide and iron (Thamdrup 6c Canfield, 1996). In the case of Thio-
ploca cells in marine sediments, this may be achieved through close physical association
with the sulfate-reducing bacteria that generate the sulfide (Thamdrup 6c Canfield,
1996). However, this does not seem to be the case for Achromatium species, where other
bacterial cells have not been found associated with Achromatium cells. Achromatium
may therefore possess high-affinity sulfide-uptake systems that are capable of compet-
ing effectively with chemical reactions with iron to permit the utilization of low levels
of dissolved sulfide. The kinetics of sulfide oxidation by Achromatium species have yet
to be studied in detail.

On the other hand, Achromatium cells may not use dissolved sulfide directly, but
may instead use solid-phase iron sulfides as their principal source of electron donor
(Gray et al., 1997; Head et al., 2000a, b). If this is the case, the precipitation of intra-
cellular calcium carbonate by these bacteria may provide a means by which they could
dissolve solid-phase iron sulfides and thus drive reoxidation of sulfur at greater rates
than would be predicted from very low free-sulfide concentrations. Several explanations
for calcite deposition in Achromatium species, i.e. neutralization of acid formed by
sulfur oxidation, regulation of cell buoyancy or maintenance of high C0 2 partial
pressures to facilitate autotrophic growth, have been proposed (Head et al., 2000b).
However, this physiological feature, which is unique to Achromatium species in the
bacterial domain, suggests that it represents an adaptation to an ecophysiological
challenge not faced by other organisms.

Calcite precipitation linked to sulfide dissolution may therefore be an ability unique to
members of the genus Achromatium. Protons exported during chemiosmotic energy
generation are consumed in the mineral dissolution reaction

FeS + H + ^Fe 2+ + HS"

and the sulfide generated can then be oxidized to sulfate:

HS-+20 2 -HS04

As some of the exported protons will be consumed by their reaction with metal
sulfides, they will be unavailable for transport across the cell membrane by ATP
synthase. An inevitable consequence is that a surplus of hydroxyl ions will occur on the
cytoplasmic side of the cell membrane. This is exacerbated if nitrate is used as an
electron acceptor, as sulfide oxidation with nitrate consumes protons:

HS-+N03 + H + +H 2 0-*NH++SOj-

5HS-+8N03+3H + -5SOf +4N 2 + H 2

SGM symposium 65

52 N.D.Gray and I. M. Head

Mineral surface

FeS

Periplasm
Ca 2+

HCO^

H + «-

Fe 2+ + HS

Cytoplasm

H 2 + COJ

+ HC0 3 + OH

H 2

HS" + 20-

hso;

Calcite inclusion

-► Ca 2+

CaCO,

Fig. 6. Putative mechanism for the coupling of intracellular calcite precipitation and sulfide mineral
dissolution and oxidation in Achromatium species.

A potential sink for excess hydroxyl ions is through buffering with bicarbonate and the
formation and precipitation of calcite:

OH-+HCOj-*CO^ + H 2

Ca 2+ +C05--CaC0 3 (s)

The coccolithophorid algae have previously been proposed to buffer increases in
cytoplasmic pH via intracellular precipitation of calcite (Borowitzka, 1982). These
photosynthetic organisms deposit intracellular calcite to maintain an high internal
partial pressure of C0 2 to facilitate carbon fixation by ribulose-l,5-bisphosphate
carboxylase/oxygenase (RuBisCO) (Borowitzka, 1982). Under conditions of low
dissolved C0 7 and high dissolved oxygen, bicarbonate is converted to C0 1 and
hydroxyl ions by carbonic anhydrase (Borowitzka, 1982). The subsequent fixation of
C0 ? , however, raises cytoplasmic pH, due to the residual hydroxyl ions. To neutralize
the cytoplasmic pH, a second molecule of bicarbonate is precipitated as calcium
carbonate within an intracellular membrane-bound structure, the coccolith-containing
vesicle (Fig. 6). This reaction consumes hydroxyl ions with the generation of carbonate
and water:

SGM symposium 65

Ecology of giant sulfur bacteria 53

OH-+HCOj-*CO^ + H 2

Ca 2+ +CO|-^CaC0 3 (s

An alternative mechanism by which Achromatium species might maintain a neutral pH
in the cytoplasm would be to liberate protons directly from the reaction of bicarbonate
ions with calcium. This is unlikely, as calcite is precipitated not at the cell membrane,
but within membrane-bound intracellular inclusions (Fig. 6).

Ca 2+ +HC03-CaC0 3 + H +

Despite the fact that this mechanism of carbonate precipitation is encountered widely
in the literature, the direct precipitation of calcium carbonate in this manner does not
occur in nature. This is because calcite is ionic and can only be formed from its
component ions. As a result, CO^must be formed before calcite formation can take
place (Wright & Oren, 2005).

This proposed basis for calcite precipitation in members of the genus Achromatium
means that there should be a direct link between calcite precipitation and energy
generation. For this reason, the energetics of the whole process need to be evaluated.
Interestingly, the stoichiometry for iron sulfide oxidation linked to calcium carbonate
precipitation shows that there is no net gain or loss of protons and, thus, the process
should be chemiosmotically neutral:

FeS + Ca 2+ +20 2 +HC03-HS04+CaC0 3 +Fe 2+

However, there are other factors that should be considered: for instance, Achromatium
cells contain large amounts of calcite, even when calcium is below detection limits in its
environment. This implies a massive accumulation of calcium against a considerable
concentration gradient. There is thus likely to be an energetic cost of precipitating
calcite that must be balanced against the enhanced energy generation supported by the
consumption of intracellular hydroxyl ions. In environments with high levels of
dissolved sulfide, it is unlikely that this mechanism to access solid-phase sulfides would
be competitive, because of the cost of calcium carbonate precipitation. Interestingly,
the only Achromatium species reported routinely from sulfidic marine systems is
'Achromatium vo\utans\ which does not precipitate calcite.

Giant sulfur bacteria and the nitrogen cycle. Some giant sulfur bacteria have an
important role to play in the biogeochemical cycling of nitrogen. There is evidence that
some Beggiatoa species are capable of fixing dinitrogen gas (Polman 6c Larkin, 1988),
but more attention has been diverted toward the ability of a number of giant sulfur
bacteria to use nitrate as a respiratory oxidant. Initially, this was presumed to be linked
to denitrification (Fossing et ai, 1995; Sweerts et ai, 1990), but more recent evidence

SGM symposium 65

54 N.D.Gray and I. M. Head

strongly supports the notion that the principal route for respiratory nitrate reduction is
through DNRA (Otte et al., 1999). This has prompted a considerable reappraisal of the
significance of mats of giant sulfur bacteria in the nitrogen dynamics of sedimentary
environments.

In many coastal marine environments, nitrogen is the most important limiting nutrient
(Nixon, 1981; Ryther 6c Dunstan, 1971) and removal of nitrogen via denitrification can
be an important control on eutrophication (Seitzinger, 1988).

As organic carbon inputs to sediments increase in marine environments in particular,
denitrification becomes less important and DNRA more so. This has been attributed to
increased sulfate reduction and the accumulation of sulfide to concentrations that
inhibit key enzymes in the denitrification pathway (Brunet &C Garcia-Gil, 1996), but, as
discussed above, energetic considerations may also play a role. Nevertheless, it is clear
that a shift in nitrogen-reduction pathways from denitrification to a conservative route
of nitrate reduction, such as DNRA, will have important consequences for the overall
budget of nitrogen and the productivity of a system. In this context, the ability of giant
sulfur bacteria, such as members of the genera Thioploca and Beggiatoa, which
accumulate high levels of nitrate, to reduce nitrate to ammonium coupled with sulfide
oxidation (McHatton et al., 1996; Otte etal., 1999) clearly has significance.

Studies of sediments off the coast of Chile have shown that nitrate fluxes into sediments
with large populations of Thioploca cells are >50 % higher than in sediments with a
low abundance of Thioploca cells (Zopfi et al., 2001). It has been suggested that,
because Thioploca cells can extend into nitrate-containing waters overlying the
sediments that they inhabit, they may monopolize the available nitrate before it can
be transported to other sediment-dwelling bacteria. This would imply that other forms
of nitrate respiration might be of minimal significance in nitrogen cycling in sediments
harbouring large populations of Thioploca cells. Experimental evidence refutes
this suggestion and denitrification rates in Thioploca-rich sediments were actually
slightly higher (4-5-9 mmolrrT 2 day -1 ) than typical values for coastal sediments
(Herbert, 1999). Thus, there was still sufficient nitrate reaching the sediment to support
significant denitrification rates, despite considerable assimilation of water-column
nitrate by Thioploca cells. Furthermore, nitrous oxide was also consumed rapidly,
suggesting that complete denitrification occurred in the sediments (Zopfi et al., 2001).
Although the sole route of nitrate reduction in these sediments is not through DNRA
linked to sulfide oxidation and not all sulfide oxidation is mediated by Thioploca
species, it has been estimated that if 25 % of sulfide oxidation is mediated by nitrate-
reducing Thioploca cells [toward the lower range of reported estimates (16-91 %)], this
would result in an increase in the net flux of ammonium from the sediments of S3 %.

SGM symposium 65

Ecology of giant sulfur bacteria 55

This represents a significant increase in the amount of nitrogen retained in the system
and is likely to have a significant effect on the overall primary productivity in such
coastal ecosystems, which are typically nitrogen-limited. Indeed, although Graco et al.
(2001) found that the main source of ammonium in Chilean coastal sediments was in
fact mineralization of organic matter and only 17 % of the total recycled ammonium
resulted from the DNRA activity of giant sulfur bacteria, the activity of the sulfur
bacteria still made a significant contribution to the recycled nitrogen and could account
for >10 % of the primary productivity supported by recycled nitrogen. Similar con-
clusions have been drawn from studies of sediments in Tokyo Bay, which have also
found that nitrate-accumulating, vacuolated sulfur bacteria have an important role in
nitrogen retention in this coastal ecosystem. Interestingly, conditions of high
productivity are likely to exacerbate the situation, as higher inputs of organic matter
stimulate sulfate reduction, leading to higher sulfide concentrations that may repress
denitriflcation and favour DNRA (Brunet & Garcia-Gil, 1996), resulting in greater
nitrogen retention in the system with potential effects on long-term eutrophication. It is
interesting to speculate about the wider implications of nitrogen retention on the bio-
geochemical cycling of sulfur and carbon. Effective recycling of nitrogen within
the system will ultimately lead to greater sulfate regeneration and, consequently, the
importance of carbon mineralization through sulfate reduction may also be increased.
This would be further enhanced by the greater stoichiometric efficiency of DNRA
relative to denitriflcation.

Giant sulfur bacteria and the carbon cycle. The vast accumulations of biomass in
mats of giant sulfur-oxidizing bacteria represent a significant pool of organic carbon. It
has been shown that sedimentary carbon dynamics are affected by the presence of mats
of Thioploca and Beggiatoa cells, and sediments characterized by sulfur-oxidizing
bacterial mats have been shown to promote higher rates of carbon burial (Graco et al.,
2001). It has been suggested that this may be a consequence of sheaths and cells of
Thioploca and Beggiatoa containing refractory carbon and limited grazing on the giant
bacterial cells, although no direct measurements of the resistance of sulfur bacterial
sheaths have been made (Graco et al., 2001).

Most studies of the effect of sulfur bacterial mats on carbon cycling have focused on
cold-seep environments, where methane and hydrocarbons are important carbon
sources driving the microbial ecosystem. Mats of different colour have often been
reported from such sites and recent studies have shown that the colour of the mats,
which is believed to reflect high cytochrome content of the orange-mat bacteria,
correlates with the type of carbon metabolism exhibited by the sulfur bacteria
(Nikolaus et al., 2003). However, in this case, it was indicated that the water-soluble
pigments, which had an absorbance maximum at 390 nm, were unlike typical cyto-

SGM symposium 65

56 N. D. Gray and I. M. Head

chromes (Nikolaus et al., 2003). The non-pigmented mats expressed high but variable
levels of RuBisCO activity, suggesting that the non-pigmented sulfur bacteria had
chemoautotrophic potential. RuBisCO activity in the pigmented filaments was several
orders of magnitude lower, suggesting that they were more likely to be heterotrophic,
perhaps utilizing hydrocarbons or hydrocarbon-oxidation products prevalent at the
seep site (Nikolaus et al., 2003).

Sulfur bacterial mats are characteristic of many methane-rich environments and it has
even been suggested that the mats may form a relatively impermeable barrier that
restricts the movement of methane out of the sediment (Orphan et al., 2004). This may
help to promote anaerobic oxidation of methane by holding methane in surface
sediments, close to the large pool of sulfate in the overlying water and also to a source of
sulfate from reoxidation of sulfide by the sulfur bacteria themselves. Consequently, high
levels of anaerobic methane oxidation appear to be associated with conspicuous sulfur
bacterial mats (Joye et al., 2004; Treude et al., 2003). The high levels of anaerobic
methane oxidation, which are typically linked to sulfate reduction in marine systems,
appear to be supported by the effective reoxidation of sulfide by the bacterial mats (Joye
et al., 2004). Isotopic data, obtained from a combination of fluorescent in situ hybridi-
zation and secondary-ion mass spectrometry (FISH-SIMS; Orphan et al., 2001),
suggest that the filamentous mat bacteria also incorporate methane-derived light
carbon (Orphan et al., 2004). It is clear from these data and the precipitation of iso-
topically light carbonates in seep environments that much of the carbon from active
seep sites associated with sulfur bacterial mats is retained in the system by the
combined activities of the anaerobic methane-oxidizing archaeal-sulfate-reducer
consortia and mats of giant sulfur bacteria (Boetius & Suess, 2004). It is not only the
microbial component of the ecosystem that is affected by the carbon dynamics in seep
environments; higher trophic levels are also influenced by the presence of sulfur
bacterial mats. One study has shown that both the density and diversity of fauna are
greater at seep sites with extensive sulfur bacterial mats, compared with similar settings
that lack the bacterial mats. Mats of Tbioploca cells in the Gulf of Mexico in particular
appeared to support a more diverse foraminiferan community, compared with mats of
Beggiatoa cells and off-mat sites (Robinson et al., 2004).

EVOLUTIONARY AND ECOLOGICAL DIVERSITY IN THE GIANT
SULFUR BACTERIA

Giant sulfur bacteria have been shown by comparative 16S rRNA gene sequence
analysis to be distributed between two classes of the phylum Proteobacteria (Teske
et al., 1996). The majority (Beggiatoa-Thioploca-Thiomargarita, Achromatium and
Thiothrix) form well-defined clades within the l Gammaproteobacteria\ However, the
genus Thiovolum clusters with the 'Epsilonproteobacteria' (Fig. 2).

SGM symposium 65

Ecology of giant sulfur bacteria 57

The different genera of giant sulfur bacteria have each evolved major morphological
and physiological adaptations that allow them to exploit sulfur oxidation in different
geochemical settings. Within each of these genera, however, there is evidence for more
recent diversification and adaptation. It has been possible to correlate fine-scale
patterns of genetic diversity with ecological function in giant sulfur bacteria and, by
extrapolation, this is providing insights into the ecological mechanisms that underpin
diversification in the microbial world more generally.

Within-genus diversity in morphology and physiology

On the basis of comparative 16S rRNA gene sequence analysis, the genera Beggiatoa,
Thioploca and Tbiomargarita occupy a monophyletic group within the Proteobacteria.
Within this, four distinct lineages are evident (Teske & Nelson, 2004). The large-
vacuolated, marine Beggiatoa and Thioploca species, Tbiomargarita namibiensis and
a recently described, vacuolated, rosette-forming sulfur bacterium (Kalanetra et al.,
2004) occupy one group, whilst freshwater Thioploca species form a second group of
closely related organisms. The non-vacuolated, marine Beggiatoa species form the third
group, with the fourth comprising freshwater Beggiatoa isolates (Teske & Nelson,
2004). Until relatively recently, our knowledge of the diversity of these bacteria has
largely been defined by a few recognized species, identified on the basis of filament
diameter, the presence of nitrate-storing vacuoles and the presence or absence of sheath
material (Teske 6c Nelson, 2004). Some of these morphological features map onto
the 16S rRNA gene sequence-based phylogeny (e.g. filament diameter and presence of
vacuoles), whereas others (sheath formation) do not. As more 16S rRNA gene sequence
data become available for giant sulfur bacteria, it is becoming apparent that there
is considerably greater genetic diversity within each of the described genera than the
original, morphologically based descriptions would suggest. For instance, a recent
study of the diversity and population structure of filamentous sulfur bacteria in the
Danish Limfjorden and the German Wadden Sea showed high diversity in nitrate-
storing, vacuolated Beggiatoa species coexisting within the same sediments (Mufonann
et al., 2003). These organisms represented novel phylogenetic clusters distinct from
those characterized previously. It was predicted from the wide spectrum of filament
diameters encountered in these environments and the relatively restricted diameter of
individual Beggiatoa species identified by using FISH probes that the entire diversity
that was present within the genus Beggiatoa had not been surveyed completely.
Interestingly, the coexisting Beggiatoa species identified by FISH showed depth-related
differences in their distribution (MuEmann et al., 2003). This was attributed either to
the ability of species with a wider average filament diameter to store nitrate in greater
quantities and hence reside at greater depths for longer, or the ability of species with a
smaller average diameter to scavenge more efficiently for dissolved sulfide, which is in
limited supply in the near-surface environment. This is echoed by the findings that the

SGM symposium 65

58 N. D. Gray and I. M. Head

depth distribution of Thioploca chileae, Thioploca araucae and a third Thioploca
species, designated SCM (short-cell morphotype), in South American coastal sedi-
ments was different, and the nitrate and sulfur content of the different Thioploca
species correlated with the different location of the bacteria in the sediment (Zopfi
et al., 2001). As in most other locations, nitrate-accumulating Beggiatoa cells from
organic-rich sediments in Tokyo Bay also had varying widths of trichome (Kojima &
Fukui, 2003). The morphotype with wider trichomes was related most closely to
uncultured Beggiatoa species from other geographical localities, identified from 16S
rRNA gene sequence analysis, and distinct from the Tokyo Bay morphotype with
narrow trichomes. Among the narrower types, which formed a new branch within
the Beggiatoa-Thopmargarita-Thioploca clade, a sample of cells from a tidal flat
harboured a narrow Beggiatoa-like morphotype that was genetically distinct from
those identified in samples taken in water depths of 10 and 20 m. This suggests that
there may be some ecological differentiation between closely related Beggiatoa geno-
types. Analysis of 16S rRNA gene sequences from Chilean Thioploca isolates revealed
several coexisting, genetically distinct species that can be differentiated by filament
diameter. Not only were these found in the same sediment sample, but they also
occurred commonly within the same sheath (Teske et al., 1996). In contrast, a more
recent study of freshwater Thioploca isolates from Lake Biwa, Japan, and Lake
Constance, Germany, indicated that two freshwater Thioploca isolates from geo-
graphically separated environments were barely distinguishable on the basis of
trichome diameter and had almost identical 16S rRNA gene sequences (Kojima et al.,
2003).

As a relatively large number of Beggiatoa species have been isolated in axenic culture, a
great deal more is known about the physiology of members of the genus Beggiatoa than
other large sulfur bacteria. Members of the genus Beggiatoa have a diverse carbon
metabolism, ranging from obligate chemolithoautotrophy, facultative chemolitho-
autotrophy and mixotrophy through to lithoheterotrophy This phenotypic diversity is
observed at the species level, so it appears that carbon metabolism has been a key driver
of ecological and evolutionary diversification. For instance, although it appears that the
marine Beggiatoa species (vacuolated and unvacuolated) display predominantly
chemolithoauto trophic nutrition (Teske & Nelson, 2004), it also appears that they can
be either obligate or facultative autotrophs, as demonstrated by the related marine
strains MS-81-lc and MS-81-6 (Hagen & Nelson, 1996). For instance, in the presence of
acetate, carbon fixation by the obligately chemolithotrophic strain MS-81-lc was not
reduced significantly (Hagen & Nelson, 1996). In addition, the use of radiolabeled
substrates demonstrated that this organism was unable to respire acetate and that
2-oxoglutarate reductase, an enzyme necessary for the respiration of organic substrates
in the tricarboxylic cycle, was absent. Even in the presence of organic compounds, 80 %

SGM symposium 65

Ecology of giant sulfur bacteria 59

of cell carbon was obtained from C0 9 . Strain MS-81-lc also obtained more energy
from the oxidation of reduced sulfur species than the facultatively chemolitho-
autotrophic marine strain MS-81-6 (Hagen & Nelson, 1997). MS-81-6, although able
to grow chemolithoautotrophically, was found to be metabolically more versatile and
could utilize acetate freely for energy or biosynthesis (Hagen & Nelson, 1996).
Beggiatoa sp. strains MS-81-lc and MS-81-6 were originally isolated from the same
environment (Great Sippewissett salt marsh, Woods Hole, MA, USA) and the physio-
logical differences between the strains may be a factor in the maintenance of the
diversity of Beggiatoa species in this single geographical location.

In contrast to marine organisms, there has been considerable debate as to whether
genetically differentiated freshwater Beggiatoa strains are capable of genuine
autotrophic or lithoheterotrophic growth (Hagen & Nelson, 1997; Strohl, 2005; Teske
& Nelson, 2004). Whilst mixotrophic nutrition has been claimed for a large number
of freshwater strains (Hagen & Nelson, 1996), it has only recently been shown that a
freshwater isolate can grow lithoautotrophically (Grabovich et al., 2001). Nevertheless,
it is clear that the genetic differentiation between the marine and freshwater Beggiatoa
species is correlated with fundamental differences in carbon metabolism, given the
predominance of heterotrophic/mixotrophic metabolism in the freshwater strains and
autotrophic metabolism in the marine species. The ecological sense of these differences
may be explained by the availability of sulfide in marine and freshwater environments.
The metabolic versatility in the freshwater strains is consistent with the potentially low
and variable supply of sulfide in these environments and hence the requirement to
supplement growth by the utilization of organic carbon.

At present, very little is known about the genetic diversity of the genus Thiomargarita
and 16S rRNA gene sequence data are only available for Thiomargarita namibiensis
from Walvis Bay, Namibia, south-west Africa (Schulz et al., 1999). Unlike most com-
munities of Achromatium, Beggiatoa and Thioploca species, samples of Thiomargarita
namibiensis from Walvis Bay only produced a single 16S rRNA gene sequence, suggest-
ing either that Thiomargarita namibiensis is genetically much more homogeneous than
other giant sulfur bacteria or that any genetic differentiation of Thiomargarita species
cannot be resolved on the basis of 16S rRNA gene sequence analysis. It may be of
significance that the habitats where several distinct species of giant sulfur bacteria
coexist tend to be spatially structured, relatively undisturbed sediments. The different
genotypes observed in these systems may therefore reflect species adapted to different
conditions in the sediment environment (Gray et al., 1999b). Members of the genus
Thiomargarita are believed to have quite a different lifestyle from sulfur bacteria that
exhibit genetic diversity at a single location. It is thought that Thiomargarita cells
survive by living in an environment that is subject to periodic disturbance and mixing.

SGM symposium 65

60 N. D. Gray and I. M. Head

The mixing brings sessile cells in sulfide-rich sediment into contact with water
containing nitrate and oxygen, the oxidants that it uses for sulfide oxidation (Schulz &
Jorgensen, 2001). It is often considered that mixed environments sustain less diversity
than structured environments, as they provide a narrower range of environmental
conditions that can be exploited by ecologically distinct organisms. This may in part
explain the apparent homogeneity of communities of Thiomargarita cells. None-
theless, chains of Thiomargarita cells also fall into distinct cell-size classes that, by
analogy with other giant sulfur bacteria, may correlate with different species (Schulz
et al., 1999). Alternatively, the failure to identify different Thiomargarita species may
be due to limited sampling of the communities.

The genus Thiothrix forms a deep branch within the ' Gammaproteobacteria' ; however,
there is considerable physiological variation within the genus (Aruga et al., 2002;
Howarth et ai, 1999; Rossetti et ai, 2003). Thiothrix species, like Beggiatoa species, are
varied in their carbon metabolism and heterotrophs, mixotrophs and chemolitho-
autotrophs are all represented by different Thiothrix species. At present, no studies have
established how this genetic and physiological diversity relates to the distribution and
abundance of Thiothrix species in engineered or natural ecosystems. However, a
number of studies have shown that bacteria related closely to Thiothrix unzii, which
can grow with organic carbon sources but has an obligate requirement for sulfide or
thiosulfate (Unz &C Head, 2005), appear to be prevalent in some sulfidic cave environ-
ments (Brigmon et ai, 2003; En gel et al., 2004).

Genetic and ecological diversity in the genus Achromatium

The tantalizing link between fine-scale patterns of genetic diversity and ecological
differentiation in giant sulfur bacteria has been explored most thoroughly in the genus
Achromatium. Natural communities of Achromatium cells that were originally
thought to be genetically homogeneous in fact comprise a number of phylogenetically
distinct subpopulations, distinguishable by FISH (Glockner et al., 1999; Gray et ai,
1999b). The degree of identity observed in Achromatium-dtrivtd 16S rRNA gene
sequences (<97-5%) from individual samples of freshwater sediment indicates
that different species of Achromatium exist both in geographically separated locations
and within a single sediment (Gray et al., 1999b). As with the marine Beggiatoa
species described above, it has been shown that individual, coexisting Achrom-
atium species exhibit physiological and ecological differentiation. Coexisting
Achromatium species in sediment from the margins of Rydal Water in the English Lake
District, like Beggiatoa and Thioploca species, fall into distinct size classes (Gray et ai,
1999b). Furthermore, the composition of the Achromatium community was different
in oxidizing and reducing zones within the sediment (Gray et al., 1999b). On this basis,
it was hypothesized that genetic diversity in coexisting Achromatium communities

SGM symposium 65

Ecology of giant sulfur bacteria 61

reflected exploitation of different redox-related niches. This was based on the competi-
tive-exclusion principle, which states that if two similar species coexist in a stable
environment, they do so as a result of niche differentiation. If, however, there is no niche
differentiation, the species will compete and one will eliminate or exclude the other
(Begon et aL, 1996; Gause, 1934; Hardin, 1960). Redox-related niche differentiation
in Achromatium communities was demonstrated experimentally (Gray et aL, 2004). It
was reasoned that, in sediment microcosms harbouring Achromatium communities,
growth or maintenance of a particular Achromatium species exposed to particular
redox conditions, accompanied by neutral or antagonistic effects on other coexisting
Achromatium species, would be indicative of niche differentiation. In anoxic micro-
cosms, Achromatium sp. RY8 decreased and Achromatium spp. RY5 and RYKS
increased over time. Addition of increasing concentrations of nitrate, however, main-
tained the population of Achromatium sp. RY8. When high levels of nitrate were
maintained throughout the incubation, the composition of the Achromatium
community remained stable over time (Fig. 7). This suggested that all of the coexisting
Achromatium species are obligate or facultative anaerobes that can utilize nitrate as an
electron acceptor, but it was not possible to establish whether Achromatium species
utilized DNRA or denitrification. Achromatium sp. RY8 was clearly more sensitive to
sediment redox conditions than the other Achromatium species. These results give a
much clearer picture of the mechanism that supports coexistence of several
Achromatium species adapted to different redox conditions (Gray et aL, 2004). Redox
conditions in sediments vary with depth, and coexisting Achromatium species are
adapted to different redox conditions and thus avoid direct competition. Clearly, then,
the genetic diversity observed in Achromatium communities correlates to functional
diversity. This in turn may have an impact on geochemical cycling. The inherently low
levels of sulfate that typify sediments inhabited by freshwater Achromatium species
suggest that regeneration of sulfate is an important process and may lead to a greater
flow of electrons from organic carbon through sulfate reduction than would be
suggested by the steady-state sulfate concentration. Consequently, Achromatium
community composition may have a direct effect on the efficiency of the ecosystem
process by allowing efficient reoxidation of sulfide under a range of different redox
conditions.

GENOMICS MEETS ECOLOGY: UNDERSTANDING ECOLOGICAL
DIVERSITY BY USING METAGENOMICS

Many giant sulfur bacteria exist as communities comprising several related species that
are ecologically and physiologically distinct. Although techniques such as combined
microautoradiography and FISH can provide some information on physiological
differences between coexisting species in natural communities (Gray et aL, 2000), this
provides at best a shuttered view of the properties that distinguish the coexisting

SGM symposium 65

62 N. D.Gray and I. M. Head

(a) (i)

(ii)

(iii)

.1
I

Oh.

—

I
s

*» 1^=1180

1600

0-8
0-7
06
0-5
04
f>3
0-2
r>1

n= 1377

rth»

Time 1

0-5 01 0-01

Initial nitrate concentration (mM)

Achromatium or Thioploca species. With the advent of genomic technologies and the
emergence of metagenomic analysis (DeLong, 2002), the potential now exists to recover
large amounts of genomic information directly from natural microbial communities.
Recent forays into metagenomics have helped to frame some of the limitations of
the approach. The largest-scale metagenomic analysis conducted to date examined the
microbial diversity of the Sargasso Sea (Venter et al., 2004). This breathtaking work
resulted in the recovery of 1 Gb non-redundant sequence, yet still barely scratched the
surface of the diversity present. Even with extreme brute force, the metagenome of
marine ecosystems is a powerful adversary In contrast, a more comprehensive sampling
with less intensive sequencing effort was obtained by focusing on an ecosystem known
to have very limited diversity (acid minewaters from Iron Mountain, CA, USA; Tyson
et al., 2004). It proved possible to reconstruct the genomes from the principal players

SGM symposium 65

Ecology of giant sulfur bacteria 63

(b) (i)

n=1134

100
60

(ii)

E
*-'

"5
IS

0-5
0-4
03
02

0-1

n= 1041 n=883

n=354

n = 126S

Time

0-01

Initial nitrate concentration (mM)

Fig. 7. The effect of nitrate availability on different Achromatium species in a freshwater sediment, (a)
Relative abundance (i) and absolute numbers (ii) of different Achromatium species within sediment
microcosms incubated for 7 days under anoxic conditions with different initial nitrate concentrations.
Empty bars, Achromatium sp. RY5; shaded bars, Achromatium sp. RYKS; filled bars, Achromatium sp.
RY8; diagonally hatched bars, Achromatium sp. RY1; horizontally hatched bars, Eub338-positive cells.
Ammonium (□), sulfate (A) and nitrate (■) in sediment microcosms incubated for 7 days under
different redox conditions (iii). (b) Relative abundance of different Achromatium species in sediment
microcosms incubated for 13 days and supplied repeatedly with nitrate at different concentrations (i).
Empty bars, Achromatium sp. RY5; shaded bars, Achromatium sp. RYKS; filled bars, Achromatium sp.
RY8; diagonally hatched bars, Achromatium sp. RY1; horizontally hatched bars, Eub338-positive cells.
Sulfate (A) and nitrate (■) in sediment microcosms after 1 3 days incubation (48 h after last addition of
nitrate-containing overlying water) (ii). Time-zero analyses were conducted on replicate microcosms
sampled sacrificially at the beginning of the experiments. Error bars on all data indicate sem of triplicate
incubations. Where error bars are not shown, the error bars were smaller than the symbols, n, No.
cells counted for each replicate set of microcosms. Reproduced from Gray etal. (2004) with
permission from Blackwell Publishing.

in the highly acidic environment and, consequently, to link the genomic information
with the specific metabolic and ecological role of the most abundant organisms present.
The analyses of the Sargasso Sea and the acid minewater were both essentially
exploratory in nature. Whilst this is a valid way to proceed and leads to many novel
and exciting discoveries, different opportunities are offered by hypothesis-driven
research. Testing of hypotheses is only possible within the framework of a rigorous
experimental design. The poor sampling afforded by metagenomic analysis of many
environments precludes such an experimental design. The work of Tyson et al. (2004),

SGM symposium 65

64 N. D. Gray and I. M. Head

however, illustrates that it is possible to obtain meaningful sampling of a microbial
community that would lend itself to a defensible comparative analysis. Communities of
giant sulfur bacteria offer an ideal system for the pursuit of hypothesis-driven
metagenomics. Giant sulfur bacteria occur naturally as highly enriched, high-biomass
communities. The communities typically comprise several coexisting species that can
be distinguished genetically at the level of comparative 16S rRNA gene sequence
analysis and the different genotypes can be shown to be physiologically and ecolo-
gically distinct. Extraction of total nucleic acids and generation of high-coverage clone
libraries from such communities have a high chance of reconstructing the genomes of
the most abundant community members. Comparative genome analysis would then
provide the opportunity to identify the genetic basis for the ecological differentiation
observed in natural communities of giant sulfur bacteria.

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Vaucher, J. P. (1803). Histoire des conferves d'eau douce, contenant leurs different modes

de reproduction, et la description de leurs princi pales especes, pp. 1-285. Geneva:

J. Paschoud (in French).
Venter, J. C, Remington, K., Heidelberg, J. F. & 20 other authors (2004). Environmental

genome shotgun sequencing of the Sargasso Sea. Science 304, 66-74.
Winogradsky, S. (1887). Uber Schwefelbakterien. BotZtgAS, 489-610 (in German).
Winogradsky, S. (1888). Beitrage zur Morphologie und Physiologie der Bacterien. Heft 1.

Zur Morphologie und Physiologie der Schwefelbacterien, pp. 1-120. Leipzig: Arthur

Felix (in German).
Wright, D. & Oren, A. (2005). Nonphotosynthetic bacteria and the formation of

carbonates and evaporites through time. Geomicrobiol J 22, 27-53.
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Dubinina, G. A., Dulov, L. E. & Wada, E. (2001). Ecophysiological characteristics of

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northern Baikal. Microbiology (English translation ofMikrobiologiia) 70, 335-341 .
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5530-5537.

SGM symposium 65

Soil micro-organisms in Antarctic
dry valleys: resource supply and
utilization

D. W. Hopkins, 1 B. Elberling # 2 L G. Greenfield, 3

E. G. Gregorich, 4 P. Novis, 5 A. G. O'Donnell 6 and
A. D. Sparrow 3 ' 7

1 School of Biological and Environmental Sciences, University of Stirling, Stirling FK9 4LA,
Scotland, UK

institute of Geography, University of Copenhagen, 0sterVoldgade 10, DK-1350,
Copenhagen K., Denmark

3 School of Biological Sciences, University of Canterbury, Private Bag 4800, Christchurch, New
Zealand

4 Agriculture Canada, Central Experimental Farm, Ottawa, Canada K1A 0C6

5 Manaaki Whenua - Landcare Research, PO Box 69, Lincoln 81 52, New Zealand

institute for Research on Environment and Sustainability, University of Newcastle upon Tyne,
Newcastle upon Tyne NE1 7RU, UK

department of Natural Resources and Environmental Sciences, University of Nevada,
1 000 Valley Rd, Reno, NV 895 1 2, USA

INTRODUCTION

In 1903, the explorer Robert Scott was one of the first humans ever to see the dry valleys
of Antarctica. He called them 'valley(s) of the dead' in which 'we have seen no sign of
life, . . . not even a moss or lichen'. A century later, we know that the soils and rocks are
home to many microscopic organisms that Scott could not have seen.

The dry valleys are part of the small percentage of the land surface of the Antarctic
continent that is ice-free, amounting to about 4000 km 2 , and thus have rock and soil
surfaces that can be colonized by terrestrial organisms. They are an ancient polar
desert, perhaps as much as 2 million years old, located in Victoria Land between about
77 and 79° south (Fig. 1). The valleys are in a precipitation shadow caused by the
Transantarctic Mountains, which rise over 4000 m. The Antarctic dry valleys are now
recognized as one of the harshest terrestrial environments on Earth, characterized by
summer maximum temperatures that rarely exceed 0°C and only a few tens of
millimetres of precipitation, most of which falls as snow and is ablated by strong winds
carrying dry air from the polar plateau - potential evaporation far exceeds precipi-
tation (Fig. 1). The long periods of winter darkness are punctuated by a short summer,
when 24 h daylight is reached for a few weeks either side of the summer solstice and
when the ground surface temperature may rise to a few degrees above zero. This harsh
environment has led to the Antarctic dry valleys being considered as analogues of

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

72 D. W. Hopkins and others

90° W

90° E

1 FMAMJ J ASOND

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Fig. 1. Map of Antarctica showing the location of the dry valleys (•) and the Transantarctic
Mountains (dotted line), and summary temperatures for the dry valleys (filled bars) and McMurdo
Sound (empty bars). McMurdo Sound data were measured at Scott Base on the Ross Sea coast at
78° S; mean annual temperature, -20 °C; mean annual precipitation, 1 80 mm. Dry valley data
were measured at Vanda Station in the Wright Valley at 78° S; mean annual temperature, -20 °C;
mean annual precipitation, 45 mm.

extraterrestrial habitats by the 'astrobiology/exobiology' community (Mahaney et al.,
2001 ;Onoiri etaL, 2004).

The dry valleys are highly significant sites for studies of ecosystem processes because of
their relative biological simplicity (low diversity and small terrestrial biomass), and for
monitoring the effects of environmental changes because they operate under extreme
(both coldness and dryness) climatic conditions; hence, the Taylor Valley is included in
the US National Science Foundation network of Long-Term Ecological Research sites
(e.g. Virginia & Wall, 1999; Wall & Virginia, 1999; Greenland & Kittel, 2002; Hobbie,
2003; Turner et al., 2003). The terrestrial ecosystem comprises micro-organisms, mosses,
lichens and a restricted invertebrate community. Given the absence of a conspicuous

SGM symposium 65

Soil micro-organisms in Antarctic dry valleys 73

community of terrestrial autotrophs in these resource-poor ecosystems, understanding
the carbon and energy supply to terrestrial organisms is an important geomicro-
biological question.

ORGANISMS IN THE DRY VALLEYS

Although there have been no systematic surveys of the diversity of organisms in the dry
valleys, it is known that the soils of the dry valleys support sparse moss and lichen
communities (Bargagli et al., 1999), a low diversity of invertebrates (e.g. Treonis et al.,
1999; Stevens & Hogg, 2002) and small microbial communities, the diversity of which is
largely uncharacterized (Wynn-Williams, 1996; Cowan &C Tow, 2004). Although there is
little information about the soil micro-organisms, there is evidence of microbial activity
or potential activity in the soil (Burkins et al., 2002; Treonis et al., 2002; Parsons et ai,
2004; B. Elberling, E. G. Gregonch, D. W. Hopkins, A. D. Sparrow, P. Novis & L. G.
Greenfield, unpublished results). Probably the most specialized terrestrial microbial
community in the dry valleys is the endo/cryptoendolithic community of lichens, algae
and fungi, living a few millimetres inside relatively coarse-grained rocks. The temp-
erature and moisture regimes inside the rocks are less hostile and the organisms are
sustained by light penetrating the translucent mineral grains (Friedmann, 1982;
Friedmann et al., 1993; Nienow & Friedmann, 1993). However, the most researched
group of terrestrial organisms is the invertebrate animals - the largest residents of the
dry valleys. One hundred years after Scott proclaimed the dry valleys sterile, Wilson
(2002) referred to the soil invertebrates as 'McMurdo's equivalent of elephants and
tigers'. The invertebrate community includes rotifers, tardigrades, acari, collembola
and, most notably, nematodes, which are the largest terrestrial consumers in the dry
valleys (Treonis et al., 1999; Virginia & Wall, 1999; Courtright et al., 2001; Doran et al.,
2002; Stevens & Hogg, 2002, 2003). The most abundant nematode is usually Scottnema
lindsayae, which feeds on bacteria and algae and typically numbers between about 500
and 2000 (kg soil) -1 , whilst at particularly favoured sites, its number may rise to several
thousand kg (Table 1). The enduring presence of relatively large consumers indicates
active energy processing and nutrient cycling.

MICROBIAL ACTIVITY IN DRY VALLEY SOILS

The soils in the dry valleys are characterized by the absence of structure, cohesion and
moisture, typically alkaline pH, the absence of leaching and localized salt accumu-
lations (Campbell & Claridge, 2000). The soils contain only small reserves of organic
matter and nitrogen (Beyer et al., 1999; Burkins et al., 2002; Moorhead et al., 2003;
Barrett et al., 2005; B. Elberling and others, unpublished results). With the exception of
moist soils at the lake margins and in transient stream beds, the soils generally contain
less (typically two orders of magnitude) carbon and nitrogen than the normal range for
temperate soils (Table 2; Barrett et al., 2002; B. Elberling and others, unpublished

SGM symposium 65

74 D. W. Hopkins and others

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SGM symposium 65

Soil micro-organisms in Antarctic dry valleys 75

0-3 i

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Fig. 2. Soil respiration over 1 20 h for soils from different landscape units in the Garwood Valley,
incubated at 5 °C in the laboratory in unamended soil (basal) and in response to addition of water,
glucose, alanine and samples of the microbial mat from the lake margin. Each bar is the mean of
three replicates and the bars are ± sem.

results). Nevertheless, in situ microbial respiration is consistently measured as C0 9
efflux at the soil surface (Parsons et at., 2004; B. Elberling and others, unpublished
results) and Barrett et al. (2005) reported the presence of a significant amount of labile
soil organic matter. Assuming a steady state, the estimated turnover times for organic
matter (mass/flux) in the dry valley soils is remarkably fast: approximately 23 years in
the Taylor Valley (Burkins et al., 2002) and 30-123 years in the Garwood Valley, depend-
ing on soil type (B. Elberling and others, unpublished results). By contrast, turnover
times in most temperate region soils are generally several centuries or even millennia
(Hopkins & Gregorich, 2005). In the absence of a conspicuous community of ter-
restrial autotrophs in the dry valleys, the rapid turnover of organic carbon suggests
either that the soil organic carbon is not in a steady state or the presence of modern
additions of labile organic matter (Barrett et al., 2005). The presence of labile organic
matter is supported by in situ measurements of soil respiration (B. Elberling and others,
unpublished results). In the Garwood Valley, soil respiration at the lake margin was
greater than that in the drier soils from elsewhere in the valley (Fig. 2). This difference

SGM symposium 65

76 D. W. Hopkins and others

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Fig. 3. Relationship of in situ soil respiration and organic C : N ratio for different soils in the Garwood
Valley (B. Elberling and others, unpublished results).

was not related solely to the water content (Fig. 2) and experimental additions of glu-
cose and alanine and of detritus from the lake margin microbial mats (C:N ratio of
10-12) indicated that soil respiration is limited by both carbon and nitrogen (Fig. 2). In
the case of the addition of microbial-mat material, the soil amendment may also have
contained viable organisms that may have contributed to the respiratory response. The
influence of nitrogen on soil respiration is consistent with in situ measurements of soil
C0 2 respiration rates (Fig. 3), which indicate a correlation between the ratio of the total
organic C : total N and soil C0 2 effluxes, with the largest effluxes associated with the
smallest ratios of organic C : total N (B. Elberling and others, unpublished results).

SOURCES OF RESOURCES IN DRY VALLEY SOILS

For most terrestrial habitats, there is little debate over the source of organic resources
for soil micro-organisms. In the dry valleys, however, the sparse and discontinuous
presence of large photoautotrophs means that alternative sources of organic resources
need to be considered. The possible sources of fixed carbon are: (i) modern autotrophic
activity in situ, which would include the cryptoendolithic communities, mosses,
cyanobacteria and heterotrophic algae in the soils, and autotrophic bacteria such as
nitrifiers; (ii) ancient in situ, or 'legacy', organic deposits from a time when the climate
was warmer and conditions were wetter, and organic lake sediments accumulated on
the surfaces that developed into the modern dry valleys; (iii) spatial subsidies from the
coastal regions, where there are abundant marine and ornithogenic deposits carried
into the dry valleys by aeolian dispersal. This would represent relatively long-range
transport of up to about 50 km inland; and (iv) spatial subsidies from the margins of
modern lakes, where microbial mats accumulate under favourable conditions. This

SGM symposium 65

Soil micro-organisms in Antarctic dry valleys 77

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Fig. 4. Isotopic signals for soils from different locations at three sites in the Taylor Valley, shown with
boxes covering the range of isotopic signals for materials from the marine and lacustrine environments
and for cryptoendolithic materials. The sites are each shown by a number indicating the altitude in m
above sea level. Data extracted and redrawn from Burkins etal. (2000) with permission from The
Ecological Society of America.

SGM symposium 65

78 D. W. Hopkins and others

10 000

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Fig. 5. Potential nitrification activities for soils from different landscape units in the Garwood
Valley. The soils were incubated in a shaking incubator at 5 °C in a suspension containing 0-25 mM
(NH 4 ) 2 S0 4 as described by Hopkins et al. (1 988).

material may subsequently be dispersed onto the surrounding soils, representing
relatively short-range aeolian dispersal.

These sources are not necessarily mutually exclusive and, depending on particular site
details, they may all contribute to the overall carbon cycling, albeit in different
proportions. There have been few detailed investigations of the provenance of organic
matter in the dry valley soils (Burkins et al., 2000; B. Elberling and others, unpublished
results) and it is not therefore possible to partition soil biological activity between
different resource drivers. Nevertheless, from the evidence available, it is possible to
evaluate the likelihood of some of the different sources of organic resources being
important contributors to carbon cycling.

Burkins et al. (2000) used the carbon and nitrogen isotopic signatures of soils and
organic materials from within and around the dry valleys to investigate the provenance
of organic matter in the Taylor Valley (Fig. 4). These studies did not provide evidence of
substantial ornithogenic sources of soil organic matter. They did show that, depending
on altitude, organic matter with cryptoendolithic and lacustrine signals was present,
with the cryptoendolithic material occurring at greater altitude, i.e. at drier and colder
sites further from the valley floor and lakes. However, the conclusions were not clear-cut
because, counter-intuitively, the strongest evidence for marine-derived organic matter
occurred at the most inland site (33 km from the coast) examined, rather than at sites
closer to the coast.

SGM symposium 65

Soil micro-organisms in Antarctic dry valleys 79

The data shown in Fig. 5 provide evidence for a chemoautotrophic contribution to
organic matter input that is supported indirectly by the soil NOT concentrations (Table
2). However, given the low productivity (mol NH4 oxidized by autotrophic nitrifying
bacteria) -1 (Wood, 1986) and the fact that the measurements in Fig. 5 were obtained
under laboratory conditions where NH4 nitrogen was not limiting, it is not likely that
nitrification is a substantial source of soil organic carbon.

The 'legacy' model for the Taylor Valley proposes that the origin of much of the soil
organic carbon is material laid down in ancient lake beds. It is supported indirectly by
geomorphological data on landscape origin (e.g. Burkins et aL, 2000; Higgins et aL,
2000), in particular the presence of the palaeolake Washburn in the Taylor Valley
between 22 800 and 8500 years ago, and by some isotopic signatures in soils (Burkins
et aL, 2000). However, it is difficult to reconcile the relatively rapid estimated rates of
carbon turnover with the persistence of an ancient but active reserve of carbon in the
soil, and Barrett et aL (2005) now discuss multiple organic matter sources.

The 'legacy' model may be most relevant in large expanses of dry ground remote from
lakes, in valleys with relatively small areas covered by lakes and where the geology is not
conducive to endoliths. However, in smaller valleys with a larger proportion covered by
lakes, at lake and stream margins and in ephemeral stream beds in both small and large
valleys, which are recognized as hot-spots of biological activity (Greenfield, 1998;
McKnight et aL, 1999; Moorhead et aL, 2003), modern aquatic-derived detritus may
assume greater importance (Moorhead et aL, 2003; B. Elberling and others, unpub-
lished results). The lakes and ponds in the dry valleys are the focus of productivity by
carbon- and nitrogen-fixing cyanobacteria and eukaryotic algae (Olson et aL, 1998;
Hawes & Schwarz, 1999), representing draw-down and concentration of resources
from the atmosphere. Periodicity in lake and stream level may then cause exposure of
resource-rich cyanobacterial and algal detritus and wash-up of nitrogen-rich foams
originating from decomposition of algal biomass within the lake. This conceptual
model for nutrient cycling in such dry valleys is summarized in Fig. 6. Wilson (1965),
Parker et aL (1982) and Greenfield (1998) suggested a linkage between aquatic product-
ivity and nitrogen fixation by cyanobacteria, with terrestrial nutrient cycling consistent
with the spatial-subsidy model (Fig. 6). Whilst this model remains largely unparameter-
ized, such redistribution has been noted in the field (Wilson, 1965; Parker et aL, 1982;
Nienow & Friedmann, 1993; Moorhead et aL, 2003) and is likely to produce a gradient
in soil carbon stocks, declining with distance from lakes, and vertical concentration
gradients in soil profiles (B. Elberling and others, unpublished results).

The redistribution of lacustrine microbial mat from lake shores to soils may, as
suggested above, also lead to redistribution of viable organisms to the soils, although

SGM symposium 65

80 D. W. Hopkins and others

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photosynthesis
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Fig. 6. Conceptual model for resource transfer in one of the small Antarctic dry valleys.

M 1 2 3 19 20 21 37 38 39 55 56 57 73 74 75 +ve -ve -ve M

Primer dimer

PCR product

Fig. 7. Gel showing amplification of methanogenesis-specific methyl co-enzyme reductase (MCR) in
soils from the Garwood Valley by using the ME1/ME2 primer set (Hales et a/., 1996). Lanes 37-39 and
55-57 correspond to the hill-slope and polygon soils in Table 2, respectively.

few studies provide direct evidence for this. However, we have preliminary indirect
evidence from the incidence of methanogenesis and methanogens in the Garwood
Valley We have detected substantial in situ methanogenesis as CH 4 efflux at the surface
in soils within 1 m of lake margins in the range of 1-3-69 mgm _2 day _1 during the
2002-2003 austral summer (E. G. Gregorich, D. W. Hopkins, B. Elberling, A. D.
Sparrow, P. Novis & L. G. Greenfield, unpublished results). These rates are, incidentally,
comparable to CH 4 emissions from temperate and sub-Arctic wetlands (Moore &
Knowles, 1989, 1990). By using PCR, we have detected the gene for the methanogenesis-
specific methyl co-enzyme reductase (MCR) using ME1/ME2 methanogenic primers
(Hales et al., 1996) in soils between 20 and 50 m from lake margin (Fig. 7). DNA
amplified by using this primer set suggests the presence of methanogenic bacteria in the
dry valley soils that may have arisen from the wet, organic-rich lake margins (Table 2).
These data provide circumstantial evidence for transfer of methanogens from the lake

SGM symposium 65

Soil micro-organisms in Antarctic dry valleys 81

margin to the surrounding dry soils, where they were most unlikely to be active under
the dry and well-aerated surface conditions.

CONCLUDING REMARKS

Understanding the provenance of resources for soil micro-organisms in nutrient- and
energy-poor ecosystems, such as polar deserts, requires consideration of a wider range
of possible sources than in most other ecosystems. When the indigenous stocks of soil
organic carbon and nitrogen are as low as occurs in the dry valleys, even very modest
inputs from other sources may represent a significant subsidy. Whilst there is a parallel
in temperate ecosystems at the early stages of primary successions (Hodkinson et al.,
2002), for most ecosystems, in situ primary production is usually, and correctly,
assumed to be the energy source for the below-ground microbial community (Hopkins
& Gregorich, 2005). The Antarctic dry valleys may represent ecosystems dependent to a
greater extent on spatial and/or temporal subsidies than other terrestrial ecosystems.
Clearly, to better understand the relative contributions of resources from the different
sources, parameterization of the models is necessary Our priority is the spatial-subsidy
model (Fig. 6) and this will require measurements of the dispersal of lake-derived
organic matter and organisms and investigation of the periodicity of production and
dispersal in relation to lake level changes.

The dry valleys may also provide an opportunity to examine whether the biological
diversity of the soil community is causally related to function. This question has
emerged recently as a highly topical ecological issue. However, for most soils, the
biological diversity is so large, particularly amongst the heterotrophs, that it is difficult
to quantify and relate to the key carbon-cycling processes. In resource-poor systems,
where the diversity is likely to be lower, there may be a realistic opportunity to examine
soil biodiversity and function.

ACKNOWLEDGEMENTS

We are grateful to Antarctica New Zealand, the UK Natural Environment Research Council, the
Carnegie Trust for the Universities of Scotland, the Royal Society of London, the Transantarctic
Association, the Danish Natural Science Research Council, Agriculture and Agri-Food Canada
and the University of Canterbury (Christ church, NZ) for support for different parts of the
research outlined here. In addition, we express our gratitude to David Wardle (Lincoln, NZ) and
collaborators in the US Antarctic Program supported by the US National Science Foundation,
in particular Diana Wall, Ross Virginia, Jeb Barrett and Byron Adams, for access to some of
their data and stimulating discussions. We are grateful to Lorna English, Patrick St Georges and
M. R. Nielsen for technical assistance, and the numerous staff of Antarctica New Zealand for
logistic support. Finally, D. W. H. wishes to acknowledge the influential role of the late David
Wynn-Williams for an introduction to biological research in Antarctica.

SGM symposium 65

82 D. W. Hopkins and others

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SGM symposium 65

New insights into bacterial
cell-wall structure and physico
chemistry: implications for
interactions with metal ions
and minerals

V. R. Phoenix, A. A. Korenevsky, V. R. F. Matias and
T. J. Beveridge

Department of Molecular and Cellular Biology, College of Biological Science,
University of Guelph, Guelph, Ontario, Canada N1G 2W1

INTRODUCTION

Prokaryotes are the Earth's smallest life form and, yet, have the largest surface
area : volume ratio of all cells (Beveridge, 1988, 1989a). They are also the most ancient
form of life and have persisted on Earth for at least 3-6 x 10 9 years, even in some of the
most extreme environments imaginable, such as the deep subsurface. Most of these
early primitive (and today's modern) natural environments possess reasonably high
amounts of metal ions that are capable of precipitation under suitable pH or redox
conditions. Deep-seated in such geochemical situations is the likelihood of suitable
interfaces that lower the local free energy, so that interfacial metal precipitation is
promoted. Bacteria, being minute and having highly reactive surfaces (interfaces), are
exquisitely efficient environmental particles for metal-ion adsorption and mineral
nucleation. Metal ions interact with available reactive groups (or ligands) on the
bacterial surface and precipitates grow as environmental counter-ions interact with
more and more metal at the site (Beveridge & Murray, 1976, 1980; Beveridge et al.,
1982; Ferris & Beveridge, 1986; Fortin et al., 1998). Once formed, these precipitates are
under the influence of natural geochemical and additional microbially mediated
conditions (Lee & Beveridge, 2001) that instigate the development of fine-grain
minerals, usually via dehydration, so that crystalline phases are eventually developed
(Beveridge et al., 1983). These minerals commence as so-called 'nano-mineral phases'
and grow with time to become larger and larger. This bacterially induced mineral-
ization is probably the natural phenomenon that so encases some cells in fine-grain
minerals that they die and become bona fide 'microfossils' (Ferris et al., 1988). In
ancient times, these mineral-encased prokaryotes, enduring low-temperature

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

86 V. R. Phoenix and others

metamorphic geological conditions, survived as microfossils that are still existent in
such very old Precambrian formations as the ~ 2-0 x 10 9 years Gun Flint Chert, north of
Lake Superior in Canada.

It is certain that bacterial surfaces interact with environmental metal ions and can
provide nucleation sites for mineral precipitation, but it has been extremely difficult to
study such systems with high precision on a cell-to-cell basis, even though a wide base
of techniques exists (Beveridge et al., 1997). This is because the cells are extremely small
and the interactive structures even smaller, and because the reactive sites responsible for
adsorbing metals retain their reactivity only over a certain range of pH and £ h , which is
difficult to monitor over microscale distances. This chapter will outline our new
advances in elucidating the structure of bacterial surfaces, as well as the determination
of reactive sites via potentiometric examination. Additional techniques, such as zeta
potentials and hydrophobic/hydrophilic determinations, will also be given, so as to
provide a more general picture of how surface structure and surface reactivity affect the
natural physico-chemical traits of bacteria.

NEW OBSERVATIONS OF BACTERIAL SURFACES BY USING
CRYO-TRANSMISSION ELECTRON MICROSCOPY (cryoTEM)

By using conventional fixation, embedding and thin-section techniques, the TEM
observation of bacteria has been a most powerful method for elucidating the internal
cytoplasmic organization and the juxtaposition of encompassing envelope layers of
cells (Beveridge, 1989b; Koval 6c Beveridge, 1999). Yet, there are many drawbacks to
such conventional techniques, as the cells are fixed chemically by using harsh fixatives
(such as glutaraldehyde and osmium tetroxide) and dehydrated before embedding in a
plastic resin for thin-sectioning (Beveridge et al., 2005). Essential lipids are extracted,
proteins are denatured and nucleic acids are condensed artificially ... the cells are a
spectre of their former selves. Clearly, the images of such cells have been beneficial to
our initial perception of the structural organization of prokaryotic cells (Beveridge,
1989b), but hydration (which these embedded cells no longer have) is a necessary
prerequisite for the maintenance of native structure. With proper expertise, care and
equipment, it is now possible through the use of cryoTEM to obtain a better and more
natural view of bacteria (Dubochet et al., 1983; Umeda et al., 1987; Beveridge &
Graham, 1991 ; Paul et al., 1993) .

Freeze-substitution

One cryoTEM technique that became popular during the 1980s and 1990s was freeze-
substitution (Hobot et al., 1984; Umeda et al., 1987; Graham & Beveridge, 1990, 1994).
Here, cells are frozen rapidly at approximately -196 °C so as to vitrify them in amor-
phous ice, which is not crystalline and is a kind of glass (Koval 6c Beveridge, 1999;

SGM symposium 65

Cell-wall structure and physico-chemistry 87

Fig. 1. Thin-section image of a freeze-substituted wall of a Gram-positive B. subtilis cell, showing three
distinct regions in the cell wall: the inner (1 ), middle (2) and outer (3) regions that correspond to cell-
wall turnover and to the available reactive groups within the cell-wall network. Bar, 50 nm.

Beveridge et al., 2005). Hence, the cells are physically 'fixed', as there is no time during
freezing for structure to degrade; in fact, freezing occurs within milli- to microseconds
and all molecular motion is stopped. If the bacteria are thawed, they come back to the
living state and continue to grow and divide. This is a clear-cut measure of how well
the cells are preserved! Once vitrification is accomplished and cellular structure is
preserved, the temperature is raised from -196 to -80 °C and the cells are put into a
freeze-substitution mixture. This consists of a cryogenic fluid (such as acetone) con-
taining a chemical fixative (osmium tetroxide), a heavy-metal stain (uranyl acetate) and
a molecular sieve (for trapping water) (Graham & Beveridge, 1990, 1994). At this
low temperature, the cells do not melt; instead they remain vitrified and structure is
preserved. The cellular and surrounding ice (water) sublimes and is trapped in the
molecular sieve. Eventually, the specimen is dehydrated thoroughly and fixed in this
freeze-substitution mixture, with the structure maintaining many of its native features.
Now, it can be embedded in plastic resin, cured and thin-sectioned for viewing by TEM.

Freeze-substitution and Gram-positive (Bacillus subtilis)
cell walls

The results of freeze-substitution are breathtaking, especially when examining the
surface structures (Fig. 1). The fabric of Gram-positive cell walls is no longer featureless
(as seen in conventional embeddings; Beveridge, 2000), but is a tripartite structure
(Fig. 1). The region associated with the plasma membrane (immediately above the
bilayer) is highly contrasted, due to the acquisition of large quantities of heavy-metal
stain. The middle region is lightly contrasted, as the wall polymers (mainly peptido-
glycan) are stretched almost to breaking point, so that the mass : volume ratio is much
reduced compared with other wall regions. The outermost region consists of thin fibres

SGM symposium 65

88 V. R. Phoenix and others

that extend into the external milieu. This tripartite cell-wall structure is compatible
with current models of cell-wall turnover (Graham & Beveridge, 1994; Beveridge,
2000).

More important for this chapter is the fact that the polymeric structure of the wall has
been preserved by freeze-substitution and the available reactive sites within the wall
have been 'decorated' with the heavy-metal stain, so that we obtain a clear picture of
where the reactive sites reside. Many sites are in the region immediately apposed to the
membrane as, here, new wall polymers are being extruded and compacted via
penicillin-binding proteins in the membrane. As an area of de novo wall assembly,
where new polymers are being cross-linked into the pre-existing wall fabric, many
reactive groups are available for decoration in this region, as the mass is great and the
availability of reactive groups is high. However, the middle region is different. It is
the area in the cell wall that resists turgor pressure and maintains the integrity of the
cell. It has to be highly cross-linked to hold the cell together under such a pressure load
and is considerably stretched. This region, then, has few reactive groups left available
for interaction with the heavy-metal stain, as most have been used to cross-link the
network together for strength. The outermost region is an area where the cell wall is
being broken down by the wall's constituent autolysins. Covalent bonds are being
broken and new reactive groups are being made. This region probably has little mass (as
it is being shed during cell-wall turnover), but an excess of reactive sites that are
decorated readily by the stain.

Frozen hydrated thin sections

A more difficult and therefore less used cryoTEM technique is the use of frozen
hydrated sections (Dubochet et al., 1983, 1988). This technique requires skill and
perseverance. As in freeze-substitution, cells are vitrified, but now, instead of processing
the cells so that conventional, plastic thin sections are obtained, this frozen material is
put immediately into a 'cryo'-ultramicrotome and thin-sectioned. The cells are
sectioned whilst vitrified and the frozen sections are mounted immediately into a 'cryo'-
specimen holder and inserted into the 'cryo'-chamber of a 'cryo'TEM. We emphasize
'cryo' because the temperature must be maintained at between -196 and -140 °C during
all manipulations, otherwise the amorphous ice embedding the cells will become
crystalline and ruin the native structure to be observed. No chemical fixatives or heavy-
metal stains are used during the entire process and, as the sections remain vitrified, all
cellular macromolecules and polymers remain in a hydrous state.

Frozen hydrated Gram-positive (6. subtilis) cell walls

One of the advantages of the cryo-sectioning technique is that no artificial chemical
fixatives or heavy-metal stains need to be used. However, implicit in the use of TEM is

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Cell-wall structure and physico-chemistry 89

that the specimen must possess enough density to efficiently scatter high-voltage
electrons from the electron gun of the microscope (i.e. the electron potential is typically
approx. 100000-200000 eV). Biomaterials, once thin-sectioned, rarely have enough
density to effectively scatter such high-powered electrons, as the thin sections are only
~50 nm thick and the biomatter possesses only low-atomic-number elements (such as
H, C, O, N etc.). This is the primary reason that conventional and freeze-substitution
thin sections use stained material; the heavy-metal stains increase the density (or the
mass : volume ratio) of the specimen so that the contrast becomes great enough for
the cells to be visualized easily (Beveridge et al., 2005). Frozen hydrated thin sections of
unstained bacteria do not have this luxury, as they cannot be stained once sectioned; the
staining fluids would immediately freeze over the sample and obliterate the structure of
the cells.

Clearly, then, these frozen hydrated sections of bacteria are difficult to see, as their
contrast is close to that of the surrounding vitrified ice. For this reason, we rely on the
inherent phase function of the lenses of the cryoTEM and use phase contrast to help
imaging by underfocusing to see the bacteria. Certain microscopes (e.g. those with
energy filters) can also derive more additional contrast for the specimen. However, even
then, there are additional problems in visualizing frozen sections. The energy of the
electron beam is often high enough to locally increase the specimen's temperature,
resulting in the formation of crystalline ice (from amorphous ice) and, eventually, in ice
sublimation. As bacteria are excellent nucleation particles (remember how efficient they
are at forming fine-grain minerals), the amorphous-to-crystalline phase transition of
ice frequently occurs on the bacteria and their structure is obscured. Furthermore, as
the specimen is kept so cold, the frozen sections act as 'cold traps' for the condensation
of extraneous molecules within the high vacuum of the microscope column, thereby
often contaminating the structure of the specimen. With all these associated problems,
it is a wonder that frozen hydrated sections of bacteria can be imaged, but they can and
they are extraordinary (Fig. 2).

These images differ from what we see in freeze-substitution images. Remember, now we
have no heavy-metal stains to assist contrast and must rely on the inherent density
imparted on the cell by the constituent atoms within its molecules. Proteins will be
discerned more readily from the surrounding ice than, say, carbohydrates, because they
are usually larger, contain nitrogen and (sometimes) sulfur and tend to fold more
tightly. Most importantly, all cellular constituents remain hydrated and, therefore, not
artificially condensed because of a dehydration regimen. The ribosomes are larger and
more robust and can be seen because of their high concentration of protein and
rRNA (here, the phosphorus adds additional contrast) (Fig. 2). Even the bilayer of
the membrane can be seen, because of the inherent contrast of the phosphorus in the

SGM symposium 65

90 V. R. Phoenix and others

wfLJ *■■ , - -''-' ■

Fig. 2. Frozen hydrated thin section of a B. subtilis cell wall, showing the plasma membrane (PM), the
inner-wall zone (IWZ) and the outer-wall zone (OWZ). Here, the IWZ corresponds to region 1 and the
OWZ to region 2 of the freeze-substituted wall in Fig. 1 . Region 3 is not seen. This IWZ and OWZ have
different dimensions from the regions in Fig. 1 and are visualized entirely, due to the density of the
constituent macromolecules. Bar, 50 nm. Reproduced from Matias & Beveridge (2005) with
permission from Blackwell Publishing.

phospholipids. Most important for this chapter, though, is the cell wall. Here, we get a
clear idea of the mass distribution within the Gram-positive wall of B. subtilis (Fig. 2).
Immediately above the bilayered membrane is a low-density space with little contrast.
Experiments have shown that this is a periplasmic space (Matias & Beveridge, 2005).
Above this is a more densely contrasted region that represents the peptidoglycan-
teichoic acid network of the wall. Unlike freeze-substitutions, there is not an outermost
fibrous region (cf. Figs 1 and 2).

Correlation of freeze-substitution and frozen hydrated
images

How can we reconcile the differences seen in Figs 1 and 2, remembering that the cell in
Fig. 1 has been dehydrated and decorated with a heavy-metal stain? As it is dehydrated,
we would expect that there would be a certain amount of contraction of the wall
regions in freeze-substitution images (Fig. 1), because the structures are no longer
hydrated. We would also expect reactive groups to be labelled. On the other hand,
frozen hydrated structure would not be condensed and this is why the thickness of each
wall region is greater in Fig. 2 than in Fig. 1. This same image shows a periplasmic
space, whereas Fig. 1 does not. The periplasm has contracted and condensed in Fig. 1,

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Cell-wall structure and physico-chemistry 91

but it has been preserved in its natural state in Fig. 2. Accordingly, the periplasm in
Fig. 1 is more concentrated and (it seems) more reactive, as it stains strongly. In Fig. 2,
the periplasm has not condensed and it has low density. The conclusion, then, is that the
natural state of the periplasm in these cells is as a relatively low-density matrix of
highly reactive biomatter occupying a definite periplasmic space, defined by the
membrane and the middle region. Presumably, the periplasm in this space consists of
new wall polymers, secreted proteins (and their associated chaperones) and both
periplasmic enzymes and oligosaccharides (Matias & Beveridge, 2005).

The region above this periplasmic space is wider in frozen sections than in freeze-
substitutions (cf. Figs 1 and 2) and, as this is the hydrated structure, the increased
width shows its natural state. Both figures reveal it to be of relatively low contrast;
freeze-substitutions suggest that there are few available reactive groups and frozen
sections suggest that there is little density (Matias & Beveridge, 2005). Therefore, this
region, as the stress-bearing region of the wall, has been stretched taut (thereby
reducing its mass) and most reactive groups have been utilized to ensure that the
network is cemented firmly together. This forms a strong but elastic fabric of wall
polymers, mainly peptidoglycan (Yao et al., 1999; Pink et al., 2000).

The outermost fibrous region, which is highly decorated with stain in Fig. 1, is not seen
in Fig. 2. Accordingly, this outermost region has so little mass that it cannot be seen, but
is highly reactive. This is in accordance with cell-wall turnover, as this is a region where
autolysins are breaking down old peptidoglycan, making it soluble. This would reduce
its mass whilst, at the same time, generating many new reactive sites due to hydrolysis
(Matias St Beveridge, 2005).

How does this new interpretation of wall structure correlate
with metal-ion interaction and mineralization?

It is undeniable that metal ions interact strongly with bacterial surfaces and mineralize
them (Beveridge & Murray, 1976; Beveridge & Koval, 1981; Beveridge et al., 1983;
Beveridge & Fyfe, 1985; Beveridge, 1989c; Fein et al, 1997; Fortin et al., 1998; Fowle
et al., 2000 ; Daughney et al., 2001 ; Martinez & Ferris, 2001 ; Yee & Fein, 2001 ; Yee et al.,
2004a). Cell walls adsorb metal ions and minerals are nucleated in Gram-positive cell
walls because of metal-ion interaction with the peptidoglycan and secondary polymers
(Fig. 3; Beveridge & Murray, 1980; Beveridge et al., 1982). Our new structural
observations on Gram-positive walls now show the quantity of hydrated biomatter that
resides in the wall for metal interaction (Fig. 2). They also reveal the positions of the
most reactive and likely regions for metal-ion interaction and mineral growth (Fig. 1).
The outermost fibrous region would be the most accessible reactive region and, here,
there would be little problem of metal-ion access and of mineral-particle growth. As

SGM symposium 65

92 V. R. Phoenix and others

* -

Cytoplasm

Fig. 3. Conventional thin section of a B. subtilis cell that has been subjected to 50 mM FeCI 3
treatment for 1 5 min at 22 °C. The iron has begun to precipitate from solution onto the cell wal
(arrows). Notice that most iron is associated with the wall surface and with the periplasm. No
stains other than the iron have been used on this cell. Bar, 50 nm.

Fibrous

Middle

Periplasm
Membrane

Fig. 4. Diagram to explain the regions of the Gram-positive cell wall where metal ions will interact
with reactive sites and develop into fine-grained minerals. The fibrous region refers to region 3, the
middle region to region 2 and the periplasm to region 1 of Fig. 1 . The membrane is the plasma
membrane. The black, crystalline shapes refer to developing minerals and their size represents
mineral abundance.

many wall polymers are in the act of being solubilized, these polymers and their
precipitates would be sloughed from the cells, but could continue to grow and mature
into bona fide mineral phases in the external milieu. The middle region has less
reactivity, as it is highly cross-linked. Because turgor pressure stretches this region
almost to breaking point, the peptidoglycan would be in a relatively 'open' configur-
ation (Pink et al., 2000) so that most metal ions could penetrate through. The inner

SGM symposium 65

Cell-wall structure and physico-chemistry 93

region (or periplasmic space) is a highly reactive, loose gel of polymers and proteins
that would both interact with and precipitate metal ions. Here, though, as the peri-
plasm resides within a confined space between the plasma membrane and the middle
wall region, mineral growth would be restricted to the accessible space.

These correlations and interpretations allow certain predictions to be made as to where
environmental metal ions should interact with Gram-positive bacterial surfaces. Fig. 4
shows a model of these predictions, with most metal minerals associated with the
outermost region and the periplasmic space. Gram-negative surfaces are structurally
more complicated (Beveridge, 1999), but they have also been imaged via freeze-
substitution and frozen hydrated sections (Matias et al., 2003) and are, therefore, also
able to be correlated. Limited space for this chapter prohibits us from doing so.

POTENTIO METRIC PROPERTIES OF CELL SURFACES

The surface charge of any microbial cell is controlled by the density, distribution
and protonation state of ionizable sites in the cell wall. These chemical sites are pre-
dominantly carboxyl, phosphoryl, amino and hydroxyl groups and they deprotonate
with increasing pH, imparting a net negative charge on the surface. The deprotonation
of, say, a carboxylate can be described by the following equation:

B-COOH ^ B-COCT+ FT (equation 1)

where B-COOH is a protonated carboxyl group on the bacterial surface and B-COO~is
its deprotonated form. The log equilibrium constant (pKJ for the reaction in equation
1 is then defined as:

K a = [H + ] x [B-COOI / [B-COOH] (equation 2)

pK a = -logK a (equation 3)

In simple terms, the pX a can be considered as the pH at which a significant number of
those groups or ligands become negatively charged. High-resolution acid-base
titrations (HRABT), combined with a suitable modelling approach, can be used to
determine ligand concentrations and their corresponding pX a values (e.g. Fein et al.,
1997; Cox et al., 1999). This approach works because the adsorption and release of H +
from such ligands causes a shift in the pH away from the expected norm during titration
(Fig. 5). This difference, known as the charge excess, can be used to calculate ligand
concentrations and their pK a values. These calculations can be performed by using
a number of approaches, such as the surface complexation method with fiteql
(Fein et al., 1997; Borrok et al., 2004a), the linear programming method (LPM; Cox
et al., 1999), the fully optimized continuous (focus; Smith & Ferris, 2001; Martinez et
al., 2002) method or the Donnan shell model (Plette et al., 1995; Martinez et al., 2002;
Yeeetal., 2004b).

SGM symposium 65

94 V. R. Phoenix and others

HRABT analysis of the Gram-negative bacterium Pseudomonas aeruginosa PA01,
modelled by using the LPM approach, is shown in Fig. 6. Here, five distinct proton-
binding sites (and their corresponding concentrations) are revealed. Based on model
compounds, each pX a can be assigned to a most probable ligand type. Carboxyl groups
deprotonate over the lower pH range and exhibit a range of pK a values that occur
predominantly between pK a 2 and 6 (Perdue, 1985). Similarly, phosphoryl groups
exhibit intermediate pK a values that are commonly between pK a 54 and 8 (Martell
et al., 1987). Thiols (e.g. cysteine) exhibit pK a values around 8, amines generally
between 8 and 11 and hydroxyls between 9 and 12 (Perdue, 1985; Martell et al., 1987).
Considering this, the pK a groups in Fig. 6 can be assigned as follows: pK^ 4-7, carboxyl;
pK a 5-9, phosphoryl (or carboxyl); pK a 6-8, phosphoryl; pK a 84, amino (or thiol);
pX a 94, amino/hydroxyl.

The pK a spectrum of each bacterium is controlled by the chemical composition of the
cell wall and, thus, each detected pX a and its assigned ligand type can be attributed
to various components of this structure. Carboxyl groups can be associated with
proteins, peptide stems of peptidoglycan and (in Gram-negatives) lipopolysaccharides
[LPSs; either because of ketodeoxyoctonate (Kdo) in the oligosaccharide (unless it is
acetylated; Vinogradov et al., 2002, 2003a, 2004) or because of O side chains, depend-
ing on O serotype]. Phosphoryl groups are associated with teichoic acids in Gram-
positive cell walls (Beveridge et al., 1982), whereas, in Gram-negative outer membranes
(OMs), they are found in the core oligosaccharide and lipid A fraction of the LPS and in
the phospholipids of the membrane's inner face (Beveridge, 1999). Thiol groups, whilst
uncommon in OM proteins (OMPs), may be found in periplasmic proteins. Amines are
found in abundance within the peptidoglycan (peptide stems), as well as in the proteins
of the OM and periplasm.

Importantly, HRABT analysis provides us with both the concentration and pX a
values of these potential metal-binding ligands (Fein et al., 1997). Additionally,
some metal-adsorption modelling approaches utilize a pK a spectrum to underpin their
metal-complexation models (e.g. Fein et al., 1997).

PHYSICO-CHEMICAL STUDIES

Cell-surface complexity and its impact on potentiometric
properties

The nature of the cell wall is complex and diverse, imparting individual surface
characteristics to each bacterial species and strain. Each micro-organism displays a
unique arrangement of ligands and thus has unique potentiometric and metal-binding
properties (e.g. Beveridge & Fyfe, 1985; Martinez & Ferris, 2001; Borrok et al., 2004a).

SGM symposium 65

-300 J

Cell-wall structure and physico-chemistry 95

Base added (ml)

Fig. 5. Raw data from high-resolution acid-base titration (HRABT) analysis. The pH (y axis) is given in
mV. #, Blank titration (no bacteria); O, titration with bacteria.

c
O

c
o
o

-a

03
D5

0-8 -I

07-

0-6

0*5-

0-4-

0-3-

0-2-

0-1-

•f

Fig. 6. p/C, spectrum generated from HRABT analysis of P. aeruginosa PA01 , modelled by using a
linear programming method (LPM). This spectrum was generated from three titrations. Error bars are
SD(a = 1).

As described earlier, novel cryo-based TEM methods have provided further insights into
the distribution of reactive sites within this basic framework. Together with cryoTEM,
HRABT adds another surface-analysis tool to our arsenal for deciphering the
availability of cell-wall ligands. For example, the cyanobacterium Calothrix sp. (strain
KC97) surrounds itself in a thick (up to 1 \im) extracellular polysaccharide (PS) sheath,
which is above the cell wall. This is ion-permeable and ions are able to interact with
functional groups that reside within both the sheath and cell wall. HRABT analysis of
intact cells and isolated sheath has shown that only ~ 15 % of the available ligands are

SGM symposium 65

96 V. R. Phoenix and others

located in the sheath; the remaining 85 % are located in the wall (Phoenix et al., 2002).
Thus, the cell-surface reactivity of Calothrix sp. can be divided into a dual layer, com-
posed of a highly reactive polar cell wall enclosed within a poorly reactive apolar
sheath. The distribution of ionizable groups is likely to be controlled by the organism in
a self-serving manner. For example, the motile phase of Calothrix, called hormogonia,
does not possess a sheath and the electronegative wall is the exposed surface, making
the cyanobacterium hydrophilic. This property seems ideally suited to a free-swimming
phase, which must be in close contact with its aqueous milieu. However, the benthic,
non-motile phase assembles a sheath around cells, making them more hydrophobic and
able to stick to inanimate surfaces. The distribution of ionizable groups over the cell
surface also has implications for metal complexation. The higher concentration of
ionizable groups on the cell wall ensures that this structure remains the main sink for
metals (Yee et al., 2004a).

There are several other general types of polymeric structures besides sheaths that reside
above cell walls (Whitfield & Valvano, 1993), such as capsules and exopolymeric subs-
tances (frequently found in microbial biofilms). All of these structures, including
S layers [i.e. paracrystalline arrays of (glyco)protein; Sleytr & Beveridge, 1999] aid
in adjusting the availability of surface groups on bacteria, and such polymeric controls
on ligand distribution can differ according to bacterial genus, species and (even) strain.
HRABT analysis of eight different strains of Shewanella sp. demonstrated that
those that expressed O side-chain LPS (i.e. smooth strains) commonly exhibited higher
total ligand concentrations than those with just core oligosaccharide (i.e. rough
strains) (V. R. Phoenix, A. A. Korenevsky, T. J. Beveridge, Y. A. Gorby & F. G. Ferris,
unpublished data). Furthermore, comparison of individual pK a values revealed that
Shewanella strains possessing O side chains were relatively enriched in ligands with a
pK a of ~5, suggesting that these chains contained available carboxyl groups. Their
existence was corroborated by structural nuclear magnetic resonance (NMR) analyses
of selected rough and smooth LPS strains (Vinogradov et al., 2002, 2003a, b, 2004).
From these data, one may speculate that strains that exhibit O side chains should
display higher metal-binding capacity than their rough counterparts.

The complex nature of cell surfaces, however, ensures that we cannot assume that all
organisms exhibiting O side chains will contain higher ligand concentrations than
rough strains. This may be due to more significant differences in ligand concentrations
in other components of the cell wall, such as OMPs, or may simply be due to a low
concentration of O side chains or O side chains that are extremely short (and therefore
do not have many ligands on them). This is exemplified by comparing two strains of
P. aeruginosa, one that expresses O side chains (PAOl) and one that does not (rd7513).
Importantly, the density of O side chains on PAOl is typically quite low, with only

SGM symposium 65

Cell-wall structure and physico-chemistry 97

~20 % of the LPS containing O side chains (Kropinski et al., 1987). When the total
ligand concentrations for each strain were evaluated by using HRABT analysis, both
PAOl and rd7513 revealed ligand concentrations that were not statistically different
(V. R. Phoenix & T. J. Beveridge, unpublished data). This reflects the low density of O
side chains dispersed over PAOl's surface, which was insufficient to significantly
enhance the total ligand concentration.

With this in mind, a combined HRABT plus NMR approach can be further exploited
to reveal the density of certain surface polymers. As described above, although some
LPS molecules may express O side chains, many will not and it can therefore be difficult
to estimate total O side-chain concentration over the cell surface. We have combined
NMR and HRABT analyses of isolated LPS to approximate O side-chain coverage on
Shewanella alga BrY. In this example, NMR analysis of the lipid A and LPS core
indicated a 1 :5 stoichiometric relationship between carboxyl and phosphoryl groups.
Additionally, NMR of the O side chain revealed that this component contained one
carboxyl group per repeating structural unit (no other ionizable groups were present).
However, HRABT analysis of the same LPS revealed a 1 : 1 stoichiometry between
carboxyl and phosphoryl groups. Thus, considering a 1 : 5 stoichiometry in the lipid
A and core and the single carboxyl per repeat unit in the O side chain (from NMR
analysis), each LPS polymer must contain, on average, an O side chain of four repeating
structural units (V. R. Phoenix, A. A. Korenevsky, T. J. Beveridge, Y. A. Gorby &
F. G. Ferris, unpublished data).

When performed on whole cells, HRABT analyses probe functional groups embedded
deep within the cell wall (although exactly how deep is uncertain). This provides
additional information about metal-adsorption properties, as metal ions, like protons,
will be able to migrate into the wall matrix. However, HRABT analysis may not be
suitable for evaluating all cell-interface processes. For example, when considering
surface charge to determine electrostatic interactions with mineral surfaces [e.g. during
adhesion, an approach that measures the charge properties of the very outermost
fractions of the cell surface, such as zeta-potential analysis, may be more suitable (see
later section for more details)]. This is because changes in the composition of these
outermost cell-wall polymers can have a marked impact on the zeta potential of a
bacterium, and yet have a considerably smaller impact on the HRABT-determined
properties. This is especially true if the compositions of ionizable groups embedded
deeper within the cell wall are similar for the organisms in question. For example, in a
recent study of several strains of Shewanella sp., the strains displayed a notable diversity
in zero point of charge (ZPC) as determined from zeta-potential analysis, whereas
the ZPCs from HRABT analysis were similar (V. R. Phoenix, A. A. Korenevsky,
T. J. Beveridge, Y. A. Gorby & F. G. Ferris, unpublished data). Thus, as highlighted

SGM symposium 65

98 V. R. Phoenix and others

0-8

(a)

0-6

0-4-

£ 0-2-

4
0-5n (b)

0-4

03-

0-2-

0-1 o

~i 1

9 10

7
p/t

Fig. 7. p/C a spectra from HRABT analysis (modelled by using an LPM approach), (a) pK a spectra of
whole cells of 5. algae BrY (■) and 5. oneidensis MR-1 (O) (n =3, a = 1 ). (b) p/C a spectra from LPS
isolated from 5. algae BrY (■) and 5. oneidensis MR-1 (O).

throughout this chapter, it can be pertinent to understand not only the concentration
and types of reactive groups within the cell wall, but also their distribution.

This is further emphasized by comparing HRABT analyses of whole cells with those
of polymers isolated from the outer surface of the cell wall. HRABT-determined pK a
spectra for S. alga BrY and Shewanella oneidensis MR-1 are shown in Fig. 7(a) ; the two
organisms clearly display quite similar pK a spectra. However, pX a spectra obtained
from LPS isolated from MR-1 and BrY are very different (Fig. 7b). Because LPS is located
on the outer face of the OM, it has a significant impact on the microbe's interaction
with the surrounding environment. Considering the difference in LPS pK a spectra, it is
unsurprising that these two strains display quite different zeta potentials and mineral-

SGM symposium 65

Cell-wall structure and physico-chemistry 99

adhesion properties (A. A. Korenevsky & T. J. Beveridge, unpublished data). This
would not have been anticipated from the pX a spectra of whole cells. Note the
approximate 1 :1 stoichiometry between the carboxyl {pK a ~4) and phosphoryl (pK a
6-8) groups in the LPS of BrY used, as described above, to determine the O side-
chain density on this bacterium. Furthermore, the dominance of phosphoryl groups
(pK a 6-S) in the LPS of MR-1 (Fig. 7b) has also been corroborated by NMR
(Vinogradov et al, 2003a) and further emphasizes the compatibility of these two
methods. The overall lower ligand concentrations exhibited by the LPS of MR-1 are
also noteworthy. These are probably due to the presence of capsular polysaccharide
(CPS) of low ligand concentration associated with the LPS (Korenevsky et al, 2002).

Several studies have noted changes in potentiometric and metal-adsorption properties
in Gram-positive and -negative bacteria in response to environmental conditions
(Daughney et al, 2001; Borrok et al., 2004a; Haas, 2004). For example, B. subtilis walls
exchange their major secondary polymer (teichoic acid) for teichuronic acid under
phosphate starvation (Beveridge, 1989a, c). However, for the purposes of environ-
mental modelling (to predict the transport and fate of metals in aqueous systems), it is
desirable to describe the potentiometric and metal-adsorption properties of a wide
range of bacteria by using a few bulk average parameters (Yee & Fein, 2001; Borrok et
al., 2004a, b). This is viable, providing the error associated with using universal para-
meters is smaller than other errors inherent in environmental modelling (e.g. uncertain-
ty in cell-number distribution throughout the system). However, understanding the
diversity of potentiometric and metal-adsorption properties displayed by different
species is still pertinent because it (i) allows better determination of the average
universal parameters and their associated error and (ii) provides high-resolution data
for specific environments where more accurate modelling may be required.

Hydrophobicity studies

Many studies using electron microscopy of natural biofilms or planktonic cells from
soils, sediments, mine-tailing wastes, hot springs, etc. have revealed that particulate
mineral deposits are commonly associated with microbial cells (Fig. 8; Xue et al., 1988;
Konhauser et al, 1993, 1994, 2004; Konhauser & Ferris, 1996; Small et al, 1999;
Martinez et al, 2003). These minerals may arise not only because of the interaction of
metal ions with ionized functional groups on the cell surface, but also as a result of the
binding of pre-formed, finely dispersed minerals, such as colloidal silica, metal oxides
and clays. Laboratory experiments on the interaction of dissimilatory metal-reducing
bacteria, such as Shewanella putrefaciens, with nano-sized particulate iron oxides
showed strong association between bacterial surfaces and mineral particles (Fig. 9),
so strong that it resulted in irreversible adsorption (Caccavo et al, 1997; Glasauer et
a/., 2001).

SGM symposium 65

100 V. R. Phoenix and others

Fig. 8. Thin section of a natural freshwater-stream sediment taken from Table Mountain in Grosse
Morne Park, Newfoundland, Canada. No electron-microscopic stains have been used and three cells
that are surrounded by minerals have been indicated by asterisks. Bar, 250 nm.

The binding of particulate minerals by cell surfaces is a result of the interplay of
electrostatic (or electric double layer), van der Waals and Lewis acid-base interactions.
The latter are responsible for such interfacial effects as hydrophobic attraction and
hydrophilic repulsion (Van Oss et al., 2001). It has long been recognized that so-called
'hydrophobicity' plays a significant role in the binding of particulate minerals to the
bacterial surface, yet it is clearly a most poorly understood aspect of surface physico-
chemistry Hydrophobicity is a relatively non-specific term that has limited utility, as it
does not provide any scale of magnitude and is comparative; the term draws no actual
line between 'hydrophobic' and 'hydrophilic' surfaces, as one surface is simply more
unwettable than the other. The only reliable quantitative criterion for defining the
relative terms 'hydrophobicity' and 'hydrophilicity' is the sign and value of free energy
of the interfacial interaction between molecules, particles or surfaces immersed in
water (Van Oss & Giese, 1995).

Microbial-surface hydrophobicity is determined by the availability and number of
apolar functional groups (e.g. -CH 3 ) on cell-wall macromolecules. Although there are

SGM symposium 65

Cell-wall structure and physico-chemistry 101

Fig. 9. Conventional thin section of 5. oneidensis MR-1 that has been reacted with a fine-grained
hydrated iron oxide, which has adsorbed to the cell surface. Bar, 200 nm.

no set rules, proteins are considered a major factor for cell-surface hydrophobicity,
whilst PSs determine surface hydrophilicity. The influence of these two major
components of cell walls on surface properties can be illustrated in the Gram-negative
OM, where phospholipids and LPSs assemble into a membrane bilayer and where,
extending from the OM face, PSs strongly affect surface physico-chemistry and
adhesiveness. Surface PSs (capsule or O side chains of LPS) extending from the cell
surface screen hydrophobic OMPs. Smooth strains (expressing O side chains or
possibly CPS) are more hydrophilic than their rough counterparts, which possess only
core oligosaccharide on their LPS (Williams et al., 1986; Makin & Beveridge, 1996;
Williams & Fletcher, 1996; Flemming et aL, 1998; DeFlaun et al, 1999; Faille et al,
2002). However, this is not always true; sugar residues of microbial PS are often
substituted with apolar methyl and acyl groups (Whitfield & Valvano, 1993), which can
contribute to surface hydrophobicity (Makin & Beveridge, 1996).

We have recently conducted studies on a number of different Shewanella strains in
which a number of different techniques were used to determine cell-surface
hydrophobicity, including contact angle measurement (CAM) and hydrophobic
interaction chromatography (HIC) (A. A. Korenevsky & T. J. Beveridge, unpublished
data). The first method, CAM, yielded quantitative parameters of cell-surface
hydrophobicity and assessed the free surface energy by using the Lifshitz-van der
Waals/Lewis acid-base approach. Here, measurements showed the overall character
of all Shewanella surfaces to be hydrophilic and essentially monopolar, with low
electron-acceptor but high electron-donor parameters. The second method, HIC, was
chosen as the least perturbing cell-surface protocol, as it measures live bacteria in their
natural, fully hydrated state (Pembrey et al, 1999). It is based on the microbial
interaction with a hydrophobic substrate (octyl Sepharose) and appeared to be much

SGM symposium 65

102 V. R. Phoenix and others

more sensitive to ultrastructural variations of the bacterial surfaces than CAM. The
HIC values of Shewanella showed a strong relationship with known LPS/PS composi-
tions (A. A. Korenevsky & T. J. Beveridge, unpublished data) and were in good agree-
ment with previous studies where increased cell-surface hydrophobicity was found in
strains expressing short O side chains or rough LPS (Hermansson et aL, 1982; Williams
et aL, 1986; Makin & Beveridge, 1996; Williams Sc Fletcher, 1996; Flemming et aL,
1998; Hanna et aL, 2003). The higher hydrophobicity of rough strains is considered to
be the consequence of increased exposure of hydrophobic OMPs.

In our studies, electrostatic interaction chromatography (ESIC) and microelectro-
phoresis (zeta potential) were also used and provided information on cell-surface
electronegativity (A. A. Korenevsky & T. J. Beveridge, unpublished data). Both types
of measurements showed that the presence of smooth LPS or CPS decreased cell-
surface electronegativity Even though encapsulated bacterial strains are often reported
to be more electronegative than rough LPS phenotypes, the opposite was found
with our Shewanella strains. Both methods showed the cells with PS to be significantly
less electronegative. This tendency has been reported previously for a number of
Gram-negative bacteria (Hermansson et aL, 1982; Makin & Beveridge, 1996; Flemming
et aL, 1998; Razatos et aL, 1998; DeFlaun et aL, 1999). Our recent LPS/PS structural
analyses help to explain these present electronegativity results (Vinogradov et aL,
2002, 2003a, b, 2004). The common repeating motif of the S. alga BrY O side chain
consists of four monosaccharides and contains 3-hydroxybutyric acid and malic acid.
Here, the carboxyl group of malic acid is the only available ionizable group in the
structure (Vinogradov et aL, 2004). A similar situation was seen in the CPS of
S. oneidensis MR-4; here, the repeating unit contains only one glucuronic acid as an
ionizable residue (E. Vinogradov, A. A. Korenevsky & T J. Beveridge, unpublished
data). In contrast, Shewanella 's core oligosaccharides were found to be highly phos-
phorylated (Vinogradov et aL, 2002, 2003, 2004b; Moule et aL, 2004) and thus very
electronegative at circumneutral pH. Accordingly, Shewanella^ O side chains and CPS
seem to be only weakly charged and are capable of screening the main surface charge
located in the core-lipid A region of LPS. Unlike most other shewanellae, S. oneidensis
MR-1 possesses a surprisingly low surface charge at circumneutral pH, even though it
expresses rough LPS. The core of its LPS contains an unusual component, 8-amino-
Kdo (Vinogradov et aL, 2003), which may account for the lowered surface charge, as the
amino group masks the carboxylate. Furthermore, this strain possesses a microcapsule
of 20-30 nm that, like the PSs of MR-4, could possess a low charge density (Korenevsky
etal., 2002).

It at first seems to be a contradiction that HIC suggests that the more polar
(electronegative) strains are more hydrophobic than the less electronegative strains. The

SGM symposium 65

Cell-wall structure and physico-chemistry 103

answer to this contradiction could reside in our traditional perception of OM structure
and the arrangement of surface macromolecules. We too often consider this membrane
to be a homogeneous, static assembly of macromolecules that are dispersed randomly
over the bilayer's surface. Instead, all molecules are in high motion (usually over
extremely short timescales) and it is probable that, at distinct time intervals, a mosaic
of molecular patches of definite polarity and charge exists (Sokolov et al., 2001;
Korenevsky et al., 2002; Vadillo-Rodriguez et al., 2003, 2005). Given an inanimate
surface of consistent hydrophobicity (such as the Sepharose beads used for HIC), the
tendency of the OM surface would be to congregate and align compatible macro-
molecules towards the inanimate surface. Therefore, although the overall Shewanella
cell surface was found to be hydrophilic by other analyses (e.g. CAM; A. A. Korenevsky
& T. J. Beveridge, unpublished data), it still may be capable of hydrophobic interactions
through such apolar congregations. This would lead to adhesion to hydrophobic sub-
strata, such as was seen in our HIC experiments.

It is clear that hydrophilicity and hydrophobicity have a strong bearing on the para-
meters that affect geomicrobiology. The accessibility of metal ions to reactive sites on
bacterial surfaces, biomineral development on such surfaces, adsorption of pre-formed
fine-grained minerals and adhesion of bacteria to larger mineral phases all depend on
the surfaces' polar and apolar properties. In our Shewanella study (A. A. Korenevsky &
T. J. Beveridge, unpublished data), a strong correlation between such macroscopic
surface parameters as surface negativity, relative hydrophobicity and bacterial adhesion
to haematite was observed. Rough strains exhibited higher affinity and maximal sorp-
tion capacity (by more than an order of magnitude) to haematite when compared with
encapsulated strains. It follows, then, that hydrophobic interactions do not make a
significant contribution to Shewanella^ adhesion to haematite. This is in accord with
acid-base titrations, which demonstrated that hydrous ferric oxide interacts directly
with carboxyl sites on the surface of S. putrefaciens CN32 (Smith & Ferris, 2003;
Martinez etal., 2003).

The assessment of the cell-surface hydrophobicity is a challenging task because
microbial walls have high chemical and structural complexity and are degraded readily
by native enzymes, such as autolysins (Fig. 1). Yet, a good understanding of the
bacterial surface, together with a thorough knowledge of the techniques being used in
the determination of surface physico-chemistry, can be illuminating and can provide
valuable information about the interactions between microbes and the environment. In
this chapter, we have integrated together a variety of techniques (from electron
microscopy to HRABT to hydrophobicity/hydrophilicity studies) into an amalgam to
help explain metal ion-bacterial surface interactions, fine-grained biomineral
development and mineral-sorption/adhesion phenomena. Certainly, there is much more

SGM symposium 65

104 V. R. Phoenix and others

to be deciphered, especially as microbes are dynamic systems that can quickly react and
respond to changing environmental systems, such as altered redox conditions, nutrient
limitation, pH and temperature. Microbial surfaces alter with these environmental
fluxes and therefore can present an entirely different set of interfacial parameters. This
dynamic cellular behaviour and the range of interdisciplinary fields that are required to
study these animate-inanimate systems make geomicrobiology particularly
challenging and exciting.

ACKNOWLEDGEMENTS

The research presented in this chapter was made possible through funding provided by a
Canadian National Science and Engineering Research Council (NSERC) Discovery grant and
a United States Department of Energy Natural and Accelerated Bioremediation Research
Program (US-DOE-NAB1R) grant to T.J.B. The electron microscopy was performed in the
Guelph Regional Integrated Imaging Facility (GRIIF), which is partially funded by an NSERC
Major Facility Access grant to T J. B.

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SGM symposium 65

Horizontal gene transfer of metal
homeostasis genes and its role in
microbial communities of the
deep terrestrial subsurface

Jonna Coombs and Tamar Barkay

Department of Biochemistry and Microbiology, Cook College, Rutgers University, 76 Lipman Dr.,
New Brunswick, NJ 08901, USA

INTRODUCTION

Both basic and applied science issues drive our interests in the microbiology of the deep
terrestrial subsurface. As an environment that is disconnected from the Earth's surface,
the deep subsurface is less subject to variations in temperature and light and, in un-
saturated zones, to intense gradients across interfaces created at the microscale level.
These characteristics dictate an average growth rate that is very slow, up to thousands of
years per cell division (Kieft 6c Brockman, 2001), and an ecosystem where change occurs
over very long time scales (Fredrickson 6c Onstott, 2001). Thus, the subsurface is one of
the most extreme environments on Earth, and identifying what limits life in the sub-
surface has value as a model for life on other planets (Chapelle et al., 2002; Nealson 6c
Cox, 2002). The inadvertent release of contaminants from industrial processing plants
and storage tanks, as well as the possibility of permanently depositing nuclear wastes
deep below the Earth's surface (Pedersen, 2001), raise questions about how microbial
activities might exacerbate or mitigate contamination problems in the subsurface.

The terrestrial subsurface is the habitat for diverse microbial communities that,
together with the oceanic subsurface, may be the habitat for the largest proportion of
Earth's biomass (Whitman et al., 1998). As subsurfaces are characterized by a range
of physical and chemical properties, from fully aerated sedimentary shallow aquifers to
deep igneous rocks devoid of oxygen and elevated in temperatures, their microbial
communities are equally varied (Fredrickson 6c Fletcher, 2001). Microbial life in
the subsurface is greatly constrained by temperature, pressure, limited space and
availability of water and scarce resources of electron donors, acceptors and micro-

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

110 J. Coombs and T. Barkay

nutrients, challenging microbial life to its limit (Colwell, 2001). Nevertheless, studies
over the last 30 years have revealed metabolically and phylogenetically diverse microbial
communities in the subsurface (Amy et al., 1992; Balkwill et al., 1997).

Microbial biomass and diversity in the subsurface

Microbial biomass has been estimated by direct and viable counts and by the quanti-
fication of total phospholipid fatty acids (PLFA) (Kieft et al., 1997; Ringelberg et al.,
1997). Biomass estimates show variability that corresponds to the heterogeneity of the
geological strata that were sampled. Direct counts range from 10 7 cells (g soil) -1 in
the sediments of the Atlantic coastal plain in North America, Rainier Mesa in Nevada
and Witwatersrand Basin in South Africa to 10 4 cells g _1 in deep sediments of the
western USA, and groundwater samples rarely contain more the 10 4 cells ml -1 (sum-
marized by Fredrickson 6c Onstott, 2001). Thus, the microbial biomass of both
the soil and the aqueous subsurface is orders of magnitude lower than those of the
corresponding surface environments.

Representatives belonging to the major prokaryotic lineages have been detected in the
subsurface, as revealed by culturing methods (Balkwill et al., 1997) and by molecular
signatures of both PLFA and 16S rRNA clone libraries (Chandler et al., 1998; Feris et
al., 2004). Interesting unique observations emerge, however, when community structure
and diversity are examined within the context of the unique spatial properties of the
subsurface. For example, in the vadose zone, the area located between the top soil and
the water table, water availability is a limiting resource not so much due to desiccation
(water potentials of >-0-l MPa, sufficient for hydration, are common), but mostly
because water is trapped in small spaces, creating discontinuous environments limiting
transport of microbes, nutrients and toxicants (Kieft & Brockman, 2001). This spatial
discontinuity of microbial niches determines microbial distribution and diversity
patterns and limits microbial interactions to microniches. For example, Takai et al.
(2003b) demonstrated a varied distribution of methanogens in the transition from low-
sulfate and organic- and methane-rich shale to high-sulfate and methane- and organic-
poor sandstone, thus relating community structure to geochemical gradients and
lithologies. Zhou et al. (2002) reported a much higher diversity in surface relative to
subsurface soils, but both had a high degree of species evenness rather than species
dominance, suggesting non-competitive diversity patterns. The authors proposed that,
in soils typified by discontinuity, microbial growth is limited by lack of diffusion of
essential substrates rather than by competition (Zhou et al., 2002, 2004).

Microbial metabolism in the subsurface

Diverse modes of microbial metabolism exist in the subsurface. Heterotrophic
metabolism is supported in aquifers recharged by surface water containing soluble

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Horizontal gene transfer of metal homeostasis genes 111

organic carbon, where the consumption of limited oxygen leads to anaerobic con-
ditions and the dominance of anaerobic respiratory pathways. Organic carbon accreted
during the slow process of sediment formation is another energy source in the sub-
surface (Krumholz et al., 1997; Colwell, 2001), and microbial degradation of natural
petroleum reserves is an example of heterotrophic anaerobic metabolism with far-
reaching consequences for oil quality and quantity (Aitken et al., 2004; Head et al.,
2003).

Dissimilatory iron reduction is a dominant respiratory pathway in anoxic aquifers
(Lovley et al., 2004). Iron reducers representing common (Petrie et al., 2003) and novel
(Coates et al., 2001) members of the Geobacteraceae are often detected in such
environments. Furthermore, the novel iron reducer Rhodoferax ferrireducens, the first
non-phototrophic species of its genus, was isolated from subsurface sediment
(Finneran et al., 2003). Iron reducers are important targets for bioremediation efforts in
the subsurface because they can use other electron acceptors, among them uranium
(Lovley et al., 2004) and vanadium (Ortiz-Bernad et al., 2004), inducing the formation
of insoluble precipitates. These activities, stimulated in situ by the injection of readily
oxidizable substrates such as acetate (Anderson et al., 2003) or ethanol (North et al.,
2004) into contaminated aquifers, result in the precipitation of metal and radionuclide
contaminants. Invariably, such treatments stimulate growth of Geobacteraceae in the
treated subsurface communities (Anderson et al., 2003; North et al., 2004).

Sulfate reduction is another common respiratory pathway in the subsurface (Wong
et al., 2004), driven by sulfate in groundwater and possibly by the activity of pyrite-
oxidizing bacteria (Ulrich et al., 1998). Sulfide that accumulates during sulfate
reduction may complex with metals and radionuclides (Neal et al., 2004) and retard
their mobility in the subsurface, thus making stimulation of sulfate-reducing bacteria
another strategy for bioremediation in the subsurface.

Chemolithoautotrophic metabolism is a major mode of microbial metabolism in the
subsurface, supporting vast communities of methanogens (Chapelle et al., 2002) and
possibly acetogens (Pedersen, 2001) in environments with very low levels of organic
substrates, such as deep aquifers or igneous rocks. Pedersen (2001) has proposed the
hydrogen-driven biosphere hypothesis to explain microbial life in the latter. According
to this hypothesis, hydrogen and carbon dioxide drive methanogens and acetogens,
which then support the activities of acetoclastic methanogens and acetate-utilizing iron
and sulfate reducers, resulting in the formation of organic polymers that, upon their
degradation, are converted to hydrogen and carbon dioxide. A continuous source of
hydrogen in the subsurface is required to support this hypothesis. The issue of whether
hydrogen can (Stevens & McKinley, 1995; Freund et al., 2002) or cannot (Anderson

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112 J. Coombs and T. Barkay

et al., 1998) be produced in the subsurface by the reaction of water with minerals
is presently undecided. The prevalence and diversity of chemolithoautotrophic meta-
bolism among subsurface microbes is also highlighted by the isolation of a thermo-
philic hydrogen- and sulfur-utilizing chemolithoautotroph from a thermal aquifer,
Sulfur ihydroge nib ium subterraneum, belonging to the order Aquificales (Takai et al.,
2003a). Hydrogen-driven chemoautotrophy by aerobes may be an important, as-yet
unexplored, process in vadose zones where little organic matter is available for hetero-
trophic processes. This lack of information may reflect the fact that the focus of
subsurface microbiology research has been on anaerobic processes in groundwater
aquifers because of the higher biomass and metabolic rates relative to those in the
vadose zone.

In this chapter, we address the issue of the interactions of subsurface micro-organisms
with heavy metals and how they are affected by the exchange of genetic material. Thus,
we touch on the issues of metal homeostasis and genetic diversity in subsurface
microbiology, two topics barely investigated to date.

THE INTERACTIONS OF SUBSURFACE BACTERIA WITH HEAVY
METALS

The contamination of the deep subsurface with mixtures of radionuclides, metals and
organic solvents that may leach into groundwater aquifers is one of the most detri-
mental legacies of the Cold War. Immobilization of these contaminants may be the
only feasible approach to solving this problem, and micro-organisms that convert
inorganic contaminants to less soluble precipitates play a prominent role in in situ
immobilization strategies for the subsurface (NABIR, 2001). As these strategies depend
on the activity of microbes, the presence of a mixture of toxicants may result in the
inhibition of reactions essential for immobilization. It is for that reason that strains of
Deinococcus spp. with high levels of resistance to ionizing and gamma radiation have
been engineered with the ability to degrade organic contaminants and withstand metal
toxicity (Brim et al., 2000, 2003), and that the toxicity of metals and actinides to
bacteria with a potential in bioremediation has been evaluated (Reed et al., 1999;
Ruggiero et al., 2005).

A study to determine the level of metal resistance among subsurface aerobic hetero-
trophic bacteria was initiated, reasoning that these microbes play a role in facili-
tating microbial metabolism in the subsurface (Benyehuda et al., 2003). The microbes
tested were from the subsurface microbial culture collection (SMCC) that is maintained
at Florida State University, USA, and included 261 isolates from the Savannah
River Site (SRS) in South Carolina, USA (borehole P24) and 89 strains from the
Hanford site in Washington state (borehole YB-02). The SRS isolates belonged to

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Horizontal gene transfer of metal homeostasis genes 113

the a-, /5- and y-proteobacteria, as well as to the high-G+C Gram-positive group.
The Hanford strains contained representatives of these taxonomic groups, as well
as the low-G+C Gram-positive group. Resistance to Pb(II), Hg(II) and Cr(VI) was
determined by disk-inhibition assays on solid growth media and by comparison with
the response of well-characterized reference resistant and sensitive strains (Benyehuda
et aL, 2003). Results were analysed for the relationship of metal resistance to the
properties of the tested microbial communities and the environments from which
they were isolated. The major findings were:

(i) Resistance to Pb(II) and Cr(VI) was common among subsurface strains from SRS
and Hanford sediments, while fewer, mostly Gram-positive strains, were resistant to
Hg(H).

(ii) With the exception of a high level of metal tolerance among Arthrobacter spp., there
was no relationship between the phylogeny of the microbes and their metal-resistance
patterns. This is not surprising, as metal resistance is often specified by mobile elements
such as plasmids and transposons (Silver & Phung, 1996; Kholodii et aL, 2002). Some
subsurface Arthrobacter isolates proved to be exceptionally resistant to Cr(VI) and
Hg(II). Other researchers have also reported high levels of metal tolerance among soil
Arthrobacter spp. (Roane, 1999; Megharaj et aL, 2003) and the high abundance of this
genus in soils impacted by mixed-waste contamination (Fredrickson et aL, 2004). Thus,
further investigation of Arthrobacter- -metal interactions is highly warranted.

(iii) Resistances to Hg(II) and Pb(II) were more common in the SRS collection than in
the Hanford collection (ANOVA; P<0-05) and multiple metal resistance was also
higher for the SRS, with 33 % of all strains resistant to more than one metal in this
group compared with 23% for the Hanford group (Fig. 1). Thus, toxic metals
influenced the evolution of resistance more effectively in the SRS community than in the
Hanford community. Varying geological and geochemical factors may explain this
difference. For example, metal toxicity is mitigated by low redox potential and high clay
and organic matter (Collins & Stotzky, 1989; Giller et aL, 1998), and the clay content of
Hanford sediment is higher than in the more sandy SRS sediment. These results
illustrate that, in complex environments, microbe-metal interactions are greatly
impacted by environmental factors, most likely by controlling bioavailability and thus
metal toxicity For more details of this study, see Benyehuda et at. (2003).

This study, as well as others addressing the issue of survival and activities of microbes in
mixed-waste-contaminated subsurfaces, has focused on aerobic bacteria. However,
current in situ immobilization efforts target the activities of metal- and radionuclide-
reducing anaerobic bacteria (Anderson et aL, 2003; Istok et aL, 2004). The few who

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114 J. Coombs and T. Barkay

SRS strains Hanford strains

Fig. 1. Multi-resistance to Pb(ll), Hg(ll) and Cr(VI) among bacteria from the subsurface. The proportion
of strains resistant to none (white segments), one (pale grey), two (dark grey) or all three (black) of the
test metals among strains from the SRS and the Hanford sites are shown. Redrawn with permission
from Benyehuda etal. (2003).

have examined the response of anaerobes to metal toxicity reported conflicting results.
Desulfovibrio desulfuricans G20, a model organism for the immobilization of metals
as sulfides, was susceptible to micromolar concentrations of Cu(II), Zn(II) and Pb(II)
when a medium designed to minimize metal complexation was used (Sani et al., 2003).
Mixed cultures of sulfate reducers were inhibited by Cr(VI) (Smith & Gadd, 2000) and
Cu(II) and Zn(II) (Utgikar et al., 2003). Likewise, Shewanella spp., studied for their role
in immobilizing metals and radionuclides by reducing them to insoluble forms, were
affected by U(VI) (Wade & DiChnstina, 2000) and Cr(VI) (Viamajala et ai, 2004).
These observations clearly suggest that susceptibility to metals and thus the acquisition
of metal resistance by microbes in the subsurface is critical to the success of
bioremediation in environments contaminated by mixed wastes.

THE EVOLUTION OF METAL HOMEOSTASIS GENES BY
HORIZONTAL GENE TRANSFER (HGT) IN SUBSURFACE
MICROBIAL COMMUNITIES

Gene transfer among microbes in the subsurface
environment

Genes encoding resistance to heavy metals are often transferred among micro-
organisms in much the same way that antibiotic-resistance genes travel through
microbial populations, by horizontal gene transfer (HGT) mechanisms. Like anti-
biotics, which are frequently produced by soil organisms, heavy metals are compounds
that subsurface organisms are likely to encounter as part of their environment.

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Horizontal gene transfer of metal homeostasis genes 115

Bioavailable heavy metals are produced naturally by the geochemical weathering of
ores and mobilized with the movement of groundwater. Anthropogenic contamination,
however, may increase the concentrations of toxic metals in a given ecological niche
manyfold. The presence of resistance genes on mobile genetic elements within the
subsurface community is therefore a distinct advantage. The occurrence of HGT in
topsoils, natural waters and in association with the internal and external surfaces
of plants and animals is well recognized. However, HGT has barely been examined in
the deep subsurface. Because population densities (Normander et al., 1998; Licht
et al., 1999) and active metabolism (Smets et al., 1993) stimulate HGT, while most
deep subsurface environments are notorious for low population densities and meta-
bolic rates (Balkwill, 1989; Kieft & Brockman, 2001), the deep subsurface may be the
least conducive environment for genetic exchange. To examine HGT and its role in
the evolution of metal resistance in the subsurface, we have looked for evidence of the
horizontal inheritance of genes encoding P IB -type ATPases in bacteria from subsurface
sediments of the SRS (Coombs &: Barkay, 2004).

P, B -type ATPases and their roles in metal homeostasis and
HGT

P IB -type ATPases are membrane-associated ion pumps that are responsible for
maintaining metal homeostasis by mediating the transport of heavy metals using the
energy generated by the hydrolysis of ATP. Those that are specific for monovalent
cations [Cu(I) and/or Ag(I)] are found in the three domains of life, while those specific
to divalent cations [Zn(II), Cd(II) and/or Pb (II)] are only found among the prokaryotes.
These metal pumps can function in either the import of essential ions or the export of
ions that have reached harmful levels in the cell cytoplasm, depending on the
orientation of the protein in the membrane (Rosen, 2002).

Several P^-type ATPases have been shown to be associated with mobile genetic
elements, from Gram-positive organisms such as Lactococcus lactis (O'Sullivan et al.,
2001), Staphylococcus aureus (Nucifora et al., 1989) and Artbrobacter spp. (K. Jerke
and C. Nakatsu, personal communication) and from Gram-negative organisms such as
Ralstonia metallidurans (Borremans et al., 2001) and Stenotropbomonas maltophilia
(Alonso et al., 2000). However, the occurrence of HGT and its effects on the evolution
of this locus within a specific microbial community had not previously been examined.
We have targeted the genes encoding P IB -type ATPases for a study on the role of
HGT in the evolution of metal homeostasis among subsurface bacteria because of
the importance of metal ion homeostasis for survival in the harsh environment of the
metal-contaminated subsurface. Isolates from the SMCC were selected for study
because of the large number of Pb(II)-resistant bacteria in the SRS community
(Benyehuda et al., 2003).

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116 J. Coombs and T. Barkay

Evolution of P IB -type ATPases by HGT in a subsurface microbial
community

We have used a retrospective approach, based on the recognition of genomic indicators
for evolution by HGT. We reasoned that, because HGT is estimated to occur at rates of
31 kb per million years (Lawrence & Ochman, 1997) and subsurface micro-organisms
reproduce very slowly, prospective approaches to HGT detection, such as microcosm
incubations, may be of little relevance to subsurface communities. However, because
genomic data may be interpreted in different ways, calling into question the validity of
all molecular signatures of HGT (Eisen, 2000), we employed multiple methods in
combination while determining whether or not a gene encoding a P IB -type ATPase was
horizontally transferred. These methods included (i) examining the congruence of the
P IB -type ATPase phylogeny with that of the 16S rRNA gene, (ii) looking for unusual
sequence composition (G+C content) of the P IB -type ATPase when compared with that
of the host genome and (iii) looking for shared indels (insertion/deletion events) among
P IB -type ATPase genes from different organisms.

To obtain P IB -type ATPase genes (zntA/cadA/pbrA-likt genes) from the SRS aerobic
heterotrophs, a nested PCR approach was developed. Novel primer sets for PCR
amplification were designed by aligning conserved domains in zntAIca dAlpbrA-Y\kz
genes that were available in databases. The first PCR targeted the phosphatase domain
and the ATP-binding domain and the second reaction used conserved sequences in
the transmembrane metal-binding domain and the ATP-binding domain. Using nine
PCR primer pairs, amplification products of zntA/cadA/pbrA-like genes from 48 of 105
Pb(II)- resistant subsurface strains were obtained and sequenced. These sequences and
the DNA sequences of 16S rRNA coding genes of the corresponding hosts were then
used to determine whether HGT has contributed to the evolution of zntAlcadAlpbrA-
like genes among the subsurface bacteria. For more details about this approach, see
Coombs 6c Barkay (2004). Phylogenetic incongruence using both parsimony (heuristic)
and distance (neighbour-joining) methods indicated that, in four of the isolates, zntAI
cadA/pbrA-likt genes evolved by HGT (Fig. 2), and three of these were supported by
unusual sequence composition, i.e. G+C content and the presence of indels. All trans-
fers were among the /5- and/or y-proteobacteria, which were the predominant groups
among the Pb (II) -resistant subsurface bacteria from which sequence data were
obtained. Two of these transfers were to Comamonas spp., and comparison of cluster-
ing patterns between the zntAlcadAlpbrA and the 16S rRNA trees suggested that, in
one case (in strain B0669), transfer could have occurred from another Comamonas sp.,
and in the other (in strain B0173), the origin could not be clearly identified (Fig. 2). A
third HGT event, the acquisition of a Pseudomonas-like P IB -type ATPase by Ralstonia
sp. B0665, was not supported by additional sequence features, but the phylogenetic
evidence as indicated by the bootstrap support value was very strong. Finally, a P IB -type

SGM symposium 65

Horizontal gene transfer of metal homeostasis genes 117

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SGM symposium 65

118 J. Coombs and T. Barkay

Table 1. Evidence to support HGT in the subsurface

Supporting evidence is provided either by phylogenetic incongruence of a target gene when
compared to the 1 6S rRNA gene phylogeny of the host organism or by the presence of mobile genetic
elements carrying functional genes (Plasmid). Two-letter abbreviations are used for US states.

Site of isolation Source organism(s) Gene(s) of interest Reference

Phylogenetic incongruence

Savannah River, SC Comamonas spp., P|B"tyP e ATPases Coombs & Barkay (2004)

Pseudomonas spp.

Savannah River, SC ^Proteobacteria tRNA(Leu)(UAA) Vepritskiy ef al. (2002)

South Glens Falls, NY Gram-negative bacteria Naphthalene dioxygenase Herrickef al. (1997)

Plasmid

Metal-plating lagoon Pseudomonas sp. czc,ncc Smets etal. (2003)

Savannah River, SC Heterotrophic bacteria Metal-resistance genes Fredrickson ef al. (1988)

South Glens Falls, NY Gram-negative bacteria Naphthalene dioxygenase Ghiorse ef al. (1995)

Savannah River, SC Sphingomonas sp. F1 99 Catabolic genes Romine ef al. (1 999)

Various Sphingomonas spp. Dibenzo-p-dioxin, Basta ef al. (2004)

dibenzofuran and

naphthalene sulfonates
degradation

Savannah River, SC Sphingomonas sp. F199 Catechol 2,3-dioxygenase Stillwell ef al. (1995)

Savannah River, SC Gram-negative Cryptic Brockman etal. (1989)

Savannah River, SC Sphingomonas spp. 2,3-Dihydroxybiphenyl Kim ef al. (1996)

1,2-dioxygenaseand
catechol 2,3-dioxygenase

ATPase from Acinetobacter sp. strain B0064 grouped phylogenetically within a
/2-proteobacterial clade. The G+C content supported the phylogenetic evidence, indi-
cating that this could have been a recent gene acquisition by the Acinetobacter strain.

These results indicate that HGT has occurred, albeit at low frequencies, during the
evolution of metal homeostasis genes among subsurface bacteria. Other observations
also support these conclusions (Table 1). Plasmid-borne genes for metal resistance
and the degradation of hydrocarbons have been obtained from subsurface isolates, and
phylogenetic analyses have suggested transfer of hydrocarbon-degradation genes in
a shallow aquifer contaminated with coal tar (Herrick et al., 1997). The demonstration
of conjugation in microcosms simulating low-nutrient subsurface soils (Smets et al.,
2003) suggests that HGT can affect the evolution and genetic diversity of subsurface
soil communities.

Here, we used the primary DNA sequences of zntA/cadA/pbrA-likt genes to deduce the
evolutionary pathway of an environmentally important function, metal homeostasis,

SGM symposium 65

Horizontal gene transfer of metal homeostasis genes 119

among subsurface bacteria. While it is tempting to conclude from this study that HGT
occurs in the subsurface, such a conclusion is impossible without clear evidence that the
studied strains evolved in the subsurface. Without such evidence, the alternative
possibility that transfer occurred prior to deposition in the subsurface cannot be ruled
out. Evidence for evolution in situ currently exists for a collection of Arthrobacter spp.
from the Yakima Barricade, where a coherence exists between the 16S rRNA- and recA-
based phylogenies and the geological strata from which the strains originated,
suggesting a long-term evolution, possibly as long as 8 million years, in the subsurface
(van Waasbergen et al., 2000). Examining microbial communities from this and other
similar environments may reveal HGT and other processes that affect genetic diversity
as they have occurred in the subsurface.

Evidence for HGT of P, B -type ATPases in complete prokaryotic
genomes

In order to evaluate the observed frequency of HGT among subsurface microbes, we
analysed genes encoding P IB -type ATPases of 188 bacterial and 22 archaeal genomes.
As a clear phylogenetic distinction between P IB -type ATPases specifying mono- and
divalent pumps does not currently exist, our analysis encompassed all zntA and copA-
like sequences. Only P IB -type ATPases loci that exactly matched the amino acid
sequence of the phosphatase and the transmembrane metal-binding domains of
enzymes with documented activity were included in the collection. The resulting
collections of 311 P IB -type ATPases and the 16S rRNA genes of the corresponding
genomes were subjected to phylogenetic analysis and cases of incongruence between
the two phylogenies were identified. When incongruence was detected, supportive
evidence was sought by examining sequence composition as described above. In
addition, sequences proximal, i.e. within 5 kbp, were examined for the presence
of regulatory genes that might have been co-transferred with the P IB -type ATPase genes.
When found, these were subjected to the incongruence test as above to determine
their phylogenetic affiliation (Coombs & Barkay, 2005).

Twelve instances of phylogenetic incongruence were detected, six of which were
transfers across subclasses within the Proteobacteria (Table 2). This is not surprising,
because our collection of P IB -type ATPases was dominated by proteobacteria. However,
the remaining transfer events were across a longer phylogenetic distance. In two cases,
zntA loci from low-G+C Gram-positive bacteria were found in the genomes of the
y-proteobacteria Stenotrophomonas maltopbilia and Legionella pneumophila and
a transfer of an £-proteobacterial cop A to Ureaplasma parvum, a low-G+C Gram-
positive, was also noted. These findings highlight an extensive involvement of
Gram-positive bacteria in gene exchange. Evidence that zntA and cop A were
transferred to Deinococcus radiodurans from a-proteobacteria emerged from the

SGM symposium 65

120 J. Coombs and T. Barkay

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SGM symposium 65

Horizontal gene transfer of metal homeostasis genes 121

]

Archaea

Hafobacterium NRC-1 3
Methanobacterium thermautotrophicum
Brucella melitensis 1
Mesorhizobium loti 1

Agrobacterium tumefaciens 3
Deinococcus radiodurans 1
Escherichia coli 2
Yersinia pestis 1
Vibrio cholerae 1
Listeria monocytogenes
Oceanobacillus iheyensis
Baciflus hafodurans 3

Nostoc P2 J Cyanobacterium

Bacillus anthracis 2 ~

Bacillus subtilis 1

l Thermoanaerobacter ethenogenes 1 2 Gram-positive
Clostridium perfringens 2
Clostridium acetobutylicum 1
Synechococcus PCC 6803 4 Cyanobacterium

Gamma

Gram -positive

'Pyrococcus abyssf

0-1 changes

^" Helicobacter pylori 2669 2
Helicobacter pylori J99 2
Escherichia coli CopA

Epsilon

Fig. 3. Phylogenetic evidence for the transfer of a zntA gene from a cyanobacterium to the
euryarchaeon 'Pyrococcus abyssi' (boxed). Circles at each node indicate the level of bootstrap support
obtained when analysed by both parsimony and distance methods: black, >80 %; grey, >50 %;
white, supported at > 50 % by one method only.

incongruence analysis and separate locations of these two loci in the genomes of both
D. radiodurans and a-proteobacteria suggested that two independent transfer events
were involved. A single transfer event of cop A between closely related euryarchaeota
was detected in the genome of Pyrobaculum aerophilum (Table 2) and, most excitingly,
a zntA of cyanobacterial origin was present in the genome of the archaeon 'Pyrococcus
abyss? (Fig. 3). Supportive evidence confirming HGT was available for six of the
transfers that were revealed by incongruent phylogenies (Table 2) and in three cases
these included the presence of a gene with homology to regulatory elements that are
known to control expression of zntAlcadAlpbrA loci. Phylogenetic analysis of these
regulatory genes showed that they likely shared an origin with the zntA or cop A genes
they accompanied. This latter criterion also confirmed the cross-domain transfer
between archaea and cyanobacteria.

Thus, it seems that, as we have found with the subsurface strains, the evolution of
genes encoding P IB -type ATPases in sequenced genomes has been subjected to HGT
but that their inheritance has mostly proceeded vertically. This is in contrast to the
well-documented dominance of HGT in the evolution of other traits that enhance
fitness to toxicants, such as resistance to mercury (Kholodii et ai, 2002) and antibiotics.
As P IB -type ATPases mediate metal homeostasis, they may be considered more essential

SGM symposium 65

122 J. Coombs and T. Barkay

to core metabolism than phenotypes that are exclusively involved in detoxification, thus
enhancing stable genomic inheritance. The frequency of transfer detected among the
microbial genomes, 12 of 311, was slightly lower than that among the subsurface
bacteria, 4 of 48 (see above). This difference was most likely due to differences in
the composition of the datasets. The genome study encompassed a broader phylo-
genetic range, and therefore we were able to detect transfers across large phylogenetic
distances, whereas the subsurface study detected transfer among more closely related
organisms. However, it is possible that the frequency of HGT among the subsurface
strains was underestimated. The more closely related microbes are phylogenetically, the
more likely they are to exchange genetic material (Lawrence & Ochman, 1997), but
the less likely are the transfer events to leave a detectable molecular footprint in the new
host genome (Eisen, 2000).

HGT gene microarray

DNA and expression microarrays are a powerful tool in biological research, and
applications to the study of microbial community structure (Small et al., 2001)
and function (Taroncher-Oldenburg et al., 2003; Rhee et al., 2004) have been docu-
mented. Zhou and his collaborators have identified three types of microarrays in
microbial ecology (Zhou 6c Thompson, 2002). Phylogenetic oligonucleotide arrays
(POA) consist of probes homologous to 16S rRNA genes and are used to study
community composition and its response to environmental change. Functional gene
arrays (FGA) are designed to evaluate the metabolic potential of a community by
probing for genes that specify major biogeochemical reactions, including those
essential for biodegradation and bioremediation. Community genome arrays (CGA)
target genes of pure isolates from a specific environment. Our discovery of molecular
signatures indicative of HGT in the genomes of subsurface bacteria (Coombs &
Barkay, 2004) prompted us to develop a fourth type of microarray, possibly a variation
on the CGA, the horizontal gene transfer array (HGT array). This array was designed
to answer the question: 'what are the genetic elements that transfer metal-resistance
genes among subsurface bacteria?'

The HGT array includes 158 oligonucleotide (70-mer) probes specific for genes that
encode replication/incompatibility (inc/rep) loci in 86 broad-host-range (BHR)
plasmids belonging to 13 distinct plasmid groups and 100 probes for metal-resistance
genes. The linkage of metal resistance on specific plasmids is suggested by positive
signals obtained following hybridization of Cy3- or Cy5-labelled plasmid DNA
extracted from subsurface isolates with the array. The emerging patterns classify
plasmids according to their incompatibility groupings and linkage with metal-
resistance genes.

SGM symposium 65

Horizontal gene transfer of metal homeostasis genes 123

Probe F4: designed to the
I ncP1p plasmid pEMT3

Probe F9 designed to the
IncPipplasmid pB4

Probe L5 ; designed to the

p iB'tyP e ATPase of
Deinococcus radiodurans

Fig. 4. HGT array hybridized with Cy3-labelled plasmid from the subsurface strain Comamonas sp.
B01 73. Positive hybridization signals appear as green spots.

Analysis of the four SRS subsurface isolates that were shown to have inherited Pb(II)
resistance by HGT (Fig. 2) indicated that at least two of them carried multiple small
plasmids. Plasmid DNA extracts of these strains and of additional metal-resistant
isolates from contaminated subsurface sediments in Oak Ridge, TN, USA, were
hybridized with the array following optimization and testing with 26 exact-match
reference plasmids (J. Coombs and T. Barkay, in preparation). Of these, a plasmid
extract from Comamonas sp. B0173, a strain that inherited its zntAI cadAlpbrA gene by
transfer from an unknown donor (Coombs & Barkay, 2004), hybridized to two probes
homologous to rep/inc loci of plasmids belonging to IncPl/? and to a P IB -type ATPase
probe (Fig. 4) . This finding indicates that inheritance of Pb(II) resistance in strain B0173
likely occurred by conjugal transfer of an IncPljSBHR Pb (II) -resistance plasmid. Array
studies with a plasmid extract from the other SRS isolate are currently in progress.

The three strains from the contaminated sediments in Oak Ridge, TN, are a part of a
large collection of Gram-positive bacteria where metal-resistance patterns correlated
well with plasmid carriage (Patty Sobecky, personal communication). Strain Bacillus sp.
U26 hybridized to rep/inc probes from a group of characterized Bacillus loci and
hybridized weakly to probes for arsenic-resistance genes. Interestingly, plasmid DNA
from another Bacillus strain, strain V6, hybridized to rep/inc probes homologous to

SGM symposium 65

124 J. Coombs and T. Barkay

those of plasmids that had been described previously in both Gram-positive and
-negative bacteria. In addition, the plasmid preparation hybridized to arsenic-resistance
probes. These results suggest either that two plasmids exist in strain V6 or that a single
arsenic-resistance plasmid has two different origins of replication, and imply that strain
V6 carries plasmids of a broader diversity than has previously been described in other
Gram-positive bacteria. Although our dataset is currently very small, it emphasizes that
interesting, previously uncharacterized metal-resistance plasmids exist in subsurface
soil bacteria.

CONCLUSIONS

The studies reported here focused on two issues that are critical to the activities of
microbial communities in the subsurface, metal homeostasis and HGT. Critical,
because both of these are considerations that affect strategies for controlling the
transport of metals and radionuclides in contaminated subsurface environments.
Results showed a high frequency of resistance to divalent cations and a modest, yet
significant, inheritance of a gene encoding metal homeostasis by HGT. This frequency
of HGT of metal homeostasis genes was similar to that in the microbial world at large
as suggested by the analysis of sequenced microbial genomes. While these results
suggest that HGT may have contributed to the survival of microbes in the harsh
subsurface environment, they could not determine whether transfer occurred in situ.
This question could only be addressed by examining signatures of HGT in the genomes
of subsurface microbes from communities whose evolution in the subsurface can be
documented unequivocally, enabling the histories of microbial speciation to be related
to geological processes. This opportunity may exist in certain ecological niches within
the subsurface, which are permanently isolated from the surface and where microbial
migration is restricted by the discontinuity of niches that support life.

ACKNOWLEDGEMENTS

The research described here was funded by the Natural and Accelerated Bioremediation
Research (NABIR) program, Biological and Environmental Research (BER), US Department of
Energy (grant no. DE-FG02-99ER62864).

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SGM symposium 65

Biosilicif ication: the role of
cyanobacteria in silica sinter
deposition

Liane G. Benning # 1 Vernon R. Phoenix 2 and
Bruce W. Mountain 3

1 Earth and Biosphere Institute, School of Earth and Environment, University of Leeds, UK

2 Molecular and Cellular Biology, University of Guelph, Canada

institute of Geological and Nuclear Sciences, Wairakei Research Centre, Taupo, New Zealand

INTRODUCTION

The contribution of micro-organisms to amorphous silica precipitation in modern
geothermal hot-spring environments has been the topic of intense study in the last three
to four decades. Here, we present a review on the field and laboratory studies that have
specifically addressed bacterial silicification, with a special focus on cyanobacterial
silicification. Studies related to the biogenic silicification processes in diatoms,
radiolarians and sponges are not discussed, despite the fact that, in the modern oceans
(which are undersaturated with respect to silica), the diagenetic 'ripening' of such bio-
genic silica controls the global silica cycle (Dixit et ai, 2001). It is well-known that the
amorphous silica in these organisms (particularly in size, shape and orientation) is
controlled primarily by the templating functions of glycoproteins and polypeptides
(e.g. silaffin and silicatein). For information on these issues, we refer the reader to the
extensive reviews by Simpson & Volcani (1981), Kroger et al. {1997, 2000), Baeuerlein
(2000), Perry & Keeling-Tucker (2000), Hildebrand & Wetherbee (2003) and Perry
(2003). In addition, in terrestrial environments, a large pool of amorphous silica is
cycled through higher plants (grasses and trees) that are believed to use silicification as a
protection mechanism against pathogens and insects. Information on these processes
can be found in the papers by Chen & Lewin (1969), Sangster & Hodson (1986) and
Perry & Fraser (1991).

In this review, we will focus solely on microbial silicification processes, which have
been studied extensively in active geothermal hot-spring environments. These are
characterized by geothermal waters supersaturated with respect to amorphous silica

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

132 L. G. Benning and others

derived from water-rock interaction at depth. The link between microbes and the
surface manifestations of sinter formation (both carbonate- and silica-based) was first
documented in Yellowstone National Park at the end of the 19th century (Weed, 1889).
Since that time, active geothermal systems have been studied widely, due to their
importance as geothermal energy sources and as a proxy to understanding the
formation of epithermal ore deposits, which constitute the deep-seated hydrothermal
features beneath active systems. In the last quarter of the 20th century, a multitude of
studies have been carried out to quantify the formation of silica and carbonate terraces
in active systems, with a view towards understanding whether micro-organisms play
an active or passive role in their formation (Walter et al., 1972; Ferris et al., 1986;
Schultze-Lam etal., 1995; Cady & Farmer, 1996; Konhauser & Ferris, 1996; Jones etal.,
1998, 2001; Konhauser et al., 2001; Mountain et al., 2003). Most of these studies have
focused on the relationships between microbes and the resulting morphology and
structure of modern siliceous sinters. They provide insights into the driving forces
for sinter formation in contemporary deposits and are thus relevant to processes in
Archaean and early Proterozoic settings, where microbes may have become encased and
thus preserved as microfossils (Konhauser, 2000; Cady, 2001; Toporski et al., 2002).

There have also been numerous experimental laboratory microbial silicification studies.
In single-step batch experiments, it has been shown clearly that the affinity of aqueous
silica to bind to a microbial surface is low, regardless of whether the micro-organisms
are equilibrated with solutions supersaturated or undersaturated with respect to
amorphous silica (Fein et al., 2002; Phoenix et al., 2003; Yee et al., 2003). Such single-
step experiments do not reliably mimic the processes leading to the significant silica
accumulation observed in hot springs. Other experimental studies have used high
concentrations of either organosilicon solvents, such as tetraethylorthosilicate, or
inorganic silica concentrations and/or a variety of temperatures and pressures to induce
silicification in the presence of micro-organisms and demonstrated that a complex
interplay exists between the precipitation of silica and the formed textures and
structures (e.g. Oehler & Schopf, 1971; Leo & Barghoorn, 1976; Walters et al., 1977;
Francis et al., 1978; Ferris et al., 1988; Westall et al., 1995; Konhauser et al., 2001;
Toporski et al., 2002; Mountain et al., 2003). These studies offered important insights
into the diagenetic-related fossilization processes and sinter textural development, but
they cannot provide mechanistic data pertaining to molecular-level interactions
between micro-organisms and silica accumulating in environments such as hot springs
or the ancient oceans.

Many studies of microbial silicification in active hot springs have shown that silicifi-
cation rates are rapid, but that the silicification process is controlled by purely abiotic
driving forces [i.e. boiling, cooling, evaporation, waves and splash; see Mountain et al.

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Biosilicification by cyanobacteria 133

(2003) and references therein]. Microscopic analysis of silicified micro-organisms from
active hot springs shows that the microbial surface may act as a nucleation site for silica
precipitation (Schultze-Lam et aL, 1995; Konhauser & Ferris, 1996; Jones et aL, 2000;
Phoenix et aL, 2000; Mountain et aL, 2003). Recent studies that exposed cyanobacteria
repeatedly to freshly prepared, supersaturated, polymerizing silica solution (a pseudo-
flow-through setting) have shown that extensive biomineralization, similar to that
observed in hot springs, can be induced (Phoenix et aL, 2000; Benning &: Mountain,
2004; Benning et aL, 2004a, b), with similar structures and textures to those observed in
the field (Benning & Mountain, 2004; Fig. 1). Based on detailed microscopic and, more
recently, spectroscopic measurements of samples from such laboratory experiments, it
is now believed that the accumulation of amorphous silica on the surface of cyano-
bacteria is controlled solely by silica nanoparticle aggregation, but that the contri-
bution of the microbial sheaths or cell walls in this aggregation process is considered
significant (see below; Phoenix et aL, 2000; Benning et aL 2004a, b). Lastly, recent
studies by van der Meer et aL (2002) and Pancost et aL (2005) have shown that specific
biomarker lipids can be preserved in natural modern silica sinters. Such biomarker
studies can provide insight into the complex community structure of thermophilic and
hyperthermophilic micro-organisms (including both archaea and bacteria) that are
present during silica sinter formation. The knowledge of what biomolecules remain
preserved in the rock record may provide a means to extrapolate back in time and thus
to better understand processes in ancient rocks.

In the following pages, we describe the current understanding of the abiotic and biotic
processes occurring in geothermal environments through a review of (i) the chemistry
of silica and the thermodynamic and kinetic aspects of precipitation, (ii) the role of
specific components of the microbial cell surface and (iii) the pathways of silica-colloid
interaction and aggregation on cell surfaces. Only such a synergistic approach can
provide a quantitative model for the reactions that drive microbial silicification and that
lead ultimately to sinter formation and fossil microbial preservation.

THE CHEMISTRY OF SILICA

Soluble silica or monomeric orthosilicic acid (H 4 Si0 4 ) is composed of a silicon atom
coordinated tetrahedrally to four hydroxyl groups. Amorphous silica is defined as a
non-stoichiometric, inorganic polymer made up of a mixture of Si0 2 and H 7 units in
various ratios. Monomeric silica remains stable in solution at 25 °C, as long as its
concentration is below the equilibrium concentration for amorphous hydrated silica [at
25 °C, approx. 100-125 parts per million (p. p.m.); Her, 1979]. In most natural waters,
the concentration of dissolved silica is low (between 1 and 100 ^iM; Treguer et aL, 1995)
and, specifically in marine settings, the silica concentrations are regulated by the growth
of diatoms and radiolarians. In contrast, in the surface expression of active geothermal

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134 L. G. Benning and others

(c)

Fig. 1. Silicified microbes from the New Zealand geothermal hot springs, (a) High-resolution field
emission gun scanning electron micrograph showing silica nanoparticles attached to microbial cells
from the Rotokawa Geothermal Pool; bar, 500 nm. (b) Transmission electron micrograph of silicified
micro-organisms from the Wairakei Geothermal Field; bar, 1 \xm. (c) Transmission electron micrograph
of fully silicified micro-organism from the Wairakei Geothermal Field; bar, 500 nm. Note the small
(30-200 nm) silica particles that form aggregates on the surface of the bacterial sheath.

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Biosilicification by cyanobacteria 135

systems, where temperatures are higher (approx. 30-100 °C), the dissolved silica
concentration in effluent solutions often exceeds the equilibrium solubility of amor-
phous silica. Total silica in hot-spring effluents can be as high as 1000 p.p.m. and this
represents a level many times higher than saturation, even at 100 °C. Subsurface,
geothermal fluids may be undersaturated with respect to amorphous silica but, upon
reaching the surface, drastic changes in temperature and other physico-chemical
parameters will induce the autocatalytic polycondensation/polymerization of silica
monomers, because these changes will induce amorphous silica saturation to be
surpassed. Field experimental determination of precipitation rates showed that the
ratio of monomeric to polymeric silica in the effluent solution plays an important role
in controlling silica-precipitation rates (Carroll et ai, 1998).

In a purely inorganic system, the polycondensation process follows a series of steps that
progress from the polymerization of initial monomers to form dimers, trimers etc. and,
finally, to the formation of highly soluble, critical nuclei of approximately 3 nm in size,
which correspond to approximately 800-900 silicon atoms and have an approximate
molecular mass of around 50 kDa (Her, 1980; Perry, 2003; Icopini et al., 2005). This
initial step occurs via the condensation of two silicic acid molecules and the expulsion
of water:

H 4 Si0 4 +H 4 Si0 4 - (HO) 3 Si-0-Si(OH) 3 + H 2 (equation 1)

Once the silicic acid molecules condense and Si-O-Si siloxane bonds form, cyclic ring
structures will grow and other monomers, dimers etc. will react preferentially with
these nuclei via Ostwald ripening. Dove & Rimstidt (1994) showed that the surface free
energy of such a particle, a [erg cm -2 (1 erg = 10 _7 J)], can be linked with the bulk
precipitate AGf (bulk solid) and the particle surface area (A, cmj to give the free energy of
the particle, AG f( ticle) . This in turn can be expressed as a function of the particle
radius (r, cm; assuming spherical morphology) and the molar volume (V m , cm 3 mor 1 )
via:

AG £ , = [-4^r 3 AG°/2 VI + [4 x 10 1( W 2 a] (equation 2)

f(particle) m

and, from equation 2, an expression for the solubility of a single particle can be derived
(Dove &C Rimstidt, 1994). Alexander (1975) calculated the surface free energy for
amorphous silica in equilibrium with a solution to be approximately 45 erg cm -2
(4-5 x 10 -6 J cm -2 ). This number increases dramatically with particle size and ordering
of the silica phase, reaching a value of 120 erg cm -2 (llxlO -6 ] cm -2 ) for quartz
(Rimstidt & Cole, 1983), thus confirming that smaller and less ordered particles will
dissolve as larger particles grow. Once formed, the critical nuclei will grow to form

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136 L. G. Benning and others

either large nanoparticles (from several hundred nanometres up to a micrometre) or
will aggregate to form three-dimensional complex structures (Her, 1979, 1980; Perry,
2003).

Based on the data of Gunnarsson & Arnorsson (2000), the equilibrium amorphous
silica solubility at temperatures from 20 to 95 °C lies between 100 and approximately
330 p. p.m. Conventionally, the equation representing the equilibrium between silica and
water is written as:

Si0 2 (s) + 2H 2 -> H 4 Si0 4 (aq) (equation 3)

with the equilibrium constant K expressed as the activities (a) of the species:

K=a(U 4 Si0 4 )/a(Si0 2 ) . a 2 (H 2 0) (equation 4)

This reaction is valid for all thermodynamic calculations, but fails to take into account
kinetic effects, as well as the variations in the hydration states of silica. For example, the
ratio between Si0 2 and H 2 in the aqueous species, as well as in the solid, often differs
from the ideal 1 :2, due to hydrogen-bonded waters of hydration in the stoichiometry. In
addition, in equations 3 and 4, aqueous deprotonated and polynuclear species (e.g.
H 9 Si0 4 ~, H 3 Si0 4 and H 6 Si 9 0y _ ) are not taken into account although, in some cases,
such species may contribute to up to 40% of the dissolved silica (Aplin, 1987).
Equation 3 is particularly important in geothermal systems where the geothermal
solutions are supersaturated with respect to amorphous silica, and polymerization
and precipitation are induced due to changes in physical and hydrodynamic con-
ditions.

From a thermodynamic point of view, the precipitation of amorphous silica is driven
by cooling, evaporation, boiling, solution mixing and changes in pH. These factors all
strongly affect the saturation level of amorphous silica. For a general precipitation rate,
Rate v an equation of the type:

Rate ppt =-d[nH 4 Si0 4 ]/dt =-A x k ppt [a{Si0 2 ) . a 2 (H 2 0)] (equation 5)

can be written, where n is the no. moles H 4 Si0 4 , A is the interfacial area (in m") and
k t is the pH-dependent precipitation-rate constant (Her, 1979; Rimstidt & Barnes,
1980; Carroll et ai, 1998). In solutions that are close to saturation, a nucleation barrier
that needs to be surpassed for the first nuclei to form inhibits the precipitation process.
Nielsen (1959) modelled the growth of such nuclei and showed that the flux of mono-
mers towards such nuclei is related to the collision rate, the Boltzman constant,

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Biosilicification by cyanobacteria 137

temperature and the free energy of formation of a critical nucleus. Because quartz has
a higher surface free energy than amorphous silica and its nucleation is inhibited, it
follows that the nucleation and growth of amorphous silica in geo thermal systems are
more favoured.

In most geothermal systems, the amorphous silica that precipitates is composed of
opal-A, a phase that displays varying degrees of ordering of the Si0 4 rings, as well as
varying amounts of structural Si0 2 units and degrees of hydration. Opal-A (nominally
Si0 7 . nH 2 0) is a poorly ordered, highly hydrated phase that displays only one weak,
broad Bragg diffraction band. Other silica phases observed in geothermal silica-
dominated systems are considered good indicators of a diagenetic ageing/altering
process. During this ageing, opal-A is transformed into opal-CT, opal-C, moganite,
cristobalite, chalcedony and ultimately quartz (Herdianita et al., 2000). The main
factors influencing this transformation to more stable counterparts are time, re-
equilibration with high-temperature or high-pH solutions, dehydroxylation/drying
cycles or diagenetic recrystallization. Opal-A can contain between 1 and 13 % water
in its structure; this water is present either as network water or as liquid water in
interstices bound to either internal silanols or defect sites of surface silanols (Langer &
Florke, 1974; Knauth & Epstein, 1982). During this diagenetic transformation/ripening
process, this water is expelled gradually and this is accompanied by a gradual change in

J-spacing for the main Bragg peak from 4-12 to 4-04 A (0-412-0-404 nm). This process
has been used to derive an indicator of structural ripening, as well as a measure of
depth of burial and age. During the ageing and transformation process, water content
drops and particle density increases to 2-3 gem -3 (from as low as 1-5 g cm -3 ). At the
same time, porosity (initially between 35 and 60 %) can be reduced by more than half to
a value below 30 % [see Herdianita et al. (2000) and references therein].

In an effluent solution, the saturation state and thus the precipitation rate of amor-
phous silica are dependent on a variety of parameters that include thermal gradients,
time, changes in pH, concentration of inorganic cations (i.e. Al and trace elements),
organics and ionic strength. Furthermore, this rate depends on the presence of nucle-
ation sites/surfaces, as well as hydrodynamic parameters such as evaporation, waves,
splash etc. As a result, the precipitated, amorphous silica phases will be highly variable
from site to site and the resulting morphology and textures will depend strongly on
these precipitation regimes (with resulting morphologies of the precipitated silica
varying from nanometre-sized spheroidal particles to flat sheets to bulk silica).

The first precipitated opal-A is usually made up of nanometre-sized spheroids that
are later filled in by silica cement to form bulk silica structures. Its formation is a
dynamic process and even 'fresh' sinter features can appear homogeneous; thus, they

SGM symposium 65

138 L. G. Benning and others

are sometimes difficult to distinguish from aged sinters. This has been a major stum-
bling block when purely morphological and structurally preserved biosignatures
observed in modern sinters have been used to relate and extrapolate to processes in
ancient rocks, where subsequent diagenetic or metamorphic processes have homo-
genized and altered the structures, mineral ordering and composition. Specifically,
in ancient rocks, the preservation of biogenic material is hampered by the fact that, in
most cases, the only preserved features are the mould or casings surrounding the
microbes, whilst the cell walls or sheaths have been lost. However, recent geothermal
sinters have revealed that some specific biomarkers (specifically, bacterial and archaeal
lipids) can be preserved. This may be the approach to elucidate the preservation of biota
in ancient rocks lacking unequivocal morphological indicators (van der Meer et al.,
2002; Toporski et al., 2002; Pancost et al., 2005).

CYANOBACTERIAL SURFACE PROPERTIES AND FUNCTION

The structure and composition of cyanobacterial cell walls display a number of
characteristics that are atypical of Gram-negative bacteria. They exhibit a thick, highly
cross-linked peptidoglycan layer (similar to that of Gram-positive organisms) that
makes the cell wall notably stronger (Drews & Weckesser, 1982; Hoiczyk 6c Hansel,
2000). Additionally, biomolecules commonly found in the cyanobacterial outer
membrane, such as atypical fatty acids and carotenoids, are uncommon in other Gram-
negative bacteria (Schrader et al., 1981; Resch & Gibson, 1983; Jurgens & Mantele,
1991).

As with all bacteria, each component of the cyanobacterial cell envelope plays a specific
role. The inner component, the cytoplasmic membrane, behaves as a highly selective
barrier, allowing vital nutrients into the cell and excreting waste material out of the cell.
The outer component of the Gram-negative cell wall is the outer membrane (an
asymmetrical bilayer composed of lipopolysaccharide and phospholipid). This bilayer
also acts as a selective barrier, facilitating the transport of low-molecular-mass
molecules via proteins known as porins. Housed between the two membranes are the
peptidoglycan, which provides rigidity, strength and shape, and the periplasm, which
contains functionally important enzymes. Significantly, these components of the cell
envelope are highly functional, complex and sensitive. One may then ask how these
metabolically vital, and in some cases delicate, layers would respond to encrustation in
silica precipitates.

For some Gram-negative bacteria, the outer membrane is the outermost component of
the organism and it acts as the interface between the cell and the external environment.
In other cases, the organism surrounds itself in an extracellular polysaccharide capsule
or sheath. The structure of this extracellular layer can vary considerably, ranging from

SGM symposium 65

Biosilicification by cyanobacteria 139

<-■-•

Sheath

(impermeable

to Si0 2

colloids)

♦**•*♦«
**■#*

PAR

'**"*♦ UV

Silica colloids
/% Cj in external milieu

Silica colloids accumulate
on sheath's outer surface

Fig. 2. Summary schematic illustrating extracellular silicification of an ensheathed cyanobacterium.
Silica colloids accumulate on the outer surface of the sheath, due to the impermeability of the sheath
to 'large' particles. In addition, it is shown how silica inhibits damaging UV light from reaching the
cell, whilst the photosynthetically active radiation (PAR) can pass through the silica layer with less
attenuation.

diffuse to dense and fibrous. Dense, fibrous polysaccharide layers known as sheaths are
particularly common in cyanobacteria (e.g. Rippka et al., 1979). Moreover, the cyano-
bacterial sheath is known to be devoid of metabolically vital components and it is thus
likely that the organism can withstand a higher degree of damage to this layer than the
rest of the cell envelope. The exact role of the cyanobacterial sheath is not well under-
stood but, as it encloses the more delicate components of the cell envelope, one of its
primary objectives may be to help prevent damage to these components. It has been
shown that a coating of extracellular polysaccharide can protect bacteria against
dehydration (Dudman, 1977; Scott et al., 1996; Hoiczyk, 1998; Tsuneda et al., 2003) and
predation (Dudman, 1977), or it can aid in adhesion to a solid substrate (Dudman,
1977; Scott et al., 1996). More specifically, some cyanobacterial sheaths can contain the
UV light-absorbing pigment scytonemin, aiding cyanobacterial resistance to solar
radiation (Garcia-Pichel & Castenholz, 1991). Of particular relevance to this chapter is
the ability of the cyanobacterial sheath to protect the cell from detrimental bio-
mineralization (Phoenix et al., 2000; Benning et al., 2004a) and to aid in the aggregation
of silica nanoparticles (Benning et al., 2004b).

Sheathed cyanobacteria are found in abundance in hot-spring systems, where it has
been shown that silica accumulation is restricted to the outer surface of the sheath on

SGM symposium 65

140 L. G. Benning and others

living cyanobacteria (Fig. 1; Phoenix et al., 2000; Konhauser et al., 2001; Mountain
et al., 2003). This is likely to occur because the polysaccharide meshwork of the sheath
enables it to act as a filter against colloidal silica (Fig. 2). Permeability studies demon-
strated that the sheath of Calothrix sp. was impermeable to particles of at least 11 nm
in diameter, thus preventing the colloids from biomineralizing the sensitive components
of the cell wall [Phoenix et al. (2000) and references therein].

Interestingly, in the Archaean oceans, which were enriched in silica (Siever, 1992) and
inhabited by cyanobacteria, silica biomineralization was likely to have occurred
(Cloud, 1965). This is particularly true for the shallow waters, in which intermittently
exposed environments were inhabited by stromatolitic communities and evaporation
may have controlled silica precipitation. It is possible that the sheath developed/evolved
in the early oceans as a response and protection against detrimental silica accumulation
on the cell wall. This is supported by several studies, including transmission electron
microscope- and synchrotron-based Fourier-transform IR analysis, which have demon-
strated that, in response to increased silica exposure, the sheath of cyanobacteria
thickens (Phoenix et al., 2000; Benning et al., 2004a, b). Again, this indicates that the
sheath can act as a protective layer against silicification. Naturally, when cyanobacteria
are exposed continuously to supersaturated solutions of silica, silicification eventually
becomes too extensive and this may be detrimental to the cyanobacteria. Phoenix et al.
(2000) have shown that even quite thick silica crusts (approx. 5 Jim thick for a 10 ^im
diameter cell) did not appear to be detrimental to the cells. However, whether there is
a maximum amount of extracellular silicification that cyanobacteria can withstand has
yet to be determined.

One mechanism to overcome extensive silicification may be the release of transient
motile phases (hormogonia) (Herdman & Rippka, 1988) from the ends of heavily
encrusted filaments and this may provide a pathway for survival. Benning et al. (2004a,
b) have followed bacterially mediated silica accumulation both in situ and in vivo via
the changes in the IR signature induced by the increase in silica concentration on the
organic framework of single bacterial cells. This approach allowed the quantification of
the actual and not the apparent bacterial silica-accumulation process, and they showed
that the role of the sheath is twofold. Initially, the cells react to exposure to a silica-rich
solution by producing more sheath polysaccharide as protection. As this thicker sheath
acts as a good ('sticky') substrate for further inorganically precipitated silica-colloid
aggregation, silica precipitation is enhanced, with detrimental effects to the cell.

This ability of cyanobacteria to survive and grow continually, despite extensive
extracellular silicification in modern hot springs and presumably also in the ancient
oceans, may have provided additional advantages to the microbes. This is because

SGM symposium 65

Biosilicification by cyanobacteria 141

amorphous silica biomineralization has been demonstrated to act as an effective
screen against UV radiation (Phoenix et al., 2001). This study has shown that
damaging wavelengths of UV-B and particularly UV-C are absorbed strongly by
amorphous silica, whilst photosynthetically active radiation (400-700 nm) will pass
through the silica with significantly less adsorption (Fig. 2). This process enables cyano-
bacteria to photosynthesize in environments subjected to elevated UV, a protective
mechanism particularly relevant to the Archaean (Phoenix et al., 2001), where highly
detrimental levels of UV irradiated the Earth's surface (Kasting, 1987). Furthermore,
Heijnen et al. (1992) have shown that habitation of micro-niches in bentonite clays can
protect other bacterial forms from predation by grazing protozoa and, thus, silica
encrustation may similarly protect cyanobacteria by making them inaccessible or in-
edible to protozoa. Biomineralization also plays a key role in the formation of siliceous
stromatolitic communities, both modern and ancient, by increasing their structural
integrity and thus longevity (Konhauser et al., 2001). It has also been speculated that,
because the amorphous silica matrix is highly hydrated, it may afford the organisms
an additional protection layer against dehydration. Interestingly, these potential
advantages are similar to the functions of the sheath, and it thus seems that the sheath
and enshrouding silica biominerals may work collectively to protect the organisms
within.

CYANOBACTERIAL BIOMINERALIZATION PATHWAYS AND
COLLOID AGGREGATION

Benning et al. (2004a, b) have followed the processes leading to cyanobacterial
silicification by using in situ IR microspectroscopy and imaging and have quantified the
complex interplay between the cyanobacterial cell components, specifically the sheath,
and the polymerizing silica solution. The progression of nucleation, growth and
aggregation of nanometre-sized silica particles and their effect on cyanobacterial
feedback have been described as a three-stage progression. In the first stage, in response
to the presence of polymerizing silica, the cyanobacteria will increase the formation of
new polysaccharide polymers, i.e. they will thicken their sheath. Concomitantly, silica
will form branched polymers that, upon collapse, will bind to the carbon backbone of
the hydrated polysaccharide sheath via hydrogen bridges. This step can be expressed as:

[>ROH] + [>ROH] -*2[>ROH] (equation 6a)

and

[>ROH] + [sSi-OH] - [=Si-OR<] +H 2 (equation 6b)

where >ROH represents the surface-hydroxylated sugar polymer in the sheath and
=Si-OH is the monomeric silica attached to a surface. Equation 6(b) implies a possible
site-specific silica accumulation, with the silica monomers bound via hydrogen bridges

SGM symposium 65

142 L. G. Benning and others

to another OH-containing radical. The sheath polysaccharides are the obvious candi-
dates for this step. Benning et al. (2004b) used a kinetic approach to show that this
reaction proceeds via a diffusion-limited mechanism (see below), in which polymerizing
silica units in the supersaturated aqueous environment begin to coalesce and aggregate
on the 'fresh' microbial-sheath surface. This process is enhanced once a silane group is
attached to the bacterial sheath via hydrogen bonds and, thus, a further increase in Si
load may lead to the formation of a thin, fully hydrated silica network. Subsequently,
other silane bonds may form independently of the sheath and this process will become
uncoupled from the formation of the silica-carbohydrate hydrogen bonds. This can be
expressed as:

[=Si-OR<] + [=Si-OH] - [=Si-0-Si=] + [ROH<] (equation 7)

At this stage, the formation of additional polysaccharides will no longer compete with
the polymerization of silica, a fact supported by the change in IR spectra, which
become dominated by the more ionic Si-O-Si bonds. The last stage is the formation of
inorganic silane bonds. This process has been shown to be governed by a reaction-
limited process that leads to the growth of purely inorganic Si-O-Si bonds via the
formation of an oxo bridge (Si-O-Si), whilst one water molecule is expelled. This is
similar to the purely inorganic process described in equation 1. When surface
attachment is considered, this step can be described via:

[sSi-OH] + [OH-Sis] - [=si-0-Si=] +H 2 (equation 8)

In this way, a silica network made of corner-sharing [Si0 4 ] tetrahedra is obtained when
all Si— O groups have reacted and the critical silica nuclei have formed. Their further
growth and aggregation will follow and no other connection to the polysaccharide
sheath is needed.

In general, for colloid or polymer growth and aggregation, two restrictive regimes have
been defined: diffusion-limited aggregation (DLA) and reaction-limited aggregation
(RLA) (Everett, 1988; Hunter, 1996; Jamtveit & Meakin, 1999). In the diffusion-limited
case, the limiting step is the movement of two polymer units toward each other prior
to encounter and formation of a cluster (or aggregate). In such reactions, monomers
or oligomers collide and combine instantaneously, producing a relatively porous
aggregate. For the formation of critical nuclei of silica, the DLA process has been
confirmed experimentally (Lin etal., 1990; Martin etal., 1990; Vontoni et al., 2002). For
polysaccharide polymers, however, such data are unavailable, although Rees (1977)
showed that glucose polymers - specifically amylose - grow by a DLA process. On
the other hand, in RLA, the concentrations of the encountered reactant pairs are

SGM symposium 65

Biosilicification by cyanobacteria 143

maintained at equilibrium and thus condensation occurs more slowly. In addition,
a significant repulsive barrier exists, such that the 'sticking probability' upon
oligomer-oligomer interaction is small (Everett, 1988; Gedde, 1995). For silica, the
RLA process results in a more compact aggregate structure during slow condensation
(Martin, 19S7;L'metaL, 1990).

Based on theoretical calculations for nucleation, crystallization, growth and aggre-
gation of mineral phases and organic polymers, Hulbert (1969) and Gedde (1995) have
derived hypothetical constants for the mechanistic constant n, which represents a
parameter that is related to specific mechanisms and geometric shape of the final
mineral particles or polymer. The two types of mechanisms (DLA and RLA) and
several different shapes (needles, plates, spheres, fibres, sheaths etc.) have been
investigated and values for n have been deduced. In general, n increases with increasing
'dimensionality' of the resulting particle/polymer and, in heterogeneous systems, a
change in mechanism often occurs. Hulbert (1969) and Gedde (1995) have concluded
that a particle/polymer of low geometry (one- or two-dimensional; e.g. fibre or sheath),
forming via a DLA process, will have values of n varying between 0-5 and 2, with the
highest values representing two-dimensional growth. On the other hand, if a spherical
(three-dimensional) entity grows or aggregates via a DLA mechanism, the value of n
will vary between 1-5 and 2-5, with the higher numbers indicating a switch to an RLA
mechanism. Lastly, if the same three-dimensional spherical entity grows or aggregates
via a purely RLA mechanism, values of 3-4 are expected for n.

It is well-known that the polysaccharide components of the sheath of Calothrix sp.
(composed primarily of neutral sugars; Weckesser et al., 1988) is usually found in the
form of amylose. Amylose is a linear polymer of glucose units joined by repeating
covalent C— O bonds, and it normally forms complex aggregates of linear geometry
(Rees, 1977). In the cyanobacteria silica system, a low geometric ordering of newly
formed polymers was corroborated by Benning et al. (2004b), who have derived n values
for the first step in polysaccharide growth of 0-8-1-1, thus confirming a one-
dimensional DLA growth for the carbohydrate polymers. For the second step, which is
dominated by the attachment of the formed silica nuclei to the cyanobacterial sheath,
Benning et al. (2004b) derived a value for n of 1-8-2-2, indicating a DLA mechanism for
three-dimensional growth. Finally, for the last step, an n value of 3-5-3-8 was derived,
clearly indicating three-dimensional growth via an RLA mechanism. It needs to be
noted, however, that polysaccharide and silica polymers can both form structures of
mixed or changing geometry during growth or aggregation and that the data derived by
Benning et al. (2004b), which were based on a kinetic approach, may not fully describe
all steps in this complex process. This is particularly true because, in most polymers and
colloid systems, the nucleation and aggregation behaviour in solution is affected

SGM symposium 65

144 L. G. Benning and others

strongly by pH, ionic strength, temperature, organic concentration and type. However,
for silica nucleation and growth, the aggregation steps quantified in the laboratory are
expected to be similar to processes in modern geothermal hot springs and thus can
be used as analogues to model processes in natural environments.

CONCLUSIONS

Cyanobacteria are a major group of phototrophic prokaryotes that play an important
role in the textural development of silica sinters in modern geothermal environments.
Processes that are analogous to those observed in modern hot springs may also have
been active during the fossilization of microbes and the formation of siliceous
stromatolites in the Archaean. These, in turn, can provide a proxy for the biogeo-
chemical conditions of the early biosphere. A large number of field observations and
experimental laboratory studies, as well as a few molecular-dynamic simulations, have
led to the conclusion that, in active geothermal hot springs, cyanobacteria play no
active role in the initial silica polymerization. It was shown that covalent bonds between
silica and bacterial cell-wall or sheath components are not favoured and that the
nucleation of silica from supersaturated aqueous solutions is driven by purely inorganic
polycondensation reactions, which are strongly pH-, ionic strength-, temperature- and
saturation state-dependent. Nevertheless, many field and experimental microscopic
observations showed clear evidence of a link between silica sinter structures/textures
and micro-organisms via the deposition of silica nanospheres onto the microbial
surfaces. Ultimately, this process promotes the incorporation of micro-organisms into
the sinter structure and leads to the preservation of microbial colonies as fossils.
However, despite the large variety of studies carried out so far, the question remains
as to the exact role of the microbial surface in the processes that lead to silica
precipitation. We argue here that the formation of silica sinters is a multi-step process
that is governed primarily by inorganically driven polycondensation of silica monomers
and the formation of silica nanoparticles. This is followed by the microbially enhanced
aggregation of the silica nanospheres into larger assemblages. In the first step,
experimental and theoretical evidence indicates that polymerization of silica mono-
mers leads to the formation of branching clusters that eventually collapse to form a
spherical particle. Although the parameters and mechanisms controlling this collapse
are unclear, some evidence indicates that it may be the dehydroxylation of silanol or
silane clusters. When cyanobacteria and their complex surface structures are present, it
appears that the polycondensation rates and, thus, silica nanoparticle nucleation rates,
are not enhanced. In addition, and more importantly, it is believed that the precipitation
of silica does not affect cyanobacterial metabolism or duplication rates. However,
cyanobacteria do react by increasing the amount of extracellular polysaccharide that
they produce. Once silicification is advanced and thus unavoidable, the thicker poly-
saccharide sheath will enhance the aggregation of the inorganically nucleated silica

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Biosilicification by cyanobacteria 145

nanoparticles into larger silica assemblages. This process occurs while they are alive,
but continued silicification leads to cell death, lysis and finally fossilization.

The surface features seen in active geothermal systems are often regarded as ideal
model systems for study, because they can shed light onto processes occurring in the
shallow-subsurface portions of the geothermal systems that are linked to deep-seated
epithermal ore deposits. In the streams, pools and sinters forming in modern
geothermal systems, a vast array of mesophilic and thermophilic organisms thrive at
high temperatures and varied pH, as well as high toxic-metal concentrations that are
usually detrimental to microbial growth. The knowledge gained from studying such
communities and their interaction with the minerals precipitating from the super-
saturated solutions can give valuable insights into processes of biomineralization. In
addition, our understanding of processes related to the evolution of early life forms in
the Archaean and early Proterozoic has been extrapolated from observations (both
structural and chemical) of bacterial-mineral interactions in modern hot-spring
environments or from morphological observations of Precambrian silicified micro-
fossils and stromatolites.

From field and laboratory observations, reaction pathways for the formation of silica
sinters in modern or ancient hot springs have been determined. The abiotic versus biotic
components of the silica biomineralization reaction, and thus the role of micro-
organisms in this process, were defined. These studies have shown that the silicification
process follows a series of interlinked but unavoidable steps, starting with the microbes
reacting to highly supersaturated silica solutions by increasing their production of
exopolymeric material. Simultaneously with this process, but driven inorganically,
silica polycondensation is proceeding, with monomers condensing to dimers, trimers
etc., leading to the formation of critical silica nanospheres. The newly formed, 'sticky',
exopolymeric sugars do not affect this polymerization, but will subsequently enhance
the aggregation of the inorganically formed nanospheres on the microbial surface.
Finally, this will invariably lead to the full silicification of the organic microbial
frameworks and the formation of microfossils that can thus be preserved in modern
silica sinter environments and that provide a modern analogue to processes in the
ancient past.

ACKNOWLEDGEMENTS

Support for L. G. B. from the Leverhulme Trust (ref #F/00122/F) and the Natural Environment
Research Council (GR9/04623) is gratefully acknowledged. V.R.P. acknowledges the kind
support of Professor Terry Beveridge. B. W. M. acknowledges funding from the Foundation for
Research in Science and Technology (contract C05X0201) and from the NSOF Extremophiles
Programme.

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146 L. G. Benning and others

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SGM symposium 65

Metabolic diversity in the
microbial world: relevance to
exobiology

Kenneth H. Nealson and Radu Popa

Department of Earth Sciences, University of Southern California, Los Angeles, CA 90089-0740,
USA

INTRODUCTION

Metabolic diversity is used here as a physiological or ecological concept referring to
the metabolic repertoire available to any group of organisms: in this case, microbes. At
least for now, metabolic diversity is conceptually distinct from genetic diversity,
although one imagines that, as both concepts are understood in greater depth, the
relationships between them will become clear. The metabolic repertoire encompasses,
for the most part, the entire range of redox-related energy sources that are available on
our planet, from photochemistry to organic and inorganic redox chemistry. Earthly
microbes have 'learned' to harvest the energy of nearly every useful and abundant redox
couple, revealing a nutritional versatility that to some extent could be used to describe
what the planet has to offer energetically To turn this around, one can imagine that, if
energy sources were defined for Earth, one might well predict what kinds of metabolism
should have evolved to exploit them and, in fact, for the most part, this would lead
to the correct answer. Metabolic diversity is further accentuated by various symbioses,
syntrophisms and community interactions (intracellular, intercellular and inter-
population), leading to the establishment of communities with seemingly new and
unexpected abilities. The functional diversity of the prokaryotic world is thus expressed
in terms of its redox chemistry and, with regard to geobiology, this redox chemistry/
metabolic connection defines a wide variety of relevant reactions, many of which
involve phase changes of the interacting molecules (i.e. between solid, liquid and gas
phase). In some sense, it is the world of 'accidental biominerals'; a world where
minerals are formed, altered or dissolved as a result of the forever-starved prokaryotic
world. In contrast, many eukaryotic microbes also exert significant effects on the

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

152 K. H. Nealson and R. Popa

geological landscape, and these work by completely different metabolic modes -
forming minerals by directed, energy-consuming reactions, minerals that have altered
and continue to alter the face of our planet. All in all, the way that life has adapted to
the energy regime of our planet allows the exobiological prediction that the same
should be true of any other abode where life resides. Understanding energy types and
flows may well provide a critical pathway for life detection, on or off our own planet.

MICROBIAL DIVERSITY
Genetic diversity

Genetic diversity is estimated by any of a number of methods, often related to DNA
sequence similarities and differences. For the sake of this discussion, the microbial
world has been divided into two groups, the prokaryotes and eukaryotes (Fig. 1), with
the former group consisting of the Bacteria and the Archaea (Woese, 2004; Woese et al.,
1984, 1990) and the latter including a wide array of single-celled, microscopic
eukaryotes, collectively called the protists or protoctista (Whittaker & Margulis, 1978).
This grouping is based primarily on sequence comparisons of 16S rRNA genes and
includes sequences of both cultivated and as-yet uncultivated microbes. Both the
prokaryotic and eukaryotic microbes are underrepresented with regard to cultivated
members, so that any discussion of metabolic diversity of these groups is of necessity
strongly biased towards a very small subset of the whole. For example, it is estimated
that only 0-1 to 1 % of the prokaryotes in most environments have been cultivated
(Whitman et al., 1998). In fact, there are many phylum-level bacterial taxonomic
groups that are represented by DNA sequences of 16S rRNA genes alone: i.e. groups
for which no species have yet been cultivated (Pace et al., 1986). While it is obviously
inappropriate to make ecological extrapolations based on 1 % or less of the popu-
lation, in many cases it is the best that has been possible.

Metabolic diversity

Metabolic diversity, on the other hand, can be assessed by measurement of the reactants
and products and quantified in the laboratory and the field. Just what organisms are
responsible for a given activity is often discovered after the process has been demon-
strated in natural samples. Thus, the connection between geochemists and micro-
biologists has been a strong one for many years and, in some cases, the geochemists
have observed the processes for which the causative microbes were later discovered and
characterized. Some examples of this are the processes of anaerobic methane oxidation
(Boetius etaL, 2000; Orphan et al., 2001b), anaerobic ammonia oxidation (Jetten et al.,
1998) and microbial iron and manganese reduction (Lovley & Phillips, 1988; Myers &
Nealson, 1988), all of which were known in the geochemical literature prior to the
microbiology being demonstrated. The metabolic diversity of the prokaryotes, how-

SGM symposium 65

Metabolic diversity and exobiology 1 53

Prokaryotes

Bacteria

Archaea

-A.

Eukaryotes

BjghGC

Gram-
positives

Spirochaetes

bpsiton

Alpha

Beta

Gamma

Flagellates

7 he rm odesulfotobacterium

Thermotogales

Aquifex

Fig. 1. Microbial diversity according to molecular phylogeny. This diagram shows some of the range
of microbes as estimated by sequence similarity of 1 6S rRNA gene sequences. This genetically diverse
assemblage is dominated by bacteria and archaea, but includes a large number of very important
eukaryotic microbes such as carbonate- and silica-depositing algae, bacteriovores (ciliates and

flagellates) and photosynthetic algae.

ever, extends beyond individual abilities, as some of the truly unique things done are
accomplished not by single microbes, but by collaborative efforts between microbes
that specialize in the intercellular metabolite transfer called syntrophism, thus
providing even more functional diversity than might be expected from single cells. In
fact, many of the metabolic abilities of eukaryotes (photosynthesis, respiration,
lithotrophy and nitrogen fixation) are carried out by symbionts acquired from the
prokaryo tic world (Margulis, 1981).

One way to view metabolic diversity is shown in Table 1, which illustrates the range
of redox reactions in which microbes participate. As can be seen, the prokaryotes
dominate this scene: only a few organics and no inorganics can be utilized by
eukaryotes, and, with the exception of nitrate and possibly fumarate, which can be used
by a few anaerobic fungi (Fenchel & Finlay; 1995; Tielens et al., 2002; Zumft, 1997),
only one respiratory electron acceptor (oxygen) can be utilized by the eukaryotes.
Purists might argue that even oxygen respiration is an acquired trait - one that was
'invented' by prokaryotes and arose in the eukaryotes via symbiosis (Margulis, 1981),
making respiration a hallmark trait of the prokaryotes and symbiosis, perhaps, one of
the most important parts of eukaryotic evolution. The point is, however, that, with

SGM symposium 65

154 K. H. Nealson and R. Popa

Table 1. Energy sources (fuels) and oxidants used by life

Here we see some of the range of electron donors and electron acceptors used by all of life. The
metabolic diversity of the prokaryotes (which can use all those listed) is dramatically expanded in
comparison with the rather constrained eukaryotes (which can use those entries in bold only).
Asterisks denote those molecules that form insoluble compounds (minerals) when oxidized or
reduced, connecting this redox-driven biochemistry with the geological world. DMSO, Dimethyl
sulfoxide; TMAO, trimethylamine A/-oxide.

Fuels

Oxidants

Sunlight

Fu ma rate

f Glucose

Organics <

DMSO

Organics <

Ethanol

Formaldehyde

Carbon dioxide*

TMAO

l Methanol

Sulfur*

Hydrogen

Sulfate*

Ammonia

Arsenate*

Hydrogen sulfide*

Selenite*

Sulfur*

Iron*

Manganese*

Nitrate

Carbon monoxide
Arsenite*

Oxygen

regard to redox chemistry, the prokaryotic microbes are the experts, while the eukary-
otes have shunned redox diversity in favour of a high energy yield - energy used to
power their own impressive diversity of cellular structure and behaviour. The most
pertinent features with regard to this are: (i) the use of inorganic energy sources (litho-
trophy) is the realm of the prokaryotes; (ii) the process of anaerobic respiration (i.e.
using electron acceptors other than oxygen) is, with few exceptions, the realm of the
prokaryotes; (iii) the oxidation and reduction of these inorganic compounds forms
a strong link with planetary geology, as many of the reactions either form or dissolve
solid-phase minerals and mineraloids during the process (Table 1).

Having extolled prokaryotic diversity, a few words of respect and admiration need to
be inserted with regard to the anaerobic eukaryotes. Both protists and fungi are often
abundant in many different permanently anaerobic niches (i.e. rumens, termite guts,
anaerobic sediments), where they feed on prokaryotes and their breakdown products
(Fenchel & Finlay, 1995). These eukaryotes are almost without exception devoid of
mitochondria and exist either by the use of hydrogenosomes or via nitrate reduction.
Hydrogenosomes, which are common in both anaerobic protists and fungi, are organ-
elles believed to have arisen from mitochondria (Theissen et al., 2003) - they have
replaced the enzymes involved in oxidative phosphorylation with those involved in

SGM symposium 65

Metabolic diversity and exobiology 155

(a)

Light

Cell membrane

R hod opsin =>

(b)

Light
H ; \ 2

V ±_ / 2e*2H

Q -Q,

Chlorophyll
PS II (P680)

FeS Chlorophyll
PS I (P700)

ADP+Pi

NADPH,

* ^ATPr

Fig. 2. Mechanisms for harvesting light energy, (a) Rhodopsin- mediated proton translocation
as seen in archaea (Halobacterium halobium). It should be noted that the proton gradient that is
established by this process can be used to power many cellular functions such as transport or flagellar
motion or, as shown, to power the synthesis of ATP via the membrane-bound ATPase. (b) Oxygenic
photosynthesis in chloroplasts. A and A-,, Acceptors of electrons from P700; bf, cytochrome bf
complex; Fd, ferredoxin; Fp, flavoprotein; PC, plastocyanin; Ph, phaeophytin; Pq, plastoquinone;
PS, photosystem; Q A and Q B , plastoquinone-binding proteins.

fermentation, thus creating an organelle that makes ATP by substrate-level phosphory-
lation and produces dihydrogen gas as a by-product (Martin & Muller, 1998; Muller,
1993).

A final point with regard to diversity relates to the use of light as an energy source. By
a factor of about 50000, visible light dominates our planet in terms of yearly energy
flow (Nealson & Conrad, 1999), and it is used in a number of ways. The simplest is via
a process called rhodopsin-mediated proton pumping, in which the molecule rhodopsin
is used to absorb light and the energy is used directly to pump a proton from the inside
to the outside of the cell, charging the cell membrane via an osmotic and electrical
gradient (Fig. 2a). This gradient can then be used by the cell either directly or to syn-
thesize ATP. This type of metabolism, long known in archaea (Stoeckenius 6c Bogo-
molni, 1982), has more recently been found to be widely distributed in the bacterial
world (Beja et al., 2000; de la Torre et al., 2003; Venter et al., 2004). Photosynthesis, the
more familiar use of visible light by life, was apparently 'invented' by the bacteria,
almost certainly as an anaerobic process, using sulfur compounds or iron as electron

SGM symposium 65

156 K. H. Nealson and R. Popa

Photosynthesis and the generation of oxidants

C0 2 + r donor

(CH 2 0) + oxidized e~ donor

ORGANISMS

Anoxygenic photosynthesis
Purple

Green filamentous
Heliobacter
Green sulfur

Oxygenic photosythesis
Cyanobacteria
Chloroplasts
(plants and algae)

e- donors

H 2

oxidized e~ donors

So, S0 4 2 ",
Fe3 +

Fig. 3. Generation of oxidants during photosynthesis. The general formula for photosynthesis is
shown at the top of the diagram, with the critical difference being the molecules used as electron
donors (and the resultant oxidized forms produced).

Magnetite

Vivianite (Fe.O.)

(Fe 3 (P0 4 ) 2 8H 2 0) j

p ttf e '&B? egg*

Sidcritc
(FcC0 3 )

Lactate

Acetate + CO

Fig. 4. Accidental biominerals. The redox activity of some bacteria (such as iron reducers) results in
the formation of various minerals (vivianite, magnetite, siderite), depending primarily on the chemistry
of the environment.

SGM symposium 65

Metabolic diversity and exobiology 1 57

donor (Blankenship, 2002; Xiong et al., 2000). Such anoxygenic photosynthetic bacteria
produce oxidized electron donors as a by-product of photosynthesis (Fig. 3) and were
the precursors of the now widespread oxygenic photosynthesis (Blankenship, 2002).
The more complex oxygenic photosynthesis is thought to have arisen in the cyano-
bacteria and eventually appeared symbiotically in algae and plants (Margulis, 1981),
one of the most important events in the evolution of complex life.

With the above discussion of metabolic diversity, it is easy to miss an important point:
namely that there is also an impressive unity to the metabolism of life. Virtually all
present-day metabolism on Earth operates in a similar way, with environmental redox
equivalents being harvested and used for the reduction of cellular electron or hydrogen
carriers which are similar throughout all of life (Nealson, 1997; Nealson & Conrad,
1999). These reduced carriers are then used directly for biosynthesis or for the gener-
ation of a membrane potential (combination of a pH and electrical potential) that can
be used for a variety of functions (including ATP formation, transport and motility).
Thus, despite apparent rampant diversity in terms of resource allocation, the central
energy-processing systems of all life, including those of structurally complex large
eukaryotes, operate in a similar way. With regard to this issue, there are really only
two metabolic means of extracting energy from the environment: chemiosmosis and
fermentation. Almost without exception, the interactions between the living world
and the mineral world are of the chemiosmotic type (i.e. redox chemistry). While it
is difficult to specify exactly when the ability to respire arose in time, one imagines
that it is indeed one of the earliest 'inventions' of life (Nealson 6c Rye, 2004) and, once
invented, became a component of virtually all subsequent successful experiments in
metabolic evolution.

MINERAL FORMATION: THE GEOBIOLOGY CONNECTION
Prokaryotic mineral formation

As discussed above, the array of electron donors and acceptors that the prokaryotic
world uses is, to some extent, the glue that knits microbial metabolism to geobiology.
As shown in Table 1, many of the redox-active compounds become a solid (mineral or
mineraloid) form - thus, oxidation or reduction of these components often results in a
change of state of the component (i.e. between insoluble and soluble), leading to
formation or dissolution of minerals. In this sense, the metabolism of the prokaryotes,
designed for the purpose of harvesting energy from the environment, is inadvertently
linked to the formation and/or dissolution of many minerals. The inadvertent nature of
this process is shown in Fig. 4, a summary of results in which a culture of Shewanella
putrefaciens CN-32, during the process of iron reduction, was shown to produce any of
several mineral products depending on the environment in which the bacteria are grown

SGM symposium 65

158 K. H. Nealson and R. Popa

Vesicles Organic Nucleation

(cisternae) scales of calcite

Growth of calcite
into coc eoliths

Calcite crystallites Radial ' Kxocytosis

(heterococcoliths) organization of of coccoliths

heterococcoliths

Coccoliths with
organic coating

Fig. 5. Purposeful biominerals. Biomineralization of coccolithophores (calcareous algae) making
crystals of calcite. This diagram is intended to demonstrate the complex and detailed interactions that
occur in general in the formation of mineral products by eukaryotes (after Young & Henriksen, 2003).

(Roden & Zachara, 1996). That is, while the bacteria metabolically impact the local
mineralogy, the end products produced are controlled by the environment rather
than by the bacteria themselves. Thus, to some extent, one can view the formation of
minerals by some (perhaps most) of the prokaryotes as metabolic 'accidents', depend-
ent on the metabolic chemistry, but not directed by it. As far as is known, there is
no direct advantage to the bacterium of forming a given mineral in comparison to
any other: this is determined in fact by the chemistry of the environment in which the
bacteria are living. This leaves us in the interesting position of viewing prokaryotic
minerals as a metabolic by-product - something done 'accidentally', and something,
perhaps, that occurs as an indicator of microbial metabolism, but without the morpho-
logical clues usually associated with biominerals or fossils.

This being said, while the formation of a specific mineral may have no advantage for
prokaryotes, it may well be that mineral formation in general serves a useful thermo-
dynamic/kinetic purpose. From the point of view of the bacterium, removing one of the
soluble products as an insoluble mineral could have a dynamic effect on the chemical
abilities of the bacterium. This is a strategy often employed by bacteria living in sym-
biosis with other bacteria, in which the end product of one is the substrate for another,
and this cascade makes chemical reactions proceed at speeds far beyond those predicted

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Metabolic diversity and exobiology 1 59

if the reaction products were disposed of by diffusion alone. Perhaps the most pertinent
example is that of anaerobic methane oxidation, in which the process is really driven by
sulfate-reducing bacteria that consume the hydrogen produced via methane oxidation,
pushing the reaction forward (Boetius et al., 2000; Orphan et al., 2001a, b).

Eukaryotic mineral formation

The eukaryotes, in contrast, have perfected the art of mineral fabrication, making a
wide array of biominerals with many different uses (Dove et al., 2003; Lowenstam,
1981; Lowenstam & Weiner, 1989). Biomineralization is extremely common among the
multicellular eukaryotes, many of which need structural elements to grow and function
in three dimensions, providing advantages for predation (teeth, bones etc.) and protec-
tion (shells, frustules, cell coverings) (Lowenstam, 1981; Lowenstam & Weiner, 1989).
These biominerals are often ornate and recognizable structures that provide us with a
visible fossil record, allowing the calibration of the molecular record (discussed below),
something that is extremely difficult prior to the 'invention' of these biominerals. While
it is less widely distributed among the eukaryotic microbes, organisms like coccolitho-
phores and foraminifera (organisms that make carbonate minerals) and diatoms and
radiolarians (organisms that make silica minerals) have dramatic impacts on the carbon
and silicon cycles on the planet (Lowenstam & Weiner, 1989; Skinner & Jahren, 2004).

It is clear from emerging mechanistic studies that the situation is dramatically different
with regard to eukaryotic mineral formation (Dove et al., 2003). Eukaryotic bio-
minerals are genetically directed structures, formed using protein templates that are
used to catalyse and direct the synthesis of the specific minerals. The resulting struct-
ures are reproducible enough that they can often be used to identify the organism that
produced them, often to the level of genus or even species (Fig. 5). This is in marked
contrast to prokaryotic mineral formation and dissolution, which is primarily a func-
tion of environment rather than genetics (Fig. 3). As opposed to the prokaryotes, which
form minerals while gaining metabolic energy, the formation of eukaryotic biominerals
is 'costly' in the sense of requiring specific templates, energy for synthesis and often
auxiliary systems for transport and assembly. Finally, redox chemistry is not a funda-
mental part of eukaryotic mineral formation, as redox changes do not occur and no
energy is gained. Given that the eukaryotes are unable to view minerals as either viable
electron donors or acceptors, their ability to form or dissolve redox-active minerals is
extremely limited.

A final note with regard to the minerals produced by prokaryotes and eukaryotes:
because protein templates and other structural aids are not involved in prokaryotic
mineral formation, these minerals are often 'pure', sulfides, carbonates etc., while those
of eukaryotes are 'hybrids', composed of biological material interspersed with

SGM symposium 65

160 K. H. Nealson and R. Popa

Fig. 6. Chains of magnetosomes in Magnetospirillum magnetotacticum AMB1 as seen in a high-
magnification TEM section (image provided courtesy of Dr Virginia Souza-Egipsy). Bar 100 nm.

crystalline minerals (Dove et ai, 2003). These adaptations add remarkable mechanical
properties to the eukaryotic biominerals and make them chemically distinct from
the prokaryotic biominerals (Dove et uL, 2003). This being said, it should also be noted
that a wide variety of mineral and mineraloid-like inclusions (oxalates, carbonates, uric
acid etc.) are neither species-specific nor 'hybrid' in nature.

Biomagnetite and the death of a generalization

The above discussion of mineral formation divides the world into prokaryotic redox
minerals and eukaryotic structural and functional biominerals. For example, dissimi-
latory iron-reducing bacteria can produce copious amounts of extracellular magnetite
as a function of respiration (Lovley 6c Phillips, 1988; Roden & Zachara, 1996), while
eukaryotic chitons can produce magnetite coatings on their teeth (Lowenstam, 1981;
Lowenstam & Weiner, 1989). The eukaryotic magnetite is easily recognizable, but
whether it is possible to distinguish between magnetite formed by prokaryotes and
abiotically formed magnetite is not yet clear.

The division between prokaryotes and eukaryotes with regard to biominerals is
not always so clear, however. Many strains of magnetotactic bacteria (no magneto-
tactic archaea are known) produce highly ordered, intracellular, crystalline magnetite
inclusions called magnetosomes (Fig. 6). These single domain, mineralogically nearly

SGM symposium 65

Metabolic diversity and exobiology 161

perfect, highly magnetic crystals are often arranged in chain-like arrays (Blakemore,
1975; Bazylinski & Frankel, 2000). They provide the cells that contain them with the
ability to align passively in a magnetic field, the response being that these highly motile
cells move swiftly towards one magnetic pole or the other. It appears that each bacterial
strain produces a characteristic magnetite pattern (size, shape and arrangement of the
magnetosomes) and that specific genes (and, thus, specific proteins) are involved in
the production of the magnetosomes (Schuler, 1999, 2004; Matsunaga & Okamura,
2003). Furthermore, while the function(s) of the magnetosome is not yet fully proven,
nearly everyone agrees that there must be an advantage to the cells that contain the
magnetosomes, and most agree that it is connected with environmental sensing and
location (Frankel etaL, 1997).

Are there other such examples of highly structured, genetically dictated biominerals
produced by bacteria? Recently, it was reported that highly crystalline forms of metal
sulfides were produced by sulfate-reducing bacteria (Suzuki et al., 2002). Whether these
are in fact structured with a 'purpose' is not known, but their highly crystalline
structure is an intriguing finding; perhaps such preordained and structured prokaryotic
minerals are more prevalent than was thought (Fortin, 2004).

TIME: LINKING THE PAST WITH THE PRESENT

The timing of the emergence of various metabolic abilities is one of the most enigmatic
problems in the microbial world. The evidence available to us in the ancient rock record
is, at best, akin to smudged fingerprints and, at worst, like footprints in the sand after
the tide has come and gone. One of the great hopes of molecular phylogenetic
approaches was that they would allow one to look back in time using sequence data:
using these data to estimate when major metabolic 'inventions' occurred in the past.
Such 'inventions' would include not only structural innovations visible in the fossil
record, but metabolic inventions like respiration and photosynthesis and other
prokaryotic specialities like nitrogen fixation, denitrification and sulfur oxidation. This
hope has been realized to some extent but, realistically, the methods are probably good
only as far back as there is a fossil record to support them. For example, Benner et al.
(2002) discussed the use of various types of data to understand the evolution of
metabolism, concluding that, within the last 50, perhaps, 100 million years, one can
probably do this with some confidence, utilizing a combination of fossil, isotope,
organic geochemical and molecular evolutionary clock records to infer past patterns of
evolution (in this case, the evolution of certain sugar-fermentation abilities). Thus, to
some extent, the fossil record is of limited use: true (recognizable) biominerals
produced by eukaryotes are visible only ~500 million years ago, when the first sponge
spicules and carbonate biominerals can be seen. These processes are thus geologically
young (Li etaL, 1998; Nealson & Rye, 2004).

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162 K. H. Nealson and R. Popa

Prior to the formation of these 'true' biominerals, the signatures that exist are primarily
geochemical in nature. Thus, while prokaryotic minerals may not be readily recog-
nizable by their morphologies or unique crystal structures, many can be judged to be
of biological origin via the fractionation of isotopes during the 'supply' of chemical
components for their formation. Kinetic fractionation occurs as a function of enzymic
catalysis, leading to biological materials preferentially being composed of the 'lighter'
isotopes. Thus light carbon is preferentially used by living organisms during carbon
fixation and accumulates in the resulting biomass, while light sulfide is produced during
sulfur or sulfate reduction and accumulates in sulfide minerals. These isotopic tracers
have been of value in tracing the metabolic activities of modern organisms in both the
laboratory and the field and provide a major tool for looking for indicators of
metabolic activity in ancient samples. That is, it is possible, by using C and S isotopes,
to see in the ancient rock record the appearance of processes that result in fractionation
of these isotopes. Herein lies an important distinction that can be easily missed: while
the isotopes may strongly suggest the existence of a process leading to fractionation,
they cannot tell us unambiguously which process was involved and cannot be used to
tell which organism or even group of organisms accounted for the fractionation. Unless
one accepts this tenet, one can be easily fooled.

Carbon isotopes have been among the most valuable and widely used of the isotopes
with regard to life detection and definition. Carbon fractionation occurs during its
reduction (fixation) from C0 9 to organic carbon by bacteria, algae and plants and to a
much greater degree (i.e. light carbon) when C0 2 is fixed into methane by methano-
genic archaea. However, even for this well-known and often-used system, deciphering
the isotope signatures from ancient samples is difficult because (i) these pathways are
varied and unknown, (ii) subsequent diagenetic reactions are not easy to specify and
(iii) the signature of the source of carbon is seldom known.

Even with these caveats, however, it seems clear from a variety of studies that bio-
minerals can be traced to the early phases of Earth's history. Stable-isotope signatures of
both carbon and sulfur suggest that metabolic activities were involved with the
formation of minerals from very early times. Carbon isotope ratios ( lo C/ 12 C) have been
used to suggest that carbon fixation may have existed as early as 3-8 Ga (billion years)
ago (Mojzsis et al., 1996). While this number has been challenged, few would argue
with 3-5 Ga for convincing evidence of carbon isotope signals in the ancient record.
Similarly, sulfur isotopes ( 34 S/ 33 S) suggest that sulfur reduction of some kind was
occurring 2-5 Ga and perhaps earlier (Canfield et ai, 2000; Shen et aL, 2001).

Providing definitive fossil evidence from before the time of the biomineral-producing
eukaryotes is difficult for several reasons: first, because the preservation of the materials

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Metabolic diversity and exobiology 163

is often poor, making identification difficult, and second, because virtually none of the
putative organisms seen in the samples are alive today. While they have similarities to
other organisms, the nature of their behaviour and even their metabolism cannot be
specified with reasonable certainty. Another discouraging development with regard
to molecular evolution methods is the rampant appearance of examples in which it is
now clear that the evolutionary 'clock' is neither constant (it can run at different rates)
nor predictable (Doolittle et ai, 1996). Thus, two organisms that would be described
as deeply branching and suspected of being of equal 'age' may be quite different
because of differences in evolutionary clock speeds. For ancient samples, of the order
of hundreds to thousands of millions of years and older, the situation gets even more
uncertain.

Another difficulty that is peculiar to the prokaryotes is that the 'fossils' used to identify
them are reduced to either organic geochemicals (i.e. classes of chemicals peculiar to a
certain type of cell) or isotope fractionation patterns indicative of a certain type of
metabolism. Both analyses have their limitations. For example, while some classes
of compounds can be identified as components of cyanobacteria in the modern world,
there is no way of knowing with certainty that non-cyanobacterial organisms that
contained these compounds were not present before the cyanobacteria arose. Similarly,
when one sees isotope fractionation, such as the appearance of light sulfur, it is temp-
ting to invoke the appearance of sulfate reduction (and sulfate-reducing bacteria) and,
in fact, this is often done (Canfield et al., 2000). While the activity of sulfate-reducing
bacteria would indeed explain the observed results, it is also true that sulfur can be
fractionated by many other organisms, including sulfur-, polysulfide- and thiosulfate-
reducers (Smock etal., 1998).

As a footnote to this discussion, one must remember that much of the work with both
organic biomarkers and isotope fractionations hinges on our knowledge of microbial
physiology and the composition of cultured microbes. As noted above, it is estimated
that less than 1 % of the microbial world has been obtained in culture (Whitman et al.,
1998), and therefore that we are often extrapolating from a very limited experimental
base. To this end, we note in Table 2 that much progress has been made in recent years
in uncovering novel organisms and novel metabolic abilities - microbes whose very
existence was previously doubted or unknown. Given that methods are improving with
regard to culturing microbes, one can hope that the situation will improve, but one
must be cautious when trying to extrapolate backwards in time to ancient metabolisms
based on such an incomplete database.

Finally, we must confront the issue of lateral (horizontal) gene transfer (Doolittle,
2002). With the advent of genome sequencing, it has become apparent that in the

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164 K. H. Nealson and R. Popa

Table 2. Microbial metabolic diversity

This table represents some of the metabolic abilities of prokaryotes that have been revealed since
1988.

Process

Found in

Year of discovery Reference(s)

Dissimilatory Fe and Mn
reduction

Anaerobic methane
oxidation consortium

Anaerobic ammonia
oxidation

Bacteria and archaea

Bacteria/archaea

Bacteria

1988

2000

1998

Lovley & Phillips (1 988);

Myers & Nealson (1988)

Boetiusefa/. (2000);
Orphan et al. (2001b)

Jettenefa/. (1998)

Anaerobic iron oxidation

Photosynthetic

bacteria

1993

Ehrenreich&Widdel(1994);
Widdelefa/. (1993)

Anaerobic iron oxidation

Heterotrophic I

Dacteria

1994

Straubefa/. (1996)

Perchlorate reduction

Bacteria

1996

Coatesefa/. (1999)

Phosphite oxidation

Bacteria

2000

Schink & Friedrich (2000)

Dissimilatory arsenic
reduction

Bacteria

1994

Ahmannefa/. (1994)

Dissimilatory selenate
reduction

Bacteria

1994

Oremland etal. (1994)

past there has been a large amount of mixing of genomes, such that the so-called
phylogenetic 'tree of life' is in fact much more like a cross-hatched bush (Doolittle,
2002). Each of the three 'kingdoms' has obtained, and fixed into their genomes,
ample information from the other two kingdoms, suggesting that major sharing of
genetic information has occurred, especially among the prokaryotes, where this is still
happening at rapid rates. Thus, one of the major issues in prokaryotic genomics is that
of defining the gene set that characterizes a given group of microbes (as opposed to
those genes that can apparently move in and out with non-lethal results).

With regard to the use of molecular methods, one of the intellectual 'traps' of this
approach should be noted - namely that, for the prokaryotes, all of the organisms on
which the 'phylogeny' is based are alive and evolving today. That is, while they may
contain 'ancient' traits or abilities, these are surely not as they were in the past, so that
the phylogenetic approach allows one to look at the most likely sequence of events, but
not to place accurate times on any of the events. Thus, one could argue that we are
looking at the living remnants of past microbial evolution and can have a reasonably
good idea of what preceded what (within the limitations of uncertainties introduced
by horizontal gene transfer), but trying to ascertain when any of these processes arose,
i.e. when a given process or organism first appeared, is very difficult (if not impossible)
by this method. As a precaution, one might note that almost none of the organisms
seen in the fossil record of 100 million years ago are alive today If we tried to construct

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Metabolic diversity and exobiology 165

those organisms simply from molecular biology alone, it would almost certainly be a
resounding failure. This is the dilemma we are faced with in reconstructing prokaryotic
evolution.

Kinetic biosignatures and layered communities

While the above discussion has dealt with inference of the past through reading the rock
record, much biology is concerned with measurement and manipulation of the present.
To this end, one can look for biosignatures connected to the chemistry of the organism
and its impact on the environment (substrate removal or alteration and product
deposition). Such short-term interactions define, to some extent, one of the major
differences between life and non-life: specific catalysis of reactions that would other-
wise occur at extremely slow rates. In the absence of life, many low-temperature
geochemical reactions (i.e. less than 100 °C) proceed at rates slower than molecular
diffusion, so that product accumulation and gradient formation cannot occur. Enzymic
catalysis, however, leads to the consumption of reactants at rates faster than they can
be supplied and the production of products at rates faster than they can diffuse
away, leading to the formation of gradients that are indicative of the life forces that
have produced them (Nealson & Berelson, 2003). The simple argument is that, in the
absence of life, the chemical gradients (as we call them, layered microbial communities
or LMCs) that are so dominant in anaerobic niches on Earth would simply not exist.
For example, in the Black Sea (Fig. 7), a series of redox zones are seen, each indicative
of a process that occurs in a well-defined redox zone (Nealson & Berelson, 2003). At
these interfaces or layers, the chemical profiles can be used to define the microbial
processes that are occurring to establish the gradients. For example, as shown in Fig. 7,
the abrupt disappearance of oxygen at -50 m is referred to as the oxygen-depletion
zone, where catalysis by aerobic respiration occurs so quickly that oxygen is taken to
nearly zero within a few metres. In the absence of respiration, oxygen would be nearly
constant to the bottom. Similarly, the nitrate disappears a few metres below due to a
process called denitrification - without this, the nitrate would almost certainly be
uniformly distributed in the Black Sea, as the rates of the processes that might consume
it are very slow. Below this are the zones of manganese, iron and sulfate reduction, all
of which would not exist without life catalysing each process.

It should be obvious that these LMCs are inferred not from biological measurements,
but from the very existence of the chemical layers and non-linear gradients (Nealson &
Berelson, 2003). As will be discussed below, they are more complex than is implied
above, but the gradients themselves can be used as indicators of specific catalysis and
thus of life. Of interest here is that, with regard to spatial scales, we see such LMCs at
scales of micrometres in biofilms to millimetres in algal mats and centimetres in lake
and ocean sediments - they are a universal feature of life on Earth. In environments

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166 K. H. Nealson and R. Popa

Black Sea (marine basin)
Maximum (%)
50

100

100 r

150 -

200 -

NO:

2 +

NH +

Fig. 7. Redox-related vertical profiles in the Black Sea. This profile presents a general scheme of redox
interfaces seen in the Black Sea (Nealson & Berelson, 2003). The gradients here are a function of the
organisms that accumulate at the interfaces - organisms that consume reactants and create products
at rates sufficient to account for the gradients observed. Similar chemical gradients are used as
indicators of LMCs over scales ranging from many metres, like this one, to millimetres or less in many
other environments.

where rapid mixing occurs, such as the open ocean, mixed lakes and the atmosphere,
such gradients are rapidly dissipated by convection and mixing, but the signals of the
biological processes are nevertheless there.

EXTREME ENVIRONMENTS: WHY DO WE GO THERE?

'Extremophiles': one of the most popular buzzwords of the last decade, and the cause
for great excitement among microbiologists. Finding the 'most extreme' organism with
regard to any given variable (temperature, pH, UV radiation etc.) has been a sporting
pleasure for many of us. However, our obsession with these physical and chemi-
cal extremes may be, to some extent, slightly misplaced. For many physical and
chemical variables, the prokaryotes and eukaryotes are both able to adapt reasonably

SGM symposium 65

Metabolic diversity and exobiology 167

= Eukaryotes

Physical-chemical extremes □ ■ Prokaryotes

| -20 Temperature +121

L°

pH +11-5

dw Salinity saturate^

Nutritional extremes

H 2 , H 2 S -- Inorganic energy sources- CO, Fe(ll)

Q 2 , NO3-, Mn(IV) - Oxidants - S0 4 2 ~, CQ 2 |

Fig. 8. Types of extremophily. This diagram shows the limits of prokaryotic and eukaryotic life for a
number of variables. The top three panels are physical/chemical variables, demonstrating that the
structurally simpler prokaryotes are a bit tougher and more resilient than eukaryotes, while the bottom
three panels show that some types of metabolism are 'off limits' for the eukaryotes - truly extreme,
dw, Distilled water.

well (Fig. 8), although the simpler and more robust prokaryotes almost always exceed
the more complex eukaryotes. As seen in Fig. 8, however, the true extreme conditions
might be imagined to be those of nutrition. As noted above, eukaryotes might define as
'acceptable' only those nutrients that can be converted into glucose or pyruvate, and
only oxygen to respire. Thus, nearly the entire inorganic world of electron donors is
'extreme', as is the use of any electron acceptor other than oxygen. When we move into
the anaerobic world, especially the anaerobic world at high temperature, we begin
to see the prokaryote-dominated world that we can imagine has some relevance to the
ancient Earth. It should not be surprising to see that many of us who want to think
about either the ancient Earth or the possibility of life elsewhere find ourselves studying
environments like the deep subsurface - places where nutritional extremophiles are
abundant and, for the most part, uncharacterized.

CONCLUSIONS

Metabolic diversity has become a catchphrase for the wonders of the microbial world -
wonders that may help to explain why the world is like it is through the energetic
exploitation of the geosphere via the metabolism of the biosphere. A direct, but not
entirely obvious, extension of this thinking is that it applies not only to understanding
the evolution and distribution of life on our own planet, but to the possibility of life in

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168 K. H. Nealson and R. Popa

abodes other than Earth: giving rise to such questions as, 'could we use this geobiolo-
gical logic to distinguish a living from a dead planet?' A quick perusal of the metabolic
diversity of the prokaryotic life on our planet reveals that virtually every redox-related
energetic niche is occupied. If there is fuel (an electron donor) and something with
which to burn the fuel (an electron acceptor), then some microbe has learned how to
exploit this energetic niche. This exploitation, in its simplest sense, defines the meta-
bolic diversity with which we work - microbes striving to make a living on whatever
resources are available, and being remarkably successful in plying their trade. Given
that such metabolic diversity is so dominant on our own planet, one imagines that we
should expect no less of any other planet on which life has evolved. If so, then the search
for life in all realms (the surface and subsurface of the Earth, as well as potential
extraterrestrial sites and samples) should begin with considerations of energy types,
levels and fluxes, always with an eye to finding things that should not be there in the
absence of life - the true connection of microbial diversity to exobiology. We don't
know whether life exists elsewhere than on our planet. However, given the remarkable
ability of life to adapt to and exploit earthly energy sources, it is an easy intellectual
leap to imagine that, in other abodes, with abundant energy and acceptable chemistry,
some sort of life could prosper. However, assumptions about the chemistry or other
details of any non-earthly life may lead one awry. What must be true is that life will
evolve to take advantage of the energy sources and chemical building blocks that are
available and that, to compete with the natural chemistry, it will need to invent catalysts
that elevate the rates of reactions. It is this catalysis that should be the signpost of life -
the geobiological indicator that something is amiss (reactant consumption, product
accumulation, complex molecule generation or isotope fractionation), and the enticing
possibility that life may have been responsible.

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coccoliths. In Biomineralization, Reviews in Mineralogy and Geochemistry, vol. 54,

pp. 190-215. Edited by P. M. Dove, J. J. De Yoreo & S. Weiner. Washington, DC:

Mineralogical Society of America.
Zumft, W. G. (1997). Cell biology and molecular basis of denitrification. Microbiol Mol Biol

Rev 61, 533-616.

SGM symposium 65

Biogeochemical cycling in polar,
temperate and tropical coastal
zones: similarities and differences

David B. Nedwell

Department of Biological Sciences, University of Essex, Colchester C04 3SQ, UK

INTRODUCTION

This chapter will consider biogeochemical cycling in the coastal zone. This is defined as
that area of estuarine and coastal, relatively shallow water where there is strong
benthic-pelagic linkage and exchange between the water column and the underlying
sediment. In deeper water this connection becomes increasingly tenuous as the
exchange between the euphotic zone and the benthic layer declines. Longhurst et al.
(1995) recognized the coastal boundary domain, divided into 22 provinces, as often
bounded by a shelf-break front, and included coastal upwelling regions. The coastal
zone generally exhibits high rates of primary production compared with the open
ocean (Table 1), and there is the greatest impact from inputs from the land to the coastal
sea through estuaries. Estuaries and coastal seas are highly heterotrophic systems which
are net exporters of C0 2 to the atmosphere due to the mineralization and recycling
of both autochthonous and allochthonous organic matter (Borges, 2005).

PHYSICO-CHEMICAL DIFFERENCES BETWEEN LATITUDINAL
REGIONS

The physical-biological interactions that influence marine phytoplankton production
have been reviewed by Daly &C Smith (1993). Because of the spherical shape of the
Earth, more solar energy falls per unit area of surface in equatorial regions than at
the poles (Fig. la), and the incidence of light at the equator is vertical to the surface, but
oblique at the poles. Furthermore, the distance radiation travels through the atmos-
phere is longer at the poles, thus reducing the irradiation incident at the poles compared
with equatorial regions. Consequently, equatorial regions have higher energy inputs

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

174 D. B. Nedwell

Table 1. Areas and annual primary production rates of biogeochemical provinces
based on the CZCS 1978-86 climatological data (after Longhurst et a/., 1995)

Geographical subset

Area (xKT 6 km 2 )

,-v

Primary production rate (gCm 2 year n )

Coastal domain

37-4

Polar domain

20-8

Westerlies domain

129-9

Trades domain

139-9

Arctic Ocean

Atlantic Ocean

74-0

Pacific Ocean

148-9

Indian Ocean

45-4

Southern Ocean

57-9

Upwelling provinces

8-4

385
310
126
93
645
199
132
143
141
398

Table 2. Seasonal variation in characteristics of the coastal zone in polar,
temperate and tropical coastal zones

Characteristic

Polar

Temperate

Tropical

Insolation
Day length
Water temperature
Estuarine flushing

Extremely seasonal

Constant, low

Highly seasonal
(summer melt)

Moderately seasonal
Moderately seasonal
Seasonally variable
Moderately seasonal

Weakly seasonal

Constant, high

Highly seasonal
(wet/dry season rainfall)

and are hotter than polar regions. Rainfall and humidity are also extreme in the tropics,
often with distinct rainy seasons which impose differences in the ecology of the coastal
zones. During the tropical dry season, when there is little or no rainfall or river flow,
there may be accumulation of detritus within an estuary but during the wet season
heavy rainfall flushes the estuary, exporting a spike of material to the coastal zone, and
're-setting' the estuary (Eyre & Balls, 1999). For example, north Queensland tropical
estuaries exhibited flushing times varying between 2 and 30 days during the dry season
but days in the wet, when river water essentially passed straight through the estuaries.
In contrast, temperate Scottish estuaries exhibited much less seasonal variation in
flushing. Pulses of nutrients (and particulate matter) leached from the catchment are
therefore likely in tropical estuaries compared with temperate estuaries, whose rainfall
and therefore input loads to the coastal zone are less seasonally variable. Estuaries
are extremely important in attenuating the loads of nutrients passing from land to sea
(see Nedwell et al., 1999); in the case of nitrogen particularly by denitrification, or of
phosphorus by settlement of particles to which phosphate is adsorbed. Attenuation
of nutrient fluxes from land to sea is likely to be diminished if the major loads coincide
with periods when freshwater flushing times within the estuary are very short, as in

SGM symposium 65

Biogeochemical cycling in the coastal zone 175

(a)

(b)

20 r

A I 1 1 I I I J I I L. 1 i J

JFMAMJ JASOND

Fig. 1. (a) Owing to the spherical nature of the Earth, solar energy is concentrated over a smaller area
in equatorial regions than in polar regions, and the distance travelled through the atmosphere is also
shorter at the equator. Consequently, energy input per unit area is greater at the equator than at the
poles, (b) Variations in day length throughout the year in relation to latitude (reproduced with
permission from Osborne, 2000).

some tropical estuaries. In polar regions any terrestrial input to the coastal zone occurs
only during the short summer, when ice and snow melt occurs, and is therefore also
seasonally extreme.

Day length also exhibits latitudinal changes. Because of the rotation of the Earth
around its axis there is little difference seasonally in the day length at the equator,
whereas at the poles there is extreme change in the day length between winter and
summer and intermediate seasonal changes in day length at temperate latitudes
(Fig. lb). The different environmental factors are summarized in Table 2. In essence,
high-latitude polar coastal waters are stenothermal, with relatively constant low
temperature (-1 to 4 °C), mid-latitude temperate waters are eurythermal, with seasonal
temperature cycling between approximately 2 and 20 °C, and low-latitude tropical
waters are stenothermal, with relatively constant high temperature (> 16 °C).

PRIMARY PRODUCTION

Biogeochemical cycling in the coastal zone is driven by autochthonous primary
production and also by allochthonous inputs via estuaries, both of which provide the
energy for subsequent chemo-organotrophic activity. To compare biogeochemical
cycling in different latitudinal regions we must start by examining their primary

SGM symposium 65

176 D. B. Nedwell

75
88
103
120
141
165
193
226
265
| 310
363
426
498
583
683
800

Fig. 2. Global primary production (g C rrr 2 year 1 ) modelled from satellite estimates of chlorophyll a,
photosynthesis-light relationships and local light environment (reproduced with permission from
Longhursteta/., 1995).

production. Are there consistent differences in primary production in polar, temperate
or tropical coastal zones? Longhurst et al. (1995) used satellite radiometer data to
model primary production in the oceans (Fig. 2), based on estimates of chlorophyll by
the CZCS radiometer on the NIMBUS satellite from 1979 to 1986, photosynthesis-light
relationships and the light environment. It is clear that the coastal zones are in general
areas of high production. They estimated global oceanic primary production as 44-7-
50-2 Pg C year . The annual primary production rates for the coastal domain (Table 2)
averaged 385 g C m - ^year -1 , for the Arctic Ocean 645 g C m - year -1 and upwelling
provinces 398 g C m -2 year -1 . The annual production in the coastal domain was greater
than that in all oceans except the Arctic Ocean, and the relative productivity of the
ocean province was only 2-5-4-0 times that of the coastal provinces, despite the former's
very much greater area. Tropical coastal provinces in the Indian and Pacific Oceans
averaged 279 g Cm -2 year -1 , again confirming that low-temperature polar coastal
communities are as productive as those at lower latitudes and higher temperatures.

What is beyond doubt is that the seasonal pattern of primary production in the
different regions varies enormously, reflecting the different seasonality imposed by
insolation and day length, and in turn influencing the biogeochemical cycling that goes
on. Polar regions have constantly low water temperatures (usually <4 °C) but exhibit an
extreme seasonal pattern, with short summer ice-free periods of high insolation during
which pelagic primary production rapidly increases, plankton biomass quickly attains
a maximum and a pulse of organic input to the coastal zone bottom sediments
occurs (Fig. 3) (Nedwell et al, 1993; Rysgaard et al, 1999). Rysgaard et al (1999)
showed that there was a strong relationship in Arctic waters between the annual pelagic

SGM symposium 65

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Biogeochemical cycling in the coastal zone 177

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Fig. 3. Seasonal changes in the settlement rate of organic matter from the water column to coastal

sediments at Signy Is., South Orkney Islands, Antarctica (reproduced with permission from Nedwell
etal., 1993).

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Fig. 4. Annual primary production (•) and grazing by ciliates (O), heterotrophic dinoflagellates (A)
and copepods (O) in Young Sound, northeast Greenland (reproduced with permission from Rysgaard
etal., 1999).

primary production and the duration of the open-water period when pelagic primary
production could occur. This is because an ice layer with snow cover strongly attenuates
light, and it is only after ice melt that there is significant illumination of the water. The

SGM symposium 65

178 D. B. Nedwell

Sea ice

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(reproduced with permission from Nedwell etal., 1993).

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Fig. 6. Benthic anaerobic organic matter degradation driven by sulfate reduction as a percentage
of total mineralization indicated by benthic 2 uptake (reproduced with permission from Nedwell
etal. ,1993).

seasonal cycle of pelagic primary production in Young Sound, Greenland, exhibited an
extremely short but intense period of primary production (Fig. 4) and in the coastal
zone around Signy Island, in the South Orkney islands, at the tip of the Antarctic
Peninsula, a similar short but intense period of summer production has been described

SGM symposium 65

Biogeochemical cycling in the coastal zone 179

(Nedwell et aL, 1993). Once sea ice breaks up, the rapidly increasing light penetrates the
clear water and primary production increases exponentially, the water column seeded
by ice algae from the bottom surface of the melting ice. The complete seasonal cycle
of algal production at Signy Is. occurs within 2-3 months, after which the sea ice forms
again, and may be as short as 4-8 weeks at higher latitudes (e.g. Rysgaard et al., 1998;
Karl et al., 1996). Lack of turbulence beneath fast ice causes rapid settlement of organic
matter from the water column and there is, in essence, a short but intense pulse of
organic deposition into the coastal bottom sediments. Benthic processes such as 0-,
uptake respond quickly and increase due to the input of fresh organic matter, so that
there is a strong seasonal signal of benthic activity, in direct contrast to tropical coastal
systems. In the few studies carried out, annual benthic organic degradation seems
broadly to balance net primary production in both Arctic and Antarctic coastal systems
studied so far (Nedwell et al., 1993; Rysgaard et al., 1998), indicating stable ecosystems.
However, while primary production occurs in a short summer burst, that pulsed input
of organic matter sustains benthic processes for the remainder of the year, spread
out over the period when winter ice cover minimizes further organic inputs. Thus, the
deposited organic matter represents a food or energy reserve for the organotrophic
benthic community which tides them over the winter period with no new inputs. For
example, at Signy Is. the seasonal signal of deposited organic matter could be seen only
in the top 0-5 cm of sediment (Fig. 5). As the deposited organic matter was buried into
the sediment by very active bioturbation, the proportion of benthic organic matter
degradation which was driven by anaerobic sulfate reduction increased through
the winter period, while that due to aerobic metabolism declined (Fig. 6) until, with the
input of fresh organic matter in spring onto the surface of the sediment, aerobic
metabolism again increased dramatically and the cycle repeated. Mincks et al. (2005)
have argued that this pulsed input to polar benthic communities represents a 'food
bank' for polar benthic detritivors, including micro-organisms, over the winter period
when there were no new organic inputs, and the data from Signy Is. are consistent with
this idea.

In the temperate coastal zone annual primary production still exhibits a strong season-
al signal, though less intense than that of polar regions. There are significant seasonal
changes in insolation and water temperature (e.g. approximately 3-18 °C in the
southern North Sea) and a seasonal cycle of spring bloom primary production occurs,
subsequently limited by depleted nutrient availability, and possibly with a smaller
autumn bloom if summer stratification breaks down. In estuaries and inshore waters,
where there is a continuous supply of nutrients, there may be no distinct spring bloom,
but an increase of primary production during spring which is sustained throughout the
summer (e.g. Kocum et al., 2002a, b). For example, in the north-central North Sea
during 1998-99 there was a spring bloom which decreased during summer, but, in

SGM symposium 65

180 D. B. Nedwell

the German Bight, where there is a continued supply of nutrients from the Rhine,
primary production continued in a broad peak throughout the growing season (Fig. 7).
Again, there are seasonal changes in the downward fluxes of organic matter to bottom
sediments, albeit less extreme than in polar regions. Upton et al. (1993) estimated
that benthic organic matter mineralization accounted for 17-45 % of net primary
production in the various regions of the southern North Sea.

In the tropics the seasonal insolation change is least and insolation remains at high
values throughout the year, while water temperatures also remain relatively constant
and high. Primary production is therefore less likely to be limited by insolation than at
higher latitudes, but regulated by other factors such as nutrients or turbidity. The
tropical regime typically reflects the relatively continuous formation of new organic
matter, with little seasonal variation. In a study of coastal lagoon sediments in Fiji there
were no statistically significant seasonal variations in either benthic oxygen uptake rates
or rates of pelagic/benthic nutrient exchange (Sobey, 2004), reflecting a lack of seasonal
change in inputs of organic matter. However, our knowledge of coastal zone biogeo-
chemical processes in the tropics is based on a relatively sparse dataset compared
with other regions, and requires much further study. Much of the tropical ocean has
low nutrient concentrations because of the presence of a permanent thermocline
which prevents vertical transport of new nutrients from deeper water. This conspires
to ensure that the tropical coastal zone generally operates at lower ambient nutrient
concentrations than at higher latitudes. While tropical coastal zone nutrient concen-
trations may be higher than those offshore in tropical oceans, they are still usually
lower than those in the coastal zone at higher latitudes. Tropical primary production
therefore becomes dependent upon recycling of nutrients to sustain fresh production,
and is likely to be nutrient limited. However, where nutrients become available from
upwelling or from fluvial inputs in a tropical high-light environment, primary pro-
duction may be intense (e.g. Robertson et al., 1998).

What factors regulate primary production in the different coastal environments?
Clearly, in polar regions it is light that is the principal seasonal regulator of primary
production in coastal regions, at constant very low temperature, but this is not to imply
that light is the only regulator. Primary production may be co regulated by a variety of
interacting factors, and nutrient availability is often a significant factor despite
the apparently relatively high nutrient concentrations. In the Southern Ocean the
occurrence of high-nutrient, low-chlorophyll areas, where primary production is low
despite the presence of relatively high concentrations of nitrate, has been attributed to
lack of available iron, thus limiting photosynthesis (de Baar et al., 1995; Martin et al.,
1990). This iron limitation apparently does not happen in the Arctic, where proximity
to land masses results in higher aeolian inputs of iron and nitrate is reduced to very low

SGM symposium 65

Biogeochemical cycling in the coastal zone 181

50 r

Fig. 7. Seasonal variation in pelagic primary production in different ICES regions of the North Sea
(reproduced with permission from Joint & Pomroy, 1 993).

concentrations during summer activity. Moreover, in coastal zones, in close proximity
to land, iron supply should be higher than in the open ocean. However, low tempera-
ture may have other influences on the ability of algae to sequester nutrients at low
concentrations.

My co-workers and I have shown previously (Nedwell & Rutter, 1994; Ne dwell, 1999;
Reay et aL, 1999) that low temperature apparently reduces the affinity of bacteria and
algae for substrates taken up by active transport, including nitrate, phosphate and
silicate. This has been attributed to stiffening of the cell membrane at low tempera-
tures, which reduces the efficiency of transporter proteins embedded in the membrane.
From the data available, there is no such inhibitory effect of low temperature upon
ammonium uptake, which occurs at least partly by passive or facilitated diffusion
mechanisms which do not depend upon active transport. Primary production in the
Southern Ocean seems to be very dependent upon ammonium uptake, with very low
f-ratios (see Reay et al., 2001), and it is possible that for at least part of the summer
growth season passive uptake of ammonium is sufficient to maintain the rates of
primary production measured. The concentrations of dissolved inorganic nitrogen,
predominantly nitrate, in Antarctic waters often remains comparatively high (> 10 [iM)
even during summer, but addition of nitrate still stimulates algal growth despite
significant concentrations of nitrate already being present, indicating continued limita-
tion by nitrogen even at these apparently replete concentrations.

SGM symposium 65

182 D. B. Nedwell

Table 3. Rates of benthic oxygen uptake reported in polar, temperate and tropical
regions

Site

Rate(mmol 2 m 2 day n )

Source

Polar

E. Svalbard, 226-320 m depth

Bering and Chukchi Seas, 1 9-25 m depth

E. Svalbard, 1 70-240 m depth

E. Greenland, 40 m depth

Bering and Chukchi Seas, 30-52 m depth

Off Newfoundland, 270 m depth

Signy Is., Antarctica

Temperate

Flax Pond, Long Is. Sound

Aarhus Bay

North Sea sediments, 25-81 m depth

NW Mediterranean, 60-80 m depth

Skagerrak, 190-695 m depth

Thames estuary sediments, UK

Colne estuary, UK

Tropical

Thai mangrove

Jamaican mangrove

Fiji lagoon sediment

Indus delta, Pakistan

Hinchinbrook Is., Australia

3-2-11-9

Pfannkuche&Thiel(1987)

8-7-19-2

Henriksenefa/. (1993)

3-9-11-2

Hulth etal. (1994)

17-8

Rysgaard etal. (1996)

7-3-25-5

Henriksenefa/. (1993)

8-4

Pomeroy etal. (1991)

15-85

Nedwell etal. (1993)

81-137

Mackin&Swider(1989)

Jorgensen & Revsbech (1 985)

5-3-27-8

Upton et al. (1993)

4-6-9-9

Tahey etal. (1994)

11-8-16-1

Canfieldefa/. (1993)

12-2-241

Trimmer ef al. (2000)

48-216

Dong ef al. (2000)

17-61

Kristensenefa/. (1991)

31-103

Nedwell et al. (1994)

0-24

Sobey (2004)

16-28

Kristensenefa/. (1992)

2-8-61

Alongiefa/. (1999)

If affinity for nitrate (or other nutrients) is decreased at low temperatures, algae may be
nitrogen-limited despite the presence of significant concentrations of nitrate because
the nitrate cannot be sequestered effectively. This suggestion does not necessarily
contradict the hypothesis that iron limits Southern Ocean but not Arctic Ocean
primary production, as iron uptake is itself probably an active transport process and its
uptake will therefore tend to be retarded at low temperature.

ORGANIC MATTER DEGRADATION AND BIOGEOCHEMICAL
RECYCLING

The strong pelagic-benthic linkage in the coastal zone ensures that the bottom
sediments are major sites of organic matter breakdown and biogeochemical recycling.
In comparing biogeochemical cycling in the different regions we may ask a number of
questions about the benthic communities which bring about this recycling and the

SGM symposium 65

Biogeochemical cycling in the coastal zone 183

environmental factors which control them. It is clear that primary production rates do
not show any consistent latitudinal variation with temperature (see above), being as
great in some polar regions as in the tropics. Similarly, is there any consistent relation-
ship between water temperature and the biogeochemical recycling of organic matter?
Inspection of reported rates of benthic oxygen uptake (Table 3), as a surrogate for
benthic organic matter breakdown, again suggests that there is no consistent difference
between high and low latitudes: the rates of benthic oxygen uptake in Factory Cove,
Signy Is., at water temperatures near °C all year were greater than those reported for
many tropical sediments (Nedwell et al., 1993; Glud et al., 1998). However, the overlap
of gross rates of 2 uptake in the different regions does not imply that the microbial
communities at high or low latitudes are operating in physiologically similar ways.

Pomeroy and co-workers (Pomeroy & Deibel, 1986; Pomeroy et al., 1991) proposed the
'cold water paradigm' that 'bacterial metabolism and growth are depressed to a much
greater degree than those of phytoplankton at low (<4 °C) sea water temperatures',
which would tend to lead to imbalance between the production and mineralization of
organic matter. They argued that less consumption of primary production by bacteria
at low temperature reduced the microbial loop activity and left more organic matter for
metazoan grazers. In an investigation of bacterial growth and primary production
along a north-south transect in the Atlantic Hoppe et al. (2002) demonstrated clear
gradients of primary production and bacterial growth correlated with latitudinal
temperature change. The ratio of bacterial production to primary production both
above and below the equator was most significantly correlated with water temperature.

In a study of growth rates of bacteria in cold oceans measured by thymidine incorp-
oration, Rivkin et al. (1996) concluded that there was a weak relationship between
specific growth rate and temperature, and the mean and median specific growth rates
for bacteria from cold (<4°C) and warm (>4°C) regions were not significantly
different, i.e. the growth rates of bacteria from cold and temperate oceans are similar at
their respective environmental temperatures. However, they pointed out that the
production rate is a function of both growth rate and biomass of cells in the standing
crop, and the bacterial abundance in cold ocean regions can be 10-fold lower than in
temperate or tropical seas.

EFFECT OF TEMPERATURE ON ABILITY TO SEQUESTER
SUBSTRATES BY ACTIVE TRANSPORT

Interaction between organic substrate requirement and temperature has been reported
for marine bacteria (Pomeroy et al., 1991; Pomeroy & Deibel, 1986; Wiebe et al., 1992,
1993); at low temperature a higher concentration of substrate is required to support
a given rate of growth or respiration than at a higher environmental temperature.

SGM symposium 65

184 D. B. Nedwell

This temperature-substrate concentration interaction was attributed (Nedwell 6c
Rutter, 1994; Nedwell, 1999) to decreased affinity for substrates at temperatures below
the optimum temperature for growth (T t ). At temperatures below T micro-
organisms become increasingly less able to sequester low concentrations of substrates
from their environment effectively by active uptake, apparently because as the mem-
brane stiffens as the temperature drops the transport proteins embedded in the
membrane become less effective, and affinity for substrates declines (Nedwell, 1999).
Because of the decreased affinity for substrates at low temperature, higher concen-
trations of substrates are necessary in the environment to counter loss of affinity and to
maintain a given substrate flux into the cell. This loss of affinity below T has been
demonstrated for active uptake of both inorganic solutes such as nitrate by both algae
and bacteria (Ogilvie et al., 1997; Reay et al., 1999) and organic substrates by chemo-
organotrophic bacteria (Nedwell & Rutter, 1994). The paradigm also holds true
for psychrophilic, mesophilic and thermophilic bacteria over their respective ranges of
temperature (Nedwell, 1999).

For a species to be selected by their environmental temperature they must first be
physiologically capable of growing at that temperature, and the membrane of each
species adapts to its broad temperature range by features such as the fatty acid com-
position of the membrane lipids, which affects membrane fluidity (e.g. Russell, 1990,
1992, 1998). The homeoviscous model of Sinensky (1974) proposed that these mem-
brane compositional differences adapt a species to maintain membrane fluidity, and
hence biological function, over a particular range of temperature. A certain degree of
flexibility of this temperature range appears to be possible by variations in the
membrane and key enzymes but in essence the biokinetic range for a particular species
is fixed. Within the broad temperature range for a species, however, their ability
to sequester substrates from the environment appears to decline below T . Note,
however, that we would only expect an adverse effect by a low environmental tempera-
ture on affinity for substrates if the environmental temperature was below the T for
the species, so the extent to which a species population is optimally adapted to
its environmental temperature would appear important. This raises the question of
the extent to which species populations in different temperature environments have T t
at or below their environmental temperature.

In high-temperature environments such as thermal springs the optimum temperature of
the species present conforms closely to the environmental temperature, but in low-
temperature, polar environments it appears that this is not the case. Morita (1975)
defined psychrophiles, mesophiles and thermophiles in relation to their cardinal
temperatures, with obligate psychrophiles able to grow at 0°C with T <15°C. A
category of 'psychrotolerant' types, able to grow at °C but with an optimum > 15 °C,

SGM symposium 65

Biogeochemical cycling in the coastal zone 185

was subsequently added. In reality there is probably a continuum of physiological
adaptation across the temperature range. Feller & Gerday (2003) have argued that
the terms stenothermal and eurythermal psychrophiles are better used, respectively, for
'obligate' psychrophiles able to grow over only a narrow temperature range or 'facul-
tative' psychrophiles able to grow over a wide temperature range. While stenothermal
psychrophiles are undoubtedly present in polar environments, the majority of organ-
isms present seem to be eurythermal psychrotolerants with T for growth much higher
than the environmental temperature (e.g. Morita, 1975; Franzmann, 1996; Upton
& Nedwell, 1989). Additionally, many reports of biogeochemical processes such as
sulfate reduction (Arnosti & Jorgensen, 2003 ; Nedwell, 1989) also indicate optima well
above environmental temperature in polar regions. Li and co-workers (Li, 1980, 1985;
Li et al., 1984) suggested that algal species at the equator have temperature optima for
photosynthesis near to their environmental temperature, but that at higher latitudes
and lower temperatures the environmental temperature is lower than the optimum for
photosynthesis. If this latitudinal change is the general case then it appears that in low-
temperature populations both algal and bacterial species may indeed be operating
under conditions where they are less well adapted physiologically to their environ-
mental temperature than populations at lower, warmer latitudes.

It might be predicted that it would be adaptive for an organism to operate most
efficiently at its environmental temperature, so it is somewhat surprising that, while
this seems to occur in the tropics, it does not seem to be true in high-latitude popula-
tions. This may mean that T is irrelevant to selection. A new model of the effect of
temperature on enzyme activity, the equilibrium model (Daniel et al., 2001; Peterson
et al., 2004), proposes that enzyme activity at any temperature is a function of an
equilibrium between an active and an inactive form of the enzyme; it is this equilibrium
which leads to the lowering of enzyme activity above the T t for the enzyme. It is the
inactive form of the enzyme that is irreversibly denatured at higher temperature, and
the reduction in activity may be due to a reversible change of conformation at the active
site. Similarly, inactivation of enzymes can occur at low temperature. The equilibrium
between active and inactive forms is defined by an equilibrium constant, K , and T
is the temperature when K is 1 (i.e. when active and inactive forms are equal). Peterson
et al. (2004) argue that adaptive evolution is more likely to operate through selection
for T than through T for the enzyme or thermal stability. In polar, low-temperature
populations an increase in temperature may increase the catalytic rate by the Arrhenius
activation energy, but furthermore may change the equilibrium between active and
inactive forms to increase the amount of active enzyme, so that the overall rate of
reaction is enhanced in a compound manner as temperature rises. In that case the
apparently high T for growth of polar micro-organisms may be an artefact which
does not actually reflect their degree of adaptation to their environmental temperature.

SGM symposium 65

186 D. B. Nedwell

In a review of psychrophilic enzymes Feller & Gerday (2003) argue that effective
enzyme activity at low temperature requires highly reactive reaction centres which
are less stable than those of mesophilic enzymes, which will therefore be heat labile.
Psychrophilic enzymes seem to be inactivated at temperatures much lower than those
that cause unfolding of their protein structure, unlike mesophilic or thermophilic
enzymes. This emphasizes the severe heat-lability of the psychrophilic reaction centre
compared with those of mesophiles or thermophiles. Feller & Gerday (2003) point out
that 'a mobile and flexible active site binds its substrate weakly and, indeed, most
psychrophilic enzymes have higher K m values than their mesophilic counterparts'.
Furthermore, cold-active enzymes maintain reaction rates by increasing k c . n (the
maximum enzyme reaction rate at a given temperature) at the expense of K m , i.e.
affinity as defined by K m is likely to decline (K m increase) at low temperatures. Feller &
Gerday (2003) proposed that 'psychrophilic enzymes have reached a state that is close
to the lowest possible stability... and they cannot be less stable without losing the
native and active conformation'. Does this provide a rationale for why physiological
adaptation by micro-organisms to low temperature is not as complete as that attained
at higher environmental temperatures? The assumption is extending the model for
single enzyme reactions to cell growth, but is it too fanciful to imagine that the effect of
temperature upon growth may mimic its effect upon the enzyme reactions that support
growth? If true, it may also explain why, at low environmental temperatures, micro-
organisms are more regulated by low affinity for substrates than in warmer environ-
ments, because they are operating suboptimally in terms of temperature and/or the
reaction centres of key psychrophilic enzymes (which may include transport proteins)
confer a low affinity for substrates.

The implications of this interaction between affinity for substrates and environmental
temperature are profound. Mincks et al. (2005) argue that microbial respiration and
mineralization of organic matter is a product of both substrate concentration (S) and a
first-order reaction rate constant (k) and that benthic community respiration (R) is
therefore described by

where S. and k x are the respective rate constants and concentrations for metabolizable
components (i) of the organic matter (Fig. 8a). As a given k { decreases due to low
temperature, the rate can be maintained only by an increase in S t .

Alternatively, the rate can be described (Button, 1993) using the specific affinity, a A , and
the rate of substrate utilization is

Roc2tf A(i) S f X

SGM symposium 65

Biogeochemical cycling in the coastal zone 187

c
o

■M

(Li
u

c
o

on
-Q

8
6
4 H
2 ■

(a)

1-6
14 J

as 1-2 ■

a) +j 0*6
I § °' 4

S 8 o-2

2 4 6 8

/C (respiration rate constant)

(b)

r 0035

■ 0030

■ 0025

• 0020

-C

■ 0015

■ 0-010

■ 0-005

• 0000

10 15 20

Temperature (°C)

Fig. 8. (a) Conceptual comparison between effects of substrate concentration and respiration rate
constant for benthic organic matter in polar and temperate sediments (after S. L. Mincks and others,
personal communication), (b) Effect of temperature on specific affinity (a A ) for glucose for an Antarctic
eurythermal coryneform (■) and the substrate concentration required at each temperature to
maintain a constant rate of substrate utilization of 0-001 5 umol h~ 1 (•).

where a A ^ and S 1 are the specific affinity and concentration of substrate i, and X is the
microbial biomass. As a A decreases because of lower temperature, the rate can only be
maintained by either higher substrate concentrations or a higher biomass. Fig. 8(b)
shows how a A varies with temperature for a eurythermal coryneform with a T of
22 °C, isolated from an Antarctic lake. In order for a constant rate of substrate
utilization of 0-0015 fimol h _1 to be maintained as temperature decreases below T
(assuming unit biomass, X), the decreasing a A must be countered by an increased
substrate concentration, S. This leads to the conclusion that microbial communities
might be regarded as operating either at high affinity (a A ) but low 5, as in tropical
environments, or at low affinity but high S, as in polar environments (Fig. 8a). This is
analogous to fast turnover of small substrate pools in tropical systems compared with
slow turnover of relatively large substrate pools in polar benthic environments (the
'food reserve' concept). Holmboe et al. (2001) compared anoxic decomposition in Thai
mangrove sediment and temperate Wadden Sea salt-marsh sediments. They reported
high C : N and C : P ratios in both sediments, but low nutrient release and faster rates

SGM symposium 65

188 D. B. Nedwell

of turnover of N and P by nutrient-deficient bacteria in the tropical sediment, which is
consistent with the proposed model.

Presumably, temperate environments exhibit changes between the two extremes in
response to the seasonal temperature cycle and seasonal selection of the populations
in the microbial community Previous studies (King & Nedwell, 1984; Sieburth, 1967)
have demonstrated the selection of temperate bacterial communities by seasonal
temperature change, the physiological change induced in the community lagging 2
months behind the environmental temperature change. Selection by competition is
not immediate in response to seasonal temperature change, and the 2-month lag is the
period during which the competition occurs, resulting in seasonal selection of better-
adapted populations. This presumably results in cycling between selection of low-
affinity/high-substrate-concentration populations in the spring towards higher-affinity/
low-substrate-concentration populations during summer. This continual seasonal
change in itself implies that adaptation to environmental temperature by microbial
populations in eurythermal temperate environments with seasonally varying tempera-
ture must be less complete than in stenothermal environments, and it is inevitable that
physiological adaptation with respect to environmental temperature will be less
optimized than in stenothermal environments. We have shown previously (Rutter &
Nedwell, 1994) that, in non-steady-state temperature environments, the speed of
response by a species to change may be more adaptive than their competitive ability
under constant conditions. In stenothermal, constant low or high temperature environ-
ments the speed of response to temperature change will not be an issue.

It would seem, therefore, that physiological adaptation in polar populations might be
suboptimal relative to environmental temperature because of the inherent limitation of
adaptability of enzymes to very low temperature; in temperate populations it must be
suboptimal because of the time lag between selection of populations in response to
seasonal environmental temperature change; but in stenothermal tropical systems
populations will be physiologically better adapted, with T near to environmental
temperature.

FUNCTIONAL GROUPS

While gross rates of biogeochemical processes apparently do not differ between
latitudinal regions, the microbial communities which catalyse them may differ. The
process of benthic organic matter breakdown is achieved by the integrated activities of
a variety of different functional groups of bacteria and archaea, together with meio-
and macrofauna, including aerobes, and anaerobes such as nitrate respirers, Fe and
Mn 4+ respirers, sulfate reducers and methanogenic archaea. Firstly, are there any

SGM symposium 65

Biogeochemical cycling in the coastal zone 189

differences in the relative importance of the different functional groups of micro-
organisms present in the different latitudinal regions? Secondly, are there any phylo-
genetic differences in these functional groups in the different regions: i.e. are there any
specifically low-temperature or high-temperature types?

The evidence available suggests that there are no obvious major differences in func-
tional groups in polar or temperate sediments. Benthic oxygen uptake rates in Arctic
and Antarctic coastal sediments are similar (e.g. Nedwell et al., 1993; Glud et al., 1998)
to those of temperate sediments at temperatures 20-25 °C higher. In Young Sound,
Greenland, about half of the organic input was preserved by burial in bottom sedi-
ments (Rysgaard et al., 1998) while, of the organic matter mineralized to C0 2 , aerobic
respiration accounted for 38 %, sulfate reduction for 33 %, iron reduction for 25 % and
denitrification for only 4 % (manganese reduction was unimportant in these sedi-
ments). Reduced iron (and manganese if present) will be reoxidized by 2 in the
surface, oxic layer of sediment, facilitated by bioturbation of the surface layers of
sediment, but failure to measure iron and manganese reduction directly will result in
the significance of anaerobic mineralization by iron and manganese being under-
estimated and appearing as aerobic mineralization. Similar proportions of aerobic
respiration and sulfate reduction have been reported for Antarctic sediments, although
iron and nitrate respiration were not measured directly (Nedwell et al., 1993). Rates of
sulfate reduction were similar to those reported for temperate sediments, although the
optimum temperature for sulfate reduction (15 °C) was very much higher than ambient
temperature.

Methanogenesis in high-sulfate marine sediments is usually low because of competitive
inhibition by sulfate reducers (Nedwell, 1984). While there is some evidence that low
temperature may influence carbon flow through benthic communities, inhibiting
hydrogenotrophic methanogenesis and shifting towards acetoclastic methanogenesis
and acetogenesis (Nozhevnikova et al., 1997; Schulz & Conrad, 1996; Fey & Conrad,
2000), there is no current evidence for this occurring in high-latitude coastal marine
environments.

As in most coastal zone sediments, nitrate respiration in polar sediments accounts for
only a very small proportion of organic matter degradation, largely due to the relatively
low concentrations of nitrate in bottom water. Of the denitrification in Young
Sound, Greenland, the vast majority (93 %) was derived from coupled nitrification-
denitrification (Rysgaard et al., 1998) within the sediment, and not from nitrate from
the water column, where nitrate concentrations were low. Rates of denitrification
reported from Arctic sediments (Devol et al., 1997; Glud et al., 1998; Henriksen et al.,
1993) are similar to those reported from temperate sediments (e.g. Dong et al., 2000;

SGM symposium 65

190 D. B. Nedwell

Jenkins & Kemp, 1984; Lohse et al., 1996; Nielsen, 1992; Rysgaard et ai, 1995) and give
no indication that low temperature influences nitrate respiration.

While functional group activity in coastal zone biogeochemical processes between
regions seems similar, is there any evidence of different phylogenetic composition of
polar and temperate/tropical functional group communities bringing about these
processes? Molecular studies to date have largely focused on the phylogenetically
cohesive groups (sulfate reducers and methanogenic archaea) to which the 16S rRNA
approach is applicable, rather than to the phylogenetically diverse aerobic and nitrate-
reducing communities. Molecular studies of polar marine sediments (Purdy et al., 2003;
Ravenschlag et al., 1999, 2001; Sahm et al., 1999) have shown them to have diverse
communities of sulfate reducers and methanogenic archaea, although the diversity of
both the archaeal and sulfate-reducer communities in Antarctic coastal sediment was
lower than in communities in temperate coastal sediments (Purdy et al. 2003). Only five
distinct groups of archaea were present in Signy Is. coastal sediment compared with 12
from temperate estuarine sediments. It has been suggested previously (e.g. Colinvaux,
1993) that community diversity is likely to be lower in physically or chemically stressed
environments than in communities which are not physically or chemically stressed but
where biological competition controls diversity, and this may be the case in these
Antarctic sediments.

Methanogenesis in a Signy Is. coastal sediment represented only 2 % of carbon flow,
and there was an equivalently small proportion (0-1 %) of methanogenic archaeal
rRNA in the total sedimentary rRNA. (In contrast, in Signy Is. freshwater lake sediment
archaeal rRNA represented 34% of the prokaryotic community.) The coastal
methanogens included members of both the Methanosarcinales, particularly Methano-
coccoides, which metabolize C 1 compounds, and Metbanomicrobiales, which meta-
bolize hydrogen. Inhibition of sulfate reducers immediately stimulated methanogenesis
in the coastal sediment, which suggests that, as in temperate sediments, the methano-
gens are outcompeted for substrates by sulfate reducers, which dominated the terminal
steps of organic matter breakdown. Of the total prokaryotic rRNA in the Signy Is.
coastal sediment, 15 % represented sulfate reducers, dominated by the Desulfotaleal
Desulforhopalus group. Desulfotalea are psychrophiles first isolated from Arctic
sediments (Knoblauch et al., 1999; Sahm et ai, 1999), and seem to be quantitatively
important in polar sediments. They are all incomplete oxidizers of their organic
substrates, as are Desulfovibrio, which have not been detected in either Arctic or
Antarctic sediments (Sahm et ai, 1999); Knoblauch et al. (1999) suggested that the
Desulfotaleal Desulforhopalus group is the ecological equivalent in polar sediments of
Desulfovibrio. As these groups are incomplete oxidizers, what completes the minerali-
zation of organic matter in low-temperature coastal sediments by oxidizing acetate to

SGM symposium 65

Biogeochemical cycling in the coastal zone 191

C0 2 , a step which in temperate coastal sediment is brought about by sulfate reduction
(Nedwell, 1984) and which in temperate estuarine sediments was attributed to
Desulfobacter (Purdy et ai, 1997, 2001) ?

The acetoclastic groups Desulfobacter and Desulfotomaculum reported in temperate
sediments were not detected in the Antarctic sediment. Ravenschlag et al. (1999)
reported a large number of clones in a general bacterial clone library from Svalbard
sediment that were related to S°- or Fe° + - or Mn 4+ -reducing Desulfuromonas, which
utilize acetate, and closely related clones were present in Signy Is. coastal sediment
(Purdy et al., 2003). The implication is that acetate oxidation in these low-temperature
sediments may be achieved by the acetate-oxidizing, S°-reducing Desulfuromonas (i.e.
that S° reduction rather than sulfate reduction may be important in the terminal step
of organic carbon mineralization at low temperature) or by acetate-utilizing Pelobacter
or Geobacter.

ROLE OF COASTAL WETLANDS

Biogeochemical cycling of elements in the coastal zone may be significantly influenced
by the presence of coastal wetlands, predominantly salt marshes in the temperate
region and mangroves in the tropics. The two ecosystems may be regarded as ecological
equivalents: which system dominates seems to be determined by temperature in the
coldest month > 16 °C or in the warmest month >24 °C. Salt marshes and mangroves
are highly productive ecosystems and one of the earliest hypotheses was that of
'outwelling' (Teal, 1962) ; that salt marshes export energy in the form of organic matter
to the coastal zone. The same concept was proposed for mangroves (Odum & Heald,
1972; Robertson & Duke, 1990), and also extended to the idea of net export of
nutrients such as nitrogen to the tropical coastal zone, in which primary production
tends to be nutrient limited. Subsequent research indicated that net export from both
salt marshes (Nixon, 1980) and mangroves varied enormously depending upon local
physical conditions and the productivity of the biological community. Twilley et al.
(1992) suggested that export from mangroves declined with increasing latitude as their
productivity and litter production decreased; export from both mangroves and salt
marshes has been related (e.g. Wolanski et al., 1992) to the strength of the local tidal
influence, with high-amplitude tides and strong directional water flow in estuaries
resulting in increased tidal export from estuarine salt marshes or mangroves. Export
of nitrogen to the coastal zone can undoubtedly occur from some mangroves,
predominantly as particulate or dissolved organic nitrogen (e.g. Rivera-Monroy et al.,
1995), but they are usually sinks for inorganic nitrogen. However, the zone of influence
of the mangrove or salt marsh in the adjacent coastal zone seems to be limited. Rodelli
et al. (1984) measured d 13 C values in mangrove tissue, algae and consumer organisms
and could trace the signal from mangrove detritus in the adjacent coastal zone in

SGM symposium 65

192 D. B. Nedwell

Malaysia only within about 2 km of the mangrove; similar restricted zones of influence
have been reported for Australian (Alongi, 1990), African (Hemminga et al., 1994;
Marguillier et al., 1997), Indian (Bouillon et al., 2002) and South American (Jennerjahn
& Ittekkot, 1999, 2002) mangroves. The direct impact of mangrove-derived organic
matter therefore seems to be of only limited extent in the adjacent coastal zone.

While the net export of either energy or nutrients from salt marshes or mangroves to
the coastal zone may not be significant on an annual basis, in the temperate region there
may be significant seasonal differences in their exchange and impact on the coastal
zone. Thus, a UK salt marsh was balanced overall with respect to its annual nitrogen
exchange with tidal water but there was net export of dissolved and particulate organic
nitrogen and a smaller amount of ammonium during summer, when coastal sea-
water nitrogen concentrations are low and the coastal zone phyto plankton are nutrient-
depleted, but consistent removal of nitrate from tidal water by the salt-marsh sediments
(Azni & Nedwell, 1986). This ability means that salt marshes, mangroves and estuarine
sediments are highly efficient processors of nutrients in coastal sea water and buffer the
loads of nutrients in the coastal zone, particularly when ambient inorganic nitrogen,
usually nitrate, concentrations are increased by anthropogenic inputs (Corredor 6c
Morell, 1994; Nedwell, 1975; Nedwell et al., 1999). It has been estimated that estuarine
and coastal sediments remove up to 50 % of the nitrate load through temperate
estuaries (Nedwell, 1975; Seitzinger,1988; Nixon et al., 1996). Jickells et al. (2000)
argued that the pristine Humber estuary, with its extensive wetland and salt-marsh
area, was likely to have been an effective sink for both N and P, but removal of >90 % of
these wetlands has reduced this capacity. Nonetheless, the removal of anthropogeni-
cally increased N and P loads through estuaries by estuarine sediments, including salt
marshes and mangroves, remains an important 'buffer' to ameliorate the impact of
these loads on the coastal zone.

The coastal ocean, estuaries and coastal wetlands may also be important in the global
carbon budget. In an extensive review of ocean-atmosphere C0 9 fluxes in the coastal
ocean Borges (2005) has argued that water-air fluxes of C0 2 and fluxes of dissolved
inorganic carbon (DIC) from mangroves and salt marshes are important terms in their
carbon budgets which have so far been ignored. Although the DIC fluxes do not seem to
have significant effect on the DIC and pC0 2 values in adjacent coastal waters, their
air-water C0 2 exchange, when scaled up, seems to have a significant influence upon
global C0 2 budgets. If estuaries and coastal wetlands are not taken into consideration,
the coastal ocean appears to be a sink for atmospheric C0 9 (-1-17 mol C m~ 2 year _1 ),
and the estimated uptake of C0 2 by the global ocean is -1-93 Pg C year -1 . If estuaries
and coastal wetlands, with their high heterotrophic activity, are included in the scale-up,
the coastal ocean behaves as a net source of C0 2 (0-38 mol C year ) and the C0 2

SGM symposium 65

Biogeochemical cycling in the coastal zone 193

uptake by the total global ocean decreases to -144 Pg year -1 . There are also interesting
latitudinal differences. At high latitudes and in the tropics and subtropics the model
predicts that, including estuaries and coastal wetlands, the coastal ocean is a net source
of CO? to the atmosphere, but at temperate latitudes it remains a moderate sink for
atmospheric C0 7 . Much work remains to be done to constrain these global estimates
adequately, but the significance of the coastal zone processes seems clear.

ACKNOWLEDGEMENTS

/ would particularly like to thank my colleagues Kevin Purdy and Martin Embley for their
productive, enjoyable, extensive and continued collaboration on the application of molecular
techniques to coastal zone microbial ecology over an extended period and to the Natural
Environment Research Council, UK, for their support in a number of research grants for this
work. Thanks also to the British Antarctic Survey, particularly Cynan Ellis-Evans, for their
collaboration and logistic support for the Antarctic research at Signy Island. I also acknowledge
the helpful comments of Professor Michael Danson on adaptation of extremophile enzymes to
high or low temperature.

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SGM symposium 65

Fungal roles and function in rock f
mineral and soil transformations

Geoffrey M. Gadd, Marina Fomina and Euan P. Burford

Division of Environmental and Applied Biology, Biological Sciences Institute,
School of Life Sciences, University of Dundee, Dundee DD1 4HN, Scotland, UK

INTRODUCTION

The most important perceived environmental roles of fungi are as decomposer organ-
isms, plant pathogens and symbionts (mycorrhizas, lichens), and in the maintenance of
soil structure through their filamentous growth habit and production of exopolymers.
However, a broader appreciation of fungi as agents of biogeochemical change is lacking
and, apart from obvious connections with the carbon cycle, they are frequently
neglected within broader microbiological and geochemical research contexts. While the
profound geochemical activities of bacteria and archaea receive considerable attention,
especially in relation to carbon-limited and/or anaerobic environments (see elsewhere
in this volume), in aerobic environments fungi are of great importance, especially when
considering rock surfaces, soil and the plant root-soil interface (Gadd, 2005a). For
example, mycorrhizal fungi are associated with ~ 80 % of plant species and are involved
in major mineral transformations and redistributions of inorganic nutrients, e.g.
essential metals and phosphate, as well as carbon flow Free-living fungi have major
roles in the decomposition of plant and other organic materials, including xenobiotics,
as well as mineral solubilization (Gadd, 2004). Lichens (a symbiosis between an alga or
cyanobacterium and a fungus) are one of the commonest members of the microbial
consortia inhabiting exposed subaerial rock substrates, and play fundamental roles in
early stages of rock colonization and mineral soil formation. Fungi are also major
biodeterioration agents of stone, wood, plaster, cement and other building materials,
and it is now realized that they are important components of rock-inhabiting microbial
communities, with significant roles in mineral dissolution and secondary mineral
formation (Burford et al., 2003a, b). The objective of this chapter is to outline

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

202 G. M. Gadd ; M. Fomina and E. P. Burford

important fungal roles and function in rock, mineral and soil transformations and to
emphasize the importance of fungi as agents of geological change.

WEATHERING PROCESSES AND THE INFLUENCE OF MICROBES

The composition of the Earth's lithosphere, biosphere, hydrosphere and atmosphere
is influenced by weathering processes (Ferris et al., 1994; Banfield et al, 1999; Vaughn
et al., 2002). Rock substrates and their mineral constituents are weathered through
physical (mechanical), chemical and biological mechanisms; the relative significance
of each process varying widely depending on environmental and other conditions.
Near-surface weathering of rocks and minerals, which occurs in subaerial (i.e. situated,
formed or occurring on or immediately adjacent to the surface of the Earth) and sub-
soil (i.e. not exposed to the open air) environments, will often involve an interaction
between all three mechanisms (White et al., 1992), and it is the biological component of
the overall process that provides much microbiological interest. At or near the Earth's
surface, interaction between minerals, metals and non-metallic species in an aqueous
fluid nearly always involves the presence of microbes or their metabolites (Banfield &
Nealson, 1998). Mineral replacement reactions in rocks mainly occur by dissolution-
re-precipitation processes (e.g. cation exchange, chemical weathering, leaching and
diagenesis), where one mineral or mineral assemblage is replaced by a more stable
assemblage (Putnis, 2002). Micro-organisms can influence this by mineral dissolution,
biomineralization and alteration of mineral surface chemistry and reactivity (Hochella,
2002). Mineral dissolution can also be inhibited by extracellular microbial poly-
saccharides that passivate reactive centres on minerals (Welch & Vandevivere, 1994;
Welch et al., 1999). In contrast, mineral dissolution may be accelerated by micro bially
mediated pH changes and many other changes in solution chemistry. Excretion
of organic ligands and siderophores not only enhances nutrient acquisition by micro-
organisms but can markedly affect mineral composition and dissolution reactions
(Grote & Krumbein, 1992; Maurice et al., 1995; Hersman et al., 1995; Stone, 1997;
Gadd, 1999; Kraemer et al, 1999; Liermann et al., 2000; Sayer & Gadd, 2001). In
addition, the formation of secondary minerals (biogenic crystalline precipitates) can
occur through metabolism-independent and metabolism-dependent processes, often
being influenced by abiotic and biotic factors such as environmental pH, microbial cell
density and the composition of cell walls (Ferris et al., 1987; Gadd, 1990, 1993; Arnott,
1995; Thompson & Ferris, 1990; Douglas & Beveridge, 1998; Fortin et al, 1998;
Sterflinger, 2000; Verrecchia, 2000). The term 'biomineralization' refers to biologically
induced mineralization where an organism modifies the local microenvironment,
creating conditions that promote chemical precipitation of extracellular mineral phases
(Hamilton, 2003). Secondary minerals can form by direct nucleation on cellular
macromolecules, e.g. melanin and chitin in fungal cell walls (Gadd, 1990, 1993; Fortin

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Mycotransformation of minerals 203

& Beveridge, 1997; Beveridge et al., 1997). However, indirect precipitation of secondary
minerals also frequently occurs as a result of microbially mediated changes in solution
conditions (Fortin &£ Beveridge, 1997; Banfield et al., 2000). Regardless of the mech-
anism, lithospheric weathering of rocks and minerals can result in the mobilization and
redistribution of essential nutrients (e.g. P, S) and metals (e.g. Na, K, Mg, Ca, Mn, Fe,
Cu, Zn, Co and Ni) required for plant and microbial growth. In addition, non-essential
metals (e.g. Cs, Al, Cd, Hg and Pb) may also be mobilized from mineral and soil pools
(Gadd, 1993, 2001a, b; Morley et al, 1996).

Bioweathering can be defined as the erosion, decay and decomposition of rocks
and minerals mediated by living organisms. Micro-organisms, as well as animals and
plants, can weather rock aggregates through biomechanical and biochemical attack on
component minerals (Goudie, 1996; Adeyemi 6c Gadd, 2005). Filamentous micro-
organisms, plant roots and burrowing animals can physically affect rocks and enhance
splitting and fractionation: the disruptive (hydraulic) pressure of growing roots and
hyphae is important here (Sterflinger, 2000; Money, 2001). However, biochemical
actions of organisms are believed to be more significant processes than mechanical
degradation (Sterflinger, 2000; Etienne, 2002). Microbes, e.g. bacteria, algae and fungi,
and plants can mediate chemical weathering of rocks and minerals through excretion of
for example H + , organic acids and other metabolites, while respiratory CO-, can lead to
carbonic acid attack on mineral surfaces (Johnstone & Vestal, 1993; Ehrlich, 1998;
Sterflinger, 2000; Gadd & Sayer, 2000). Biochemical weathering of rocks can result in
changes in the microtopography of minerals through pitting and etching, mineral
displacement reactions and even complete dissolution of mineral grains (Ehrlich, 1998;
Kumar & Kumar, 1999; Adeyemi St Gadd, 2005).

MICROBES IN ROCK AND MINERAL HABITATS

Micro-organisms occur in rock and building stone in a variety of microhabitats, being
classed as epilithic, hypolithic, endolithic, chasmolithic, cryptoendolithic and euendo-
lithic organisms (Fig. 1) (Gerrath et al., 1995, 2000; May, 2003). Epiliths are common
under humid conditions and occur on the surface of rocks and building stone. Hypo-
lithic micro-organisms are often found under and attached to pebbles, particularly in
hot and cold deserts. Endoliths inhabit the rock subsurface and may form distinct
masses or brightly coloured layers. Endolithic micro-organisms can occur as chasmo-
liths that grow in pre-existing cracks and fissures within rock, often being visible from
the rock surface. Conversely, cryptoendoliths grow inside cavities and among crystal
grains and cannot be observed from the rock surface. Euendolithic microbes are a
specialized group of cryptoendoliths that are capable of actively penetrating (boring)
into submerged rock (Ehrlich, 1998; Gerrath et al., 1995).

SGM symposium 65

204 G. M. Gadd, M. Fomina and E. P. Burford

Fig. 1. Terminology for rock-dwelling micro-organisms. 1, Epilithic micro-organisms, occurring on
the surface of rocks and building stone; 2, hypolithic micro-organisms, found under and attached to
pebbles; 3-5, endolithic micro-organisms, inhabiting the rock subsurface, which include chasmoliths
(3), which grow in pre-existing cracks and fissures within the rock, cryptoendoliths (4), which grow
inside cavities and among crystal grains, and euendoliths (5), which are a specialized group of
cryptoendoliths that are capable of actively penetrating (boring) into rock. Organisms shown can
represent various microbial groups including bacteria, cyanobacteria, microalgae, fungi and lichens.

Micro-organisms play a fundamental role in mineral transformations in the natural
environment, most notably in the formation of mineral soils from rock and the cycling
of elements (May, 2003; Gadd, 2005a). It is not surprising therefore that a wide variety
of micro-organisms, including bacteria, algae and fungi, inhabit rocks and stone-
work of buildings and historic monuments (Ehrlich, 2002; Burford et al., 2003a;
Gleeson et al., 2005). Exposed surfaces are not necessarily conducive to microbial
growth, as a result of moisture deficit, exposure to UV solar radiation and limited
availability of nutrients. However, complex interactions between microbes and the
mineral substrate are frequently observed, often to some distance into the mineral
(May, 2003). The rock micro-environment is subject to diurnal and seasonal changes in
for example temperature and moisture as well as in available nutrients (Gorbushina &
Krumbein, 2000; Roldan et al., 2002). Nutrients may accumulate as a result of water
interactions, wind-blown dust particles, animal faeces and death and degradation of
living organisms and be utilized by microbes. Mineral grains within the host-rock may
also serve as a source of metals essential for microbial growth. The transfer of
biological material (e.g. fungal spores and other reproductive structures) from external

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Mycotransformation of minerals 205

sources may also play a role in the colonization of subaerial environments by microbes.
Physical properties (e.g. porosity) and elemental composition of the host rock (e.g. C, P,
K, S and metal content) may govern initial establishment, growth and survival of
microbial communities (Gleeson et al., 2005). Thus, colonization of rock substrates by
microbes and the development of a microbial consortium is likely to be influenced
by physical and chemical properties and interactions based on environmental (e.g.
macro-/micro-climate) and biological factors resulting in and influencing ecological
succession at the micro-scale.

MICROBIAL PROCESSES INFLUENCED BY MINERALS

Many important microbial processes can be influenced by minerals, including energy
generation, nutrient acquisition, cell adhesion and biofilm formation (Hochella, 2002;
see elsewhere in this volume). Micro-organisms can also acquire essential nutrients for
microbial growth from mineral surfaces, which effectively concentrate these vital
nutrients far above surrounding environmental levels, e.g. C, N, P, Fe and various
organic compounds (Vaughn et al., 2002). Some environmental contaminants, which
may be concentrated on mineral surfaces by various sorption reactions, can be dis-
placed by similar microbial processes (Kraemer et al., 1999). In addition, it is likely that
potentially toxic metals, released from minerals as a result of physico-chemical and
biological processes, will have an effect on microbial communities (Gadd, 2005b).
Mineral surface properties (e.g. microtopography, surface composition, surface charge
and hydrophobicity) also play an integral role in microbial attachment and detachment.
They are therefore critical in biofilm formation and the ecology of microbial popu-
lations in and on mineral substrates (Wolfaardt et al., 1994; Fredrickson et al., 1995;
Bennett et al., 1996; Rogers et al., 1998) .

FUNGI IN THE TERRESTRIAL ENVIRONMENT

Fungi are ubiquitous components of terrestrial microbial communities, with soil being
regarded as their most characteristic habitat. Subaerial rock surfaces can be considered
an inhospitable habitat for fungal (and other microbial) growth due to their high degree
of insolation, desiccation and limited availability of nutrients (Gorbushina & Krum-
bein, 2000). Micro-organisms that thrive under these extreme conditions have been
termed 'poikilotrophic', i.e. able to deal with varying micro-climatic conditions such as
light, salinity, pH and moisture. Microbial biofilms on and in rocks are believed to be
major factors in rock decay and also in the formation of various patinas, films,
varnishes, crusts and stromatolites in rock substrates (Gorbushina & Krumbein, 2000).

Fungi have been reported from a wide range of rock types including limestone,
soapstone, marble, granite, sandstone, andesite, basalt, gneiss, dolerite, amphibolite
and quartz, from a variety of environments (Staley et al., 1982; Gorbushina et al., 1993;

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206 G. M. Gadd ; M. Fomina and E. P. Burford

Sterflinger, 2000; Verrecchia, 2000; Burford et al., 2003a, b). It is likely that they are
ubiquitous components of the microflora of all rocks and building stone, throughout
a wide range of geographical and climatic zones. Despite the apparent inhospitality
of the rock environment, the presence of organic and inorganic residues on mineral
surfaces or within cracks is thought to encourage proliferation of fungi and other
microbes. Waste products of algae and bacteria, dead cells, decaying plant material,
dust particles, aerosols and animal faeces can all act as nutrient sources for fungi
(Sterflinger, 2000). Some extremophilic fungi are especially evolved to survive and
exploit microhabitats on and within mineral substrata, occurring within the lichen
symbiosis or as free-living microcolonial fungi (Gorbushina et al., 1993; Bogomolova
et al., 1998; Sterflinger, 2000). Microcolonial fungi include those black fungi that
occur as spherical clusters of tightly packed thick-pigmented-walled cells or hyphae
(Gorbushina et al., 1993; Bogomolova et al., 1998). As well as these, other filamentous
fungi, including zygomycetes, ascomycetes and basidiomycetes, often occur on rock
surfaces (epiliths) and in cracks, fissures and pores (endoliths). Certain fungi may also
actively 'burrow' into rock substrates (cryptoendoliths). In addition, many deutero-
mycetes (Fungi Imperfecti), which only exhibit asexual reproduction, are commonly
found in mineral substrates (Kumar & Kumar, 1999; Sterflinger, 2000; Verrecchia, 2000).

In soil, fungi generally comprise the largest pool of biomass (including other microbes
and invertebrates) and this, combined with their filamentous growth habit, ensures that
fungus-mineral interactions are an integral component of biogeochemical processes in
the soil (Gadd, 1993, 1999, 2000a). They occur as free-living filamentous forms, plant
symbionts, unicellular yeasts and animal and plant pathogens, and play an important
role in carbon cycling and other biogeochemical cycles (Gadd & Sayer, 2000; Gadd,
2005a). Their ability to translocate nutrients through the mycelial network also
provides significant environmental advantages (Fomina et al., 2003; Jacobs et al., 2004).
Mycorrhizal fungi in particular are one of the most important ecological groups of soil
fungi in terms of mineral weathering and dissolution of insoluble minerals (Paris et al.,
1995; Jongmans et al., 1997; Lundstrom et al., 2000; Hoffland et al., 2002; Martino
et al., 2003; Fomina et al., 2004, 2005c).

MECHANISMS OF ROCK WEATHERING BY FUNGI
Biomechanical deterioration

Fungi are an important component of lithobiotic communities (an association of
micro-organisms forming a biofilm at the mineral-microbe interface), where they
interact with the lithic substrate, both geophysically and geochemically (de los Rios
et al., 2002; Burford et al., 2003a, b). Biomechanical deterioration of rocks may occur
through hyphal penetration (e.g. into decayed limestone) and by tunnelling into

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Mycotransformation of minerals 207

otherwise intact mineral material (e.g. along crystal planes in calcitic and dolomitic
rocks) (Kumar & Kumar, 1999; Sterflinger, 2000). Fungal hyphae can also exploit grain
boundaries, cleavages and cracks to gain access to mineral surfaces (Adeyemi & Gadd,
2005). In lichens, cleavage-bound mineral fragments as small as 5 \im in diameter can
accumulate within the lower thallus (Banfield et aL, 1999). An important feature of
hyphal growth is spatial exploration of the environment to locate and exploit new
substrates (Boswell et aL, 2002, 2003; Jacobs et aL, 2002a). This is facilitated by a range
of tropic responses that determine the direction of hyphal growth. Among the tropisms
that may occur, thigmotropism or contact guidance is a well-known property of fungi
that grow on and within solid substrates (Watts et aL, 1998). Hyphal growth can often
be influenced by grooves, ridges and pores in solid substrate and is more prevalent in
weakened mineral surfaces. Chemotropism, and other nutritional responses, may also
be important in stress avoidance, such as that raised by toxic metals (Fomina et aL,
2000a; G&ddetaL, 2001).

The process of invasive hyphal growth due to turgor pressure inside hyphae allows
fungi to acquire nutrients from many solid materials (Money, 2001). Highly pressurized
hyphae can penetrate tougher substrates than those with lower pressures, and fungi that
naturally invade hard materials generate extraordinarily high pressures (Money, 1999;
Money & Howard, 1996). Melanin has also been implicated in the penetrative ability of
plant-pathogenic fungi, as this black pigment facilitates the development of infection
structures (Wheeler & Bell, 1988; Money & Howard, 1996). Rock-dwelling fungi are
often melanized (Gorbushina et aL, 1993; Sterflinger, 2000).

Biochemical deterioration

Biochemical actions of fungi on rocks are believed to be more important than
mechanical degradation (Kumar 6c Kumar, 1999). Fungi can solubilize minerals and
metal compounds through several mechanisms, including acidolysis, complexolysis,
redoxolysis and by mycelial metal accumulation (Burgstaller & Schinner, 1993). So-
called 'heterotrophic leaching' by fungi primarily involves the first two mechanisms
and occurs as a result of several processes, including proton efflux via the plasma
membrane H + -ATPase and/or maintenance of charge balance during nutrient uptake,
the production of siderophores [for Fe(III) mobilization] or as a result of respiratory
C0 2 production. In many fungal strains, however, an important leaching mechanism
occurs through the production of organic acids (e.g. oxalic and citric acid) (Adams
et aL, 1992; Gadd, 1999, 2001a; Sayer & Gadd, 2001 ; Jarosz-Wilkolazka & Gadd, 2003 ;
Fomina et aL, 2005a). In addition, fungi excrete many other metabolites with metal-
complexing properties, e.g. amino acids and phenolic compounds (Manley & Evans,
1986; Miiller et aL, 1995). Fungal carboxylic acids can play a significant role in the
chemical attack of mineral surfaces, since the production of organic acids provides a

SGM symposium 65

208 G. M. Gadd ; M. Fomina and E. P. Burford

source of protons for solubilization and metal-chelating anions which complex metal
cations (Miiller et al., 1995; Gadd, 1999, 2001a; Sayer & Gadd, 2001; Fomina et al.,
2004). In one study on the effect of microscopic fungi on the mobility of copper, nickel
and zinc compounds in polluted Al-Fe-humus podzols, it was found that the mobility
of all studied metals increased under the impact of fungi and was predetermined mostly
by the decomposition of soil organic matter (Bespalova et al., 2002).

FUNGAL DETERIORATION OF ROCK AND BUILDING STONE

Microbial attack on minerals may be specific and may depend on the groups of micro-
organisms involved, e.g. some lichen hyphae overgrew augite and mica but avoided
quartz (Aristovskaya, 1980). Substrate acidification by a free-living and different ecto-
mycorrhizal species varied between species as well as in relation to the different
minerals. Mycena galopus and Cortinarius glaucopus produced the highest acidity
per unit biomass density, with higher substrate acidification resulting during growth
on tricalcium phosphate (Rosling et al., 2004).

In podzols, quartz and kaolin are usually overgrown by fungi and algae, with abundant
fungal hyphae also being associated with apatite particles (Aristovskaya, 1980). It
seems that alkaline (basic) rocks are generally more susceptible to fungal attack than
acidic rocks (Eckhardt, 1985; Kumar &£ Kumar, 1999). Along with other organisms,
fungi are believed to contribute to the weathering of silicate-bearing rocks, e.g. mica
and orthoclase, and iron- and manganese-bearing minerals, e.g. biotite, olivine and
pyroxene (Kumar & Kumar, 1999). Callot et al. (1987) showed that siderophore-pro-
ducing fungi were able to pit and etch microfractures in samples of olivine and glasses
under laboratory conditions. A polycarboxylate siderophore, rhizoferrin, showed the
ability to bind Cr(III), Fe(III) and Al(III) (Pillichshammer et al., 1995). Fungi can also
deteriorate natural glass and man-made antique and medieval glass (Krumbein et al.,
1991). Degradation of aluminosilicates and silicates is believed to occur as a result of
the production of organic acids, inorganic acids, alkalis and complexing agents (Rossi
& Ehrlich, 1990). It is also likely that C0 2 released during fungal respiration can
enhance silicate degradation by carbonic acid attack (Sterflinger, 2000). Aspergillus
niger can degrade olivine, dunite, serpentine, muscovite, feldspar, spodumene, kaolin
and nepheline. Penicillium expansion can degrade basalt, while Tenicillium simplicis-
sirnum and Scopulariopsis brevicaulis both release aluminium from aluminosilicates
(Mehta et al., 1979; Rossi, 1979; Sterflinger, 2000). Piloderma was able to extract K and/
or Mg from biotite, microcline and chlorite to satisfy nutritional requirements. When
grown under low-K conditions, the organism showed fibrillar growths, hyphal swellings
and hyphae devoid of ornamentation, possibly indicating nutrient deficiency.
Differences were found in growth rates, morphologies and the Mg content of hyphae
grown with chlorite and biotite, suggesting that Mg was limiting to normal growth.

SGM symposium 65

Mycotransformation of minerals 209

Energy-dispersive X-ray analysis indicated that Pilo derma extracted significantly more
K from biotite than from microcline. The high Ca and O content of hyphal orna-
mentation mainly resulted from calcium oxalate crystals (Glowa et al., 2003).

In podzol E horizons under European coniferous forests, the weathering of horn-
blendes, feldspars and granitic bedrock has been attributed to oxalic, citric, succinic,
formic and malic acid excretion by saprotrophic and mycorrhizal hyphae. Ectomycor-
rhizal fungi could form micropores (3-10 jxm) in weatherable minerals and hyphal tips
could produce micro- to millimolar concentrations of these organic acids (Jongmans et
al., 1997; van Breemen et al., 2000). In order to quantify the contribution of mineral
tunnelling to the weathering of feldspars and ecosystem influx of Ca and K, surface
soils of 11 podzols were studied by Smits et al. (2005). Tunnels were observed only
in soils older than 1650 years, with the contribution of tunnelling to mineral weathering
in the upper mineral soil being less than 1 %. Feldspar tunnelling corresponded to an
average ecosystem influx of 04 g ha -1 year -1 for K and 0-2 g ha -1 year 1 for Ca over 5000
years of soil development. These data indicate that the contribution of tunnelling to
weathering is more important in older soils, but remains low (Smits et al., 2005).

Fungal weathering of limestone, sandstone and marble is also known to occur (Kumar
& Kumar, 1999; Ehrlich, 2002). In hot and cold deserts and semi-arid regions, clump-
like colonies of epi- and endolithic darkly pigmented microcolonial fungi are common
inhabitants of limestone, sandstone, marble and granite, as well as other rock types
(Staley et al., 1982; Sterflinger, 2000; Gorbushina et al., 1993). Analysis of desert rock
samples has shown colonies or single cells in connection with pitting and etching
patterns, suggesting acid attack of the mineral surface, possibly a result of organic acid
or carbonic acid production (Sterflinger, 2000). Microcolonial fungi have also been
shown to be common inhabitants of biogenic oxalate crusts on granitic rocks (Blazquez
etal.,1997).

Acidolysis, complexolysis and metal accumulation were involved in solubilization of
zinc phosphate and pyromorphite by a selection of soil fungi representing ericoid and
ectomycorrhizal plant symbionts and an endophytic/entomopathogenic fungus,
Beauveria caledonica. Acidolysis (protonation) was found to be the major mechanism
of both zinc phosphate and pyromorphite dissolution for most of the fungi examined
and, in general, the more metal-tolerant fungal strains yielded more biomass, acidified
the medium more and dissolved more of the metal mineral than less-tolerant strains.
However, Beauveria caledonica excreted a substantial amount of oxalic acid (up to
0-8 raM) in the presence of pyromorphite that coincided with a dramatic increase in
lead mobilization, providing a clear example of complexolysis (Fig. 2) (Fomina et al.,
2004,2005a).

SGM symposium 65

210 G. M. Gadd, M. Fomina and E. P. Burford

20 -

15 --

10 --

5 --

m
U

E
O

0.8 mM oxalic acid

r"i

in p.

00
■M

Fig. 2. Importance of oxalic acid in fungal solubilization of pyromorphite. The mobilization index
(Ml), which represents the ratio of solubilized lead to the initial amount of pyromorphite, is shown
for several fungal strains and an abiotic control. Bars represent sem. Fungal cultures are identified as:
DGC3, Hymenoscyphus ericae DGC3(UZ); OmCd, Oidiodendron maius Cd; Bc4, Beauveria caledonica
4; LI8, Laccaria laccata 8; Pi23, Paxil lus involutus 23; Pi1 5, P. involutus 1 5; SI21 , Suillus luteusll; SI33,
S. luteus 33; MG1 , Suillus bovinus MG1 ; LSt8, 5. bovinus LSt8; Tt, Telephora terrestris (adapted from
Fomina et al, 2004).

Calcium carbonate (CaC0 3 ) and calcium magnesium carbonate [CaMg(C0 3 ) 2 ] occur
extensively on the Earth's surface as limestone and dolomite (Ehrlich, 2002). Near-
surface calcretes and dolocretes cover as much as 13 % of the total land surface and are
an important reservoir of carbon, accounting for almost 80 % of the total HCO3,
CO? - and C0 2 in the Earth's lithosphere (Ehrlich, 2002; Goudie, 1996). Numerous
micro-organisms, including bacteria and fungi, have been isolated from natural lime-
stone formations. Cryptoendolithic (i.e. actively penetrating the rock matrix to several
millimetres in depth) and chasmolithic or endolithic (i.e. living in hollows, cracks and
fissures) fungi are known to occur in limestone. The production of organic acids is
believed to play a major role in degradation of limestone (Ehrlich, 2002).

The chemical basis for carbonate weathering is the instability of carbonates in acid
solution:

CaC0 3 + H + - Ca 2+ + HCO3
HC03 + H + -H 2 C0 3
H 2 C0 3 - H 2 + C0 2

(equation 1)
(equation 2)
(equation 3)

SGM symposium 65

Mycotransformation of minerals 211

Since Ca(HC0 3 ) 2 is very soluble compared with CaC0 3 , the CaC0 3 dissolves even in
weakly acidic solutions. In strong acid solutions, the CaC0 3 dissolves more rapidly as
carbonate is lost from the solution as C0 2 . Any organism capable of producing acidic
metabolites is capable of dissolving carbonates and even the production of respiratory
C0 2 during respiration can have the same effect:

CO 2 + H 2 ~ H 2 C0 3 (equation 4)

H 2 C0 3 + CaC0 3 «* Ca 2+ + 2HC0 3 (equation 5)

Degradation of sandstone by fungi is also well documented and this has been attributed
to the production of for example acetic, oxalic, citric, formic, fumaric, glyoxylic,
gluconic, succinic and tartaric acids (Gomez-Alarcon et ai, 1994; Hirsch et al., 1995;
Sterflinger, 2000). Fungi can also attack rock surfaces through redox attack of mineral
constituents such as Mn and Fe (Timonin et ai, 1972; Grote & Krumbein, 1992; de la
Torre & Gomez-Alarcon, 1994).

METAL BINDING AND ACCUMULATION BY FUNGI

Fungal biomass provides a sink for metals, by either (i) metal biosorption to biomass
(cell walls, pigments and extracellular polysaccharides), (ii) intracellular accumulation
and sequestration or (iii) precipitation of metal compounds onto hyphae. As well as
immobilizing metals, the reduction in external free metal activity may shift the equili-
brium so that more metal is released into the soil solution (Gadd, 1993, 2000b; Sterf-
linger, 2000). Ectomycorrhizal fungi can release elements from apatite and wood ash and
accumulate them in the mycelia. Suillus granulatus contained 3-15 times more K (3 mg
g _1 ) and showed large calcium-rich crystals on rhizomorphs when grown with apatite,
while Paxillus involutus had the largest amount of Ca (2-7 mg g _1 ). Wood ash addition
increased the amounts of Ti, Mn and Pb in rhizomorphs (Wallander et ai, 2003).

Fungi are effective biosorbents for a variety of metals, including Ni, Zn, Ag, Cu, Cd
and Pb (Gadd, 1990, 1993). Metal binding by fungi can occur through metabolism-
dependent or metabolism-independent binding of ions onto cell walls and other
external surfaces and can be an important passive process in both living and dead
fungal biomass (Gadd, 1990, 1993; Sterflinger, 2000). Metal-binding capacity can be
influenced by external pH, with binding capacity decreasing at low pH for metals such
as Cu, Zn and Cd (de Rome & Gadd, 1987). Cell density also effects binding capacity,
with lower cell densities allowing a higher specific metal uptake per unit of biomass
(Gadd, 1993). The presence of melanin in fungal cell walls also strongly influences
biosorptive capacity (Gadd & Mowll, 1995; Manoli et ai, 1997). Metal localization
was investigated in the lichenized ascomycete Trapelia involuta growing on a range
of uraniferous minerals, including metazeunerite [Cu(U0 2 ) 2 (As0 4 ) 2 . 8H 2 0], meta-

SGM symposium 65

212 G. M. Gadd, M. Fomina and E. P. Burford

Fig. 3. Formation of metal oxalate crystals by Beauveria caledonica. (a) Dry-mode environmental
scanning electron microscope (ESEM) image of calcium oxalate (weddelite and whewellite) crystals
produced on agar containing calcium carbonate (M. Fomina and G. M. Gadd, unpublished); (b) wet-
mode ESEM image of zinc oxalate dihydrate formed by the mycelium grown on zinc phosphate and

torbernite [Cu(U0 2 ) 2 (P0 4 ) 2 . 8H 2 0], autunite [Ca(U0 2 ) 2 (P0 4 ) 2 . 10H 2 O] and uranium-
enriched iron oxide and hydroxide minerals. The highest U, Fe and Cu concentrations
occurred in the outer parts of melanized apothecia, indicating that metal biosorption
by melanin-like pigments was likely to be responsible for the observed metal fixation
(Purvis etal y 2004).

MYCOGENIC MINERAL FORMATION

Fungi have been shown to precipitate a number of inorganic and organic mineral
compounds, e.g. oxalates, carbonates and oxides (Arnott, 1995; Verrecchia, 2000;
Gadd, 2000a; Grote 6c Krumbein, 1992). This process may be important in soil, as
precipitation of carbonates, phosphates and hydroxides increases soil aggregation.

SGM symposium 65

Mycotransformation of minerals 213

covered with a mucilaginous hyphal network; (c) light microscopy images of zinc phosphate particles
adsorbed by the fungal mycelium after 7 days of growth at 25 °C in static liquid medium; (d) crystals
of zinc oxalate dihydrate in the same mycelium after 1 days of growth (adapted from Fomina etal.,
2005a). Bars, 20 urn (a, c, d) and 50 urn (b).

Cations like Si 4+ , Fe 3+ , Al 3+ and Ca 2+ (that may be released through dissolutive mech-
anisms) stimulate precipitation of such compounds, which act as bonding agents for
soil particles. Roots and hyphae can enmesh particles together, alter alignment and
release organic metabolites that assist aggregate stability (Bronick & Lai, 2005).

Oxalate precipitation

Calcium oxalate is the most common form of oxalate associated with soils and leaf
litter, occurring as the dihydrate (weddelite) or the more stable monohydrate (whewel-
lite). Calcium oxalate crystals are commonly associated with free-living, pathogenic
and plant-symbiotic fungi and are formed by the reprecipitation of solubilized calcium
as calcium oxalate (Fig. 3) (Arnott, 1995). Fungal-derived calcium oxalate can exhibit a

SGM symposium 65

214 G. M. Gadd, M. Fomina and E. P. Burford

variety of crystalline forms (tetragonal, bipyramidal, plate-like, rhombohedral or
needles). The formation of calcium oxalate by fungi is important geo chemically, since
it acts as a reservoir for calcium but also influences phosphate availability (Gadd, 1993,
1999; Jacobs e**/., 2002a, b).

Fungi can produce other metal oxalates with a variety of different metals and metal-
bearing minerals, e.g. Cd, Co, Cu, Mn, Pb, Sr and Zn (Fig. 3) (Gadd, 2000a, b; Jarosz-
Wilkolazka & Gadd, 2003; Burford et al., 2003b; Fomina et al., 2005a). Formation of
metal oxalates may provide a mechanism whereby fungi can tolerate environments
containing potentially high concentrations of toxic metals. A similar mechanism occurs
in lichens growing on copper-sulfide-bearing rocks, where precipitation of copper
oxalate occurs within the thallus (Arnott, 1995; Easton, 1997).

Lichens are one of the most common members of the microbial consortia inhabiting
exposed subaerial rock substrates and building stone. Oxalic acid excretion by lichens
can result in the dissolution of insoluble carbonates and silicates and formation
of water-soluble and -insoluble oxalates (Jones & Wilson, 1985; May et al., 1993;
Edwards et al., 1991 ; Verrecchia, 2000). Oxalates can be formed with a variety of metal
ions wherever lichens grow (Purvis, 1996). In particular, there has been concern over the
deteriorative effects of biologically formed oxalic acid on architecturally important
buildings, monuments and frescoes (Nimis et al., 1992). Conversely, however, several
studies have suggested that lichen cover protects certain rock surfaces, acting as an
'umbrella' and reducing the erosion of for example slightly soluble calcium sulfates
(Mottershead & Lucas, 2000).

Carbonate precipitation: calcified fungal filaments in
limestone and calcareous soils

Physico-chemical processes are known to play a significant role in calcrete formation
and development, although it is now recognized that micro-organisms, particularly
bacteria and algae, may also play a crucial and even dominant role in calcrete trans-
formation (Goudie, 1996) . In limestone, fungi and lichens are generally considered to be
important agents of carbonate mineral deterioration. A less studied area of fungal
action, however, is their influence on carbonate precipitation (Goudie, 1996; Sterflinger,
2000). It is already known that many near-surface limestones (calcretes), calcic and
petrocalcic horizons in soils are often secondarily cemented with calcite (CaCO^)
and whewellite (calcium oxalate monohydrate, CaC 2 4 .H 2 0) (Verrecchia, 2000).
Although this phenomenon has partly been attributed to physico-chemical processes,
the presence of calcified fungal filaments in limestone and calcareous soils from a range
of localities confirms that fungi may play a prominent role in secondary calcite
precipitation. Fungal filaments mineralized with calcite, together with whewellite, have

SGM symposium 65

Mycotransformation of minerals 215

been reported in limestone and calcareous soils from a range of localities (Kahle, 1977;
Klappa, 1979; Calvet, 1982; Callot et al, 1985a, b; Verrecchia et al, 1993; Monger St
Adams, 1996; Bruand & Duval, 1999; Verrecchia, 2000). It has also been proposed that
calcium oxalate can be degraded to calcium carbonate, e.g. in semi-arid environments,
where such a process may again act in the cementation of pre-existing limestones
(Verrecchia etal, 1990).

Calcite formation by fungi may occur through indirect processes via the fungal
excretion of oxalic acid and the precipitation of calcium oxalate (Verrecchia et al,
1990; Gadd, 1999; Verrecchia, 2000). For example, oxalic acid excretion and the form-
ation of calcium oxalate results in the dissolution of the internal pore walls of the
limestone matrix, so that the solution becomes enriched in free carbonate. During
passage of the solution through the pore walls, calcium carbonate recrystallizes as
a result of a decrease in C0 2 , and this contributes to hardening of the material.
Biodegradation of oxalate as a result of microbial activity can also lead to transform-
ation into carbonate, resulting in precipitation of calcite in the pore interior, leading to
closure of the pore system and hardening of the chalky parent material. During
decomposition of fungal hyphae, calcite crystals can act as sites of further secondary
calcite precipitation (Verrecchia, 2000). Manoli et al. (1997) demonstrated that chitin,
the major component of fungal cell walls, is a substrate on which calcite will readily
nucleate and subsequently grow.

Reduction or oxidation of metals and metalloids

Many fungi precipitate reduced forms of metals and metalloids in and around fungal
hyphae: for example, Ag(I) reduction to elemental silver [Ag(0)], selenate [Se(VI)] and
selenite [Se(IV)] to elemental selenium [Se(0)] and tellurite [Te(IV)] to elemental
tellurium [Te(0)]. Reduction of Hg(II) to volatile Hg(0) can also be mediated by fungi
(Gadd, 1993, 2000a, b). An Aspergillus sp., P37, was able to grow at arsenate con-
centrations of 0-2 M (more than 20-fold higher than that withstood by Escherichia
coli, Saccharomyces cerevisiae and Aspergillus nidulans), and it was suggested that
increased arsenate reduction contributed to the hypertolerant phenotype of this fungus
(Canovas et al., 2003a, b).

Desert varnish, an oxidized metal layer (patina) a few millimetres thick found on rocks
and in soils of arid and semi-arid regions, is believed to be of fungal and bacterial
origin. Lichenothelia spp. can oxidize manganese and iron in metal-bearing minerals,
such as siderite (FeC0 3 ) and rhodochrosite (MnC0 3 ), and precipitate them as oxides
(Grote & Krumbein, 1992). Similar oxidation of Fe(II) and Mn(II) by fungi leads to the
formation of dark patinas on glass surfaces (Erkhardt, 1985). A Mn-depositing fungus,
identified as an Acremonium-like hyphomycete, was isolated from a variety of labora-

SGM symposium 65

216 G. M. Gadd ; M. Fomina and E. P. Burford

tory and natural locations including Mn(III,iV)-oxide-coated stream-bed pebbles. A
proposed role for a laccase-like multicopper oxidase was postulated, analogous to the
Mn (II) -oxidizing factors found in certain bacteria (Miyata et al., 2004).

Other minerals associated with fungal communities

A wide range of minerals may be deposited under conditions that deviate strongly from
'normal' pressure/temperature diagrams of precipitation and stability and have been
found in association with poikilophilic communities on rock surfaces (Gorbushina
etal.,2002). Some evaporite minerals (gypsum, CaS0 4 .2H 7 0) and iron oxides (magne-
tite, Fe 3 4 ) also display distinctive morphologies indicative of the presence of microbial
communities (Gorbushina et al., 2002). Another biogenic mineral (tepius) has been
identified in association with a lichen carpet that covers high mountain ranges in
Venezuela (Gorbushina et al., 2002). Forsterite (Mg 2 Si0 4 ), the magnesium member
of the olivine [(Mg,Fe) 2 Si0 4 ] mineral solid solution series, is known to occur only in
volcanic rocks, meteorites and metamorphosized carbonates (e.g. skarn deposits). The
presence of forsterite in surficial deposits on rock surfaces can therefore be considered
to be a possible biosignature for former or extant life (Gorbushina et al., 2002).

FUNGAL-CLAY INTERACTIONS

Clay mineral formation and impact on soil properties

Silicon dioxide, when combined with oxides of Mg, Al, Ca and Fe, forms the silicate
minerals in rocks and soil (Bergna, 1994). These high-temperature minerals are unstable
in the biosphere and break down readily to form clays. Micro-organisms play a funda-
mental role in the dissolution of silicate structure in rock weathering, and therefore in
the genesis of clay minerals, and soil and sediment formation (Banfield et al., 1999). In
fact, the presence of clay minerals can be a typical symptom of biogeochemically
weathered rocks (Barker & Banfield, 1996, 1998; Rodriguez Navarro et al., 1997). For
example, in lichen weathering of silicate minerals, Ca, K, Fe clay minerals and nano-
crystalline aluminous Fe oxyhydroxides were mixed with fungal organic polymers
(Barker & Banfield, 1998), while biotite was interpenetrated by fungal hyphae growing
along cleavages and partially converted to vermiculite (Barker & Banfield, 1996). Other
studies have shown that transformation of mica and chlorite to 2 : 1 expandable clays
was predominant in the ectomycorrhizosphere compared with non-ectomycor-
rhizosphere soils, likely to be a result of the high production of organic acids and direct
extraction of K + and Mg 2+ by fungal hyphae (Arocena et al., 1999).

Soil, which can be considered to be a biologically active loose mass of weathered rock
fragments mixed with organic matter, is the ultimate product of rock weathering,
i.e. the interaction between the biota, climate and rocks. Clay minerals are generally

SGM symposium 65

Mycotransformation of minerals 217

present in soil in larger amounts than organic matter and, because of their ion-
exchange capacity, charge and adsorption powers, they perform a significant buffering
function in mineral soils (Ehrlich, 2002) and are important reservoirs of cations and
organic molecules (Wild, 1993; Li & Li, 2000; Dinelli & Tateo, 2001; Dong et aL, 2001;
Krumhansl et aL, 2001).

Biological effects of clay minerals

Fungi are in close contact with clay minerals in soils and sediments. Numerous studies
have shown that interactions of micro-organisms with solid adsorbents lead to
increases in biomass, growth rate and the production of enzymes and metabolites
(Stotzky, 1966, 2000; Martin et aL, 1976; Fletcher, 1987; Marshall, 1988; Claus & Filip,
1990; Lee & Stotzky, 1999; Lotareva & Prozorov, 2000; Lunsdorf et aL, 2000; Deman-
eche et aL, 2001; Fomina & Gadd, 2003). Some clays may stimulate or inhibit fungal
metabolism (Fomina et aL, 2000b; Fomina & Gadd, 2003). Stimulatory effects may
arise from the abilities of different clays to serve as (i) pH buffers, (ii) a source of metal
cationic nutrients, (iii) specific adsorbents of metabolic inhibitors, other nutrients and
growth stimulators and (iv) modifiers of the microbial microenvironment because of
their physico-chemical properties such as surface area and adsorptive capacity (Stotzky,
1966; Marshall, 1988; Martin et aL, 1976; Fletcher, 1987). It has also been shown that
clay minerals (bentonite, palygorskite and kaolinite) influence the size, shape and
structure of mycelial pellets in liquid medium (Fig. 4) (Fomina &C Gadd, 2002).

Fungal-clay mineral interactions in soil aggregation

Fungal-clay mineral interactions play an important role in soil evolution. An examina-
tion of soil clay aggregation by saprophytic {Rbizoctonia solani and Hyalodendron sp.)
and mycorrhizal {Hymenoscyphus ericae and Hebeloma sp.) fungi supported the
hypothesis that fungal hyphae bring mineral particles and organic materials together to
form stable microaggregates at least <2 ^im in size and enmesh such microaggregates
into stable aggregates of >50 [im in diameter (Tisdall et aL, 1997). Fungi not only
entrap soil particles in their hyphae but take part in polysaccharide aggregation as well
(Dorioz et aL, 1993; Martens 5t Frankenberger, 1992; Schlecht-Pietsch et aL, 1994;
Puget et aL, 1999; Chantigny et aL, 1997). Only a few studies have been carried out on
the sorption properties of mixtures of clay minerals (montmorillonite, kaolinite) and
microbial biomass (algae, fungi): interactions between clay minerals and fungi alter the
adsorptive properties of both clays and hyphae (Garnham et aL, 1991; Kadoshnikov
et aL, 1995; Morley & Gadd, 1995; Fomina & Gadd, 2002, 2003).

Clay and silicate weathering by fungi

Fungi and bacteria play an important role in the mobilization of silica and silicates
(Ehrlich, 2002). Their action is mainly indirect, either through the production of

SGM symposium 65

218 G. M. Gadd, M. Fomina and E. P. Burford

Fig. 4. Scanning electron micrographs of Cladosporium dadosporioides grown in the presence of
different clay minerals. (aHc) Mycelial pellets resulting from growth in medium containing 0-5 % (w/v)
(a) and 5 % (w/v) (b) palygorskite or 0-5 % (w/v) kaolinite (c). (d)-(f) Internal structure of the central
zone of fractured pellets grown in clay-free control medium (d) and in the presence of 0-5 % (w/v) (e)
and 5 % (w/v) (f) bentonite. Bars, 1 00 urn (a, b, c) and 10 urn (d, e, f). Reproduced from Fomina &
Gadd (2002).

chelates or the production of acids (mineral or organic) or, as for certain bacteria, the
production of ammonia or amines. Fungi isolated from weathered rock surfaces
{Botrytis, Mucor, Tenicillium and Tricboderma spp.) were shown to be able to solu-
bilize calcium, magnesium and zinc silicates (Webley et aL, 1963). Mobilization of
silicate from clay minerals by Aspergillus niger was found to be a result of oxalic acid
excretion (Henderson & Duff, 1963). The majority of fungal strains belonging to the
genera Aspergillus, Paecilomyces, Penicillium, Scopulariopsis and Tricboderma could
leach iron in submerged culture from a China clay sample (Mandal et aL, 2002). Large
amounts of oxalic, citric and gluconic acids were produced by Penicillium frequentans
in liquid culture. This caused extensive deterioration of clay silicates, as well as micas
and feldspars, from sandstone and granite as a result of organic salts formation such as
calcium, magnesium and ferric oxalates and calcium citrates (de la Torre et aL, 1993).
The oxalate-excreting fungus Hysterangium crassum also weathered clay minerals in
situ (Cromack et aL, 1979).

SGM symposium 65

Mycotransformation of minerals 219

Fig. 5. Deterioration of a concrete block that was exposed to fungal weathering for 2 years in an
experimental microcosm (M. Fomina and G. M. Gadd, unpublished). Cracking is evident, as well as an
extensive hyphal net over the surface. Bar, 200 urn.

MINERAL MYCOTRANSFORMATIONS AND ENVIRONMENTAL
BIOTECHNOLOGY

Concrete biodeterioration and radioactive waste disposal

Fungi play an important role in the deterioration of concrete (Fig. 5) (Perfettini et al.,
1991; Nica et al., 2000), with complexolysis suggested as the main mechanism of
calcium mobilization (Gu et al., 1998). This ability of fungi (and other microbes) to
degrade concrete and other structural materials has implications for underground
storage of nuclear waste. In high-level nuclear waste disposal, the bentonite buffer
around the copper canisters is considered to be a hostile environment for most microbes
because of the combination of radiation, heat and low water availability, but discrete
microbial species can cope with each of these constraints (Pedersen, 1999). Endolithic,
indigenous micro-organisms are capable of surviving gamma-irradiation doses
simulating the near-field environment surrounding waste canisters (Pitonzo et al.,
1999). In 1997-1998, extensive fungal growth was observed on the walls and other
building constructions in the inner part of the 'Shelter' built on the fourth unit of the
Chernobyl nuclear power plant damaged in 1986 (Zhdanova et al., 2000). It was
discovered that low-level gamma-radiation did not affect spore germination, but led
to directed growth of fungal tips towards the radiation source (so-called positive
radiotropism) (Zhdanova et al., 2001).

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220 G. M. Gadd ; M. Fomina and E. P. Burford

Bioremediation

Some of the processes detailed previously have the potential for treatment of con-
taminated land (Gadd, 2000a, b, 2001a, b, 2002, 2004; Hochella, 2002; Fomina et aL,
2005b). Solubilization processes provide a route for removal of metals from soil mat-
rices, whereas immobilization processes enable metals to be transformed into insoluble,
chemically more inert forms. Living or dead fungal biomass and fungal metabolites
have been used to remove metal or metalloid species, compounds and particulates and
organometal(loid) compounds from solution by biosorption. Fungi with Cr(VI)-
reducing activity may have potential for treatment of Cr-polluted soils (Cervantes et aL,
2001). Fungal-clay complex biomineral sorbents may combine the sorptive advantages
of the individual counterparts, i.e. the high density of metal-binding sites per unit area
and high sorption affinity and capacity of fungal biomass, and the high surface area per
unit weight, mechanical strength and efficient sorption at high metal concentrations of
the clay minerals (Fomina & Gadd, 2003). There has also been the use of extracellular
ligands excreted by fungi, especially from Aspergillus and Penicillium spp., to leach
metals such as Zn, Cu, Ni and Co from a variety of solid materials, including low-grade
mineral ores (Brandl, 2001). Mycorrhizal associations may also be used with plants for
metal clean-up in the general area of phytoremediation. Phy to extraction involves the
use of plants to remove toxic metals from soil by accumulation in above-ground parts.
Mycorrhiza may enhance phytoextraction directly or indirectly by increasing plant
biomass, and some studies have shown increased plant accumulation of metals,
especially when inoculated with mycorrhiza isolated from metalliferous environments.
However, this is a simplistic hypothesis and many complicating factors affect successful
exploitation (Meharg, 2003). Several other studies have shown reduced plant metal
uptake (Tullio et aL, 2003). Arbuscular mycorrhiza can depress translocation of zinc to
shoots of host plants in soils moderately polluted with zinc, with binding of metals in
mycorrhizal structures and immobilization of metals in the mycorrhizosphere contri-
buting to these effects (Christie et aL, 2004). Ectomycorrhizal fungi persistently fixed
Cd(II) and Pb(II), and formed an efficient biological barrier that reduced movement of
these metals in birch tissues (Krupa & Kozdroj, 2004). Such mycorrhizal metal
immobilization around plant roots, including biomineral formation, may also assist
soil remediation and revegetation. The development of stress-tolerant plant-mycor-
rhizal associations may therefore be a promising new strategy for phytoremediation
and soil amelioration (Schutzendubel & Polle, 2002). Because of the symbiosis with
ericoid mycorrhizal fungi, ericaceous plants are able to grow in highly polluted
environments, where metal ions can reach toxic levels in the soil substrate (Perotto et aL,
2002; Martino et aL, 2003). Ericoid mycorrhizal fungal endophytes, and sometimes
their plant hosts, can evolve toxic-metal resistance which enables ericoid mycorrhizal
plants to colonize polluted soil. This seems to be a major factor in the success of ericoid
mycorrhizal taxa in a range of harsh environments (Cairney & Meharg, 2003).

SGM symposium 65

Mycotransformation of minerals 221

CONCLUSIONS

It is clear that fungi have important biogeochemical roles in the biosphere. Symbiotic
mycorrhizal fungi are responsible for major mineral transformations and redistribution
of inorganic nutrients, for example essential metals and phosphate, as well as carbon
flow, while free-living fungi have major roles in the decomposition of plant and other
organic materials, including xenobiotics, as well as for example phosphate solubili-
zation. Fungi are dominant members of the soil microflora, especially in acidic
environments, and may operate over a wider pH range than most heterotrophic
bacteria. Fungi are also major agents of biodeterioration of stone, wood, plaster,
cement and other building materials, and are important components, including lichens,
of rock-inhabiting microbial communities, with significant roles in mineral dissolution
and secondary mineral formation. It is timely to draw attention to 'geomycology'
within the umbrella of geomicrobiology and to the interdisciplinary approach that is
necessary to further understanding of the important roles that all micro-organisms
play in the biogeochemical cycling of elements, the chemical and biological mech-
anisms that are involved and their environmental and biotechnological significance.

ACKNOWLEDGEMENTS

Some of the authors' research described within was funded by the BBSRC/BIRE programme
(94/BRE13640) and CLRC Daresbury SRS (SRS grant 40107), which is gratefully acknowledged.
E.RB. gratefully acknowledges receipt of an NERC post-graduate research studentship.
Thanks are also due to Mr Martin Kierans (Centre for High Resolution Imaging and Processing,
School of Life Sciences, University of Dundee, Scotland) for assistance with scanning electron
microscopy.

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SGM symposium 65

The deep intraterrestrial
biosphere

Karsten Pedersen

Deep Biosphere Laboratory, Department of Cell & Molecular Biology, Goteborg University,
Box 462, SE-405 30 Goteborg, Sweden

INTRODUCTION

Exploration of the microbial world started slowly about 350 years ago, when van
Leeuwenhoek and his contemporaries first focused their microscopes on extremely
small living things. It is only during the last 20 years, however, that exploration of the
world of intraterrestrial microbes has gathered momentum. Previously, it had generally
been assumed that persistent life could not exist deep underground, out of reach of the
sun and a photosynthetic ecosystem base. In the mid-1980s, scientists started to drill
deep holes, from hundreds to a thousand metres deep, in both hard and sedimentary
bedrock, and up came microbes in numbers equivalent to those found in many surface
ecosystems. The world of intraterrestrial microbes had been discovered.

Intraterrestrial ecosystems have been reviewed elsewhere and the content of those
reports need not be repeated here (Ghiorse & Wilson, 1988; Pedersen, 1993a, 2000;
Bachofen, 1997; Bachofen et al., 1998; Fredrickson 6c Fletcher, 2001; Amend & Teske,
2005). Instead, this chapter will focus on characteristics that distinguish the intra-
terrestrial from the terrestrial world.

Most ecosystem environments have specific, distinguishing characteristics. The environ-
ments of intraterrestrial microbial ecosystems occupy a special position, differing
substantially in many respects from those of most surface-based ecosystems. In many
ways, underground ecosystems must be approached quite differently from the way in
which those on the surface would be approached. The intraterrestrial world is huge and
diverse and this chapter can only provide a brief overview of the following matters:

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

234 K. Pedersen

Where do intraterrestrial environments begin?

Variability of intraterrestrial environments.

Strategies for exploring intraterrestrial environments.

Organisms living in intraterrestrial ecosystems.

Range of biomass in various intraterrestrial environments.

Intraterrestrial species diversity

Energy sources for intraterrestrial life.

Activity of intraterrestrial life forms.

Several different terms are used to refer to environments under the ground and seafloor
surface. Though it is difficult for one word to encompass all such environments, this
paper generally uses the term 'intraterrestrial'. Other common terms are 'subseafloor',
'subsurface' and 'underground'.

WHERE DO INTRATERRESTRIAL ENVIRONMENTS BEGIN?

There is no consensus as to the answer to this question, and various scientists would
probably suggest different answers. The view adopted in this chapter is that the intra-
terrestrial world begins where contact with surface ecosystems is lost. This lies beneath
the soil and root zones, beneath the groundwater table, beneath sediment and crust
surfaces. A long time must have passed since the last surface contact, 'a long time'
implying at least several decades and often hundreds of years or more. So, it is not depth
per se that defines an intraterrestrial ecosystem, but rather the duration of isolation
from the surface.

VARIABILITY OF INTRATERRESTRIAL ENVIRONMENTS

Many different minerals and rock types constitute our planet. These rock types lie
along a continuum, ranging from very hard rocks (such as granites and basalts) through
sedimentary rocks (such as sandstones) to fairly soft sediments not yet defined as rock.
Just two types will be defined and discussed here: (i) those that are too tight to allow the
presence of microbes elsewhere than in fractures, i.e. hard rocks, and (ii) those that are
porous enough to allow microbes inside the rock, i.e. sedimentary rocks.

When formed, hard rocks generally experience temperatures that greatly exceed the
limit for life and are therefore sterile after formation. They cool off and fracture with
time and the fractures can become colonized by microbes. Sedimentary rocks form
slowly, generally on the seafloor; over geological time scales, plate tectonics may
eventually move these rocks up from the sea to constitute land. Sedimentary rocks that
do not exceed, when formed, the temperature limit for life (currently defined as 113 °C)
can harbour microbes borne by the sediment particles during the sedimentation process
as well as microbes that arrived later.

SGM symposium 65

The deep intraterrestrial biosphere 235

Intraterrestrial environments can be described as basically solid but containing some
groundwater in fractures and pores. In this environment, huge surface areas are exposed
to groundwater; solid dissolution and precipitation processes are consequently very
important. At the opposite extreme, aquatic environments, such as sea or lake environ-
ments, consist of water with tiny amounts of dispersed solids; these solids are thus of
very limited importance for sea and lake water composition.

Mixing processes also differ between intraterrestrial and surface aquatic environments.
In seas and lakes, the water is comparatively homogeneous and well mixed by wind and
currents (although there certainly are layered environments such as meromictic lakes).
Underground, in particular in fractured rocks, there are many different water types that
mix at fracture crossroads. Each water type has it own flow-path history and has met
with different minerals and organisms on its way to a mixing point. Understanding
mixing is crucial when dealing with groundwater. Any groundwater can be charac-
terized by its age relative to when it left the ground surface and also by its origin, usually
defined as an end member of a mixed groundwater. It may seem a simple matter to
clarify the age and origin of a groundwater sample, but it is not. Different dating
methods, for example, analysing the amounts of tritium originating from nuclear bomb
tests or of 14 C or 37 C1, commonly determine the same groundwater to be of different
ages. This is because a single groundwater sample comprises a mixture of waters of
various ages and origins. For example, rainwater penetrating deep underground in the
Scandinavian region, mixing with water of glacial meltwater origin, would produce a
groundwater containing two age signatures: those of recent and of 10000-year-old
water. It would not be obvious whether any microbes detected in this groundwater were
originally borne by the rainwater or the glacial meltwater. Understanding the origins of
the various component waters of a groundwater is a challenge, but can be approached
by statistical methods. By statistical analysis of extensive datasets of ground-
water chemistry, end members have been identified for Scandinavian groundwater
(Laaksoharju et ai, 1999). Models have been built that guide scientists as to how
different end members contribute to the groundwater being analysed.

The results indicate that, as well as a continuum of types of solids in the intraterrestrial
environment, as explained above, there is also a continuum of groundwater types.
Every sample brought up from underground is more or less different from all other
such samples. This is because groundwater systems are non-homogeneous. Comparing
a surface aquatic environment, such as a sea, with most intraterrestrial environments
results in contrasts. Spatially, the sea varies little within a local area, while a ground-
water environment can vary greatly over very short distances. However, when it
comes to temporal variation, the reverse holds true: the sea has marked diurnal and
seasonal cycles, while the intraterrestrial environment shows no change over days or

SGM symposium 65

236 K. Pedersen

seasons. Observation of intraterrestrial environments may need to continue for many
years before significant changes can be resolved.

STRATEGIES FOR EXPLORING INTRATERRESTRIAL
ENVIRONMENTS

Sampling intraterrestrial micro-organisms requires a borehole or a tunnel. The
preferred method of scientific drilling is the core-drilling technique, which allows
the retrieved cores to be analysed for mineralogy and microbiology The borehole can
later be explored in several different ways.

To sample groundwater from water-conducting fractures or aquifers, part of the
borehole can be packed off and sampled. Fig. 1 depicts several possible sampling
techniques. A fairly simple sampling technique is to pump water though tubing
(Fig. la); however, in the case of deep boreholes, approaching a kilometre or more
in depth, it make take considerable time to pump out the water, and much of the
microbiology and some geochemistry may change in the process. Such changes are
induced by the pressure drop and change in temperature. A second option is to lower
samplers and open them at the sampling depth (Fig. lb, c); in this way, the pressure
of the sample can be kept at the pressure of the sampling depth. This strategy has
been applied in sampling deep Fennoscandian groundwaters (Haveman et al., 1999;
Haveman 6c Pedersen, 2002). If there is access to a tunnel or a mine beneath the surface
of the groundwater table, groundwater can be obtained from boreholes without
pumping (Fig. Id). Deep South African goldmines have allowed scientists to access
the deep intraterrestrial biosphere via boreholes drilled to depths as great as 3-5 km
(Onstott et al., 1998). Sampling of fractures that deliver groundwater to a tunnel does
not require drilling (Fig. le). This approach was used by Ekendahl et al. (2003) when
sampling for fungi in deep groundwater. Finally, it may be possible to work in situ by
installing a laboratory near the microbes in a tunnel and working under in situ
conditions. One example of an underground laboratory can be found in the Swedish
Aspo Hard Rock Laboratory (Fig. 2). There, three water-bearing fractures have been
intersected by drilling and then packed off. Groundwater from those fractures is
circulated through flow cells in the laboratory and then back to the fracture again,
maintaining the pressure, temperature and chemistry exactly at in situ levels. These
laboratory flow cells can be regarded as artificial extensions of the fractures.

The superdeep well SG-3 (12262 m deep) in the Pechenga-Zapolyarny area of the Kola
Peninsula, Russia, is currently the deepest drilled hole in the world (Butler, 1994).
Drilling a superdeep borehole becomes increasingly difficult with increasing depth, and
success depends on the qualities of the geological formation drilled, the quality of
equipment used and the skills of the drilling personnel. Other superdeep boreholes

SGM symposium 65

The deep intraterrestrial biosphere 237

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SGM symposium 65

238 K. Pedersen

Fig. 2. Artist's impression of the MICROBE underground laboratory at the 450 m level in the
Aspo hard rock tunnel. The laboratory (to the right) is situated in a steel container and connected
to three discrete fractures in the rock matrix. Tubing connects the systems in the laboratory with
the groundwater in the aquifers under in situ pressure, chemistry and temperature.

include several drilled as part of the German Continental Deep Drilling Program (KTB)
into the crystalline rock of the Bavarian Black Forest (Schwarzwald) basement in
Central Europe (Butler, 1994); the deepest of the six wells drilled reached a depth of
9100 m, where an in situ temperature of 265 °C was recorded. One of these KTB wells
was searched for hyperthermophiles at a depth of 4000 m; culturable micro-organisms
could not be demonstrated. Another very deep borehole was drilled in Gravenberg,
Sweden, in a search for deep earth gases (Gold, 1992). It reached a depth of 6800 m and
thermophilic bacteria were successfully enriched and isolated from a depth of 5278 m,
where the temperature was 65-75 °C (Szewzyk et al., 1994). The temperature data
obtained from these boreholes indicate that the temperature increase with depth varies
depending on the geographical location of the borehole. In most of Scandinavia, the
increase is very slow, ranging between 1 and 1-5 °C per 100 m; elsewhere the gradient
is much steeper, with the steepest temperature gradients obviously being found in
hot-spring areas.

ORGANISMS THAT LIVE IN INTRATERRESTRIAL
ENVIRONMENTS

Generally, representatives of most major groups of surface-living microbes should
be able to live underground -with some obvious exceptions. The underground and

SGM symposium 65

The deep intraterrestrial biosphere 239

Table 1. Range of total counts in various intraterrestrial environments

Values are means of data from various reviews: Ghiorse & Wilson (1 988), Pedersen (1 993a, 2000),
Bachofen (1997), Bachofen etal. (1998) and Fredrickson & Fletcher (2001). gdw. Gram dry weight.

Environment

Total counts

Unattached (ml -1 )

Attached (cm -2 )

Attached (gdw -1 )

Hard rock fractures

1 3 -1 6

1 5 -1 7

—

Sedimentary rock

1 3 -1 o 6

—

1 4 -1 o 8

Subseafloor sediments

—

1 5 -1 o 8

subseafloor world is dark, water-filled and usually anaerobic with reducing conditions.
The boundary between anaerobic and aerobic environments commonly follows the
groundwater table. Consequently, strictly and facultatively anaerobic archaea, bacteria,
phages, fungi and protozoa should be able to live an active life there, but not strict
aerobes, pathogens, photosynthetic organisms or symbionts of plants, animals or
insects.

RANGE OF BIOMASS IN VARIOUS INTRATERRESTRIAL
ENVIRONMENTS

Many microbes live in a planktonic state, but even more tend to live in an attached
state (Characklis & Marshall, 1990). As the intraterrestrial world is a world of solids
with some interpenetrating water, there is always a nearby surface on which to attach.
Experiments suggest that underground ecosystems in fractured rock are dominated
by attached microbes making up biofilms (Pedersen, 2001), but it remains to be demon-
strated that microbial biofilms exist on fracture surfaces in deep environments. It is
easier to obtain a reasonably undisturbed groundwater sample from a borehole than
it is to obtain a drill core sample of an undisturbed intersected fracture. The data
concerning intraterrestrial organisms, in particular in fractured rock, are thus biased
towards planktonic microbes found in groundwater samples.

Most total counts in hard rock refer to the number of unattached microbes; in
sediments, however, the data generally refer to the weight and include both attached
and unattached cells. Typically, the numbers found in fractures and sediments are
relatively equal to those found in clean surface waters and sediments, although the
range is large (Table 1). A hundred million cells may sound like a lot, but bear in mind
that this refers to cell numbers and not weight. In any case, the total intraterrestrial
biomass is very impressive: the calculations of Whitman et al. (1998) suggest that
the intraterrestrial biomass may nearly equal the surface biomass. Whether these
calculations are off by 50 % or so is immaterial, as there will still be a tremendous
amount of life dwelling deep under our feet.

SGM symposium 65

240 K. Pedersen

INTRATERRESTRIAL SPECIES DIVERSITY

The cloning and sequencing of the 16S/18S rRNA genes of microbes living in their
natural environments has revealed a genetic diversity exceeding any earlier estimates
(Pace, 1997). This methodology has been applied to microbes from various under-
ground sites and has generally revealed great species diversity. The case of the Fenno-
scandian hard-rock aquifers will be used to demonstrate typical results for hard rock. A
total of 385 clones from two sites were sequenced (Ekendahl et al., 1994; Pedersen et al.,
1996a, 1997); 122 unique sequences were found, each representing a possible species
that was not recorded in international databases at the time the analysis was conducted.
Therefore, on average, approximately one-third of the sequenced clones represented
unique species. These studies clearly have not exhausted the sequences, as new
sequences were found in nearly every additional sample repetition. This molecular
work indicates that the deep, hard-rock groundwater environments studied are
inhabited by diverse microbial populations, consistent with the great variability of
hydrogeochemical conditions. Similar results were obtained with DNA sequences
of micro-organisms from the alkaline springs of Maqarin in Jordan (Pedersen et al.,
2004), natural nuclear reactors in Oklo, Gabon (Pedersen et al., 1996b), and in nuclear-
waste buffer material (Stroes-Gascoyne et al., 1997).

The molecular work described above has provided a good insight into the phylogenetic
diversity of hard-rock aquifer micro-organisms but does not reveal species-specific
information unless 100 % identity of the full 16S rRNA gene sequence of a known and
described micro-organism is obtained. The huge diversity of the microbial world makes
the probability of such a hit very small - none of the 122 specific sequences mentioned
above had 100 % identities with described species. Even if a 100 % identity is obtained,
there may yet be strain-specific differences in some characteristics which are not
revealed by the 16S rRNA gene sequence information (Fuhrman & Campbell, 1998). If
species information is required, time-consuming methods in systematic microbiology
must be applied to a pure culture; known genera or species can be identified through
these methods.

Sequencing 16S/18S rRNA genes is a fairly blunt tool for obtaining an understanding
of microbial activity and metabolic diversity. Therefore, our laboratory has been apply-
ing cultivation methods for a long time, concurrent with molecular methods. The most
probable numbers of a range of physiologically different micro-organisms have been
analysed (Haveman et al., 1999; Haveman & Pedersen, 2002); these micro-organisms
are iron-reducing bacteria, manganese-reducing bacteria, sulfate-reducing bacteria,
autotrophic methanogens, heterotrophic methanogens, autotrophic acetogens, hetero-
trophic acetogens and methanotrophic bacteria. All those groups have been found in
hard-rock groundwater. In addition, unicellular eukaryotes, in particular fungi, have

SGM symposium 65

The deep intraterrestrial biosphere 241

Fig. 3. The deep hydrogen-driven biosphere hypothesis, illustrated by its carbon cycle. Under
relevant temperature and water-availability conditions, intraterrestrial micro-organisms are capable
of performing a life cycle that is independent of solar-driven ecosystems. Hydrogen and carbon
dioxide from the deep crust of the Earth are used as energy and carbon sources.

been detected (Ekendahl et al., 2003). Phages are probably present but remain to be
detected.

ENERGY SOURCES FOR INTRATERRESTRIAL LIFE

Where do all these organisms get their energy? Hydrogen may be an important electron
and energy source and carbon dioxide an important carbon source in deep subsurface
ecosystems. Hydrogen, methane and carbon dioxide have been found in micromolar
concentrations at all sites tested for these gases (Sherwood Lollar et al., 1993a, b;
Pedersen, 2001). A model has been proposed (Fig. 3) of a hydrogen-driven biosphere
in deep igneous rock aquifers of the Fennoscandian Shield (Pedersen, 1993b, 2000). The
organisms at the base of this ecosystem are assumed to be autotrophic acetogens
capable of reacting hydrogen with carbon dioxide to produce acetate, autotrophic
methanogens that produce methane from hydrogen and carbon dioxide and aceto-
clastic methanogens that produce methane from the acetate product of the autotrophic
acetogens. This hydrogen-driven, intraterrestrial biosphere should conceptually be
possible anywhere underground or under the seafloor where hydrogen and carbon
dioxide are available. The cycle produces acetate; this is a very versatile source of energy
and carbon for anaerobic microbes, such as iron- and sulfate-reducing bacteria and
acetoclastic methanogens. All components needed for this life cycle have been found

SGM symposium 65

242 K. Pedersen

Hydrolysis

Aerobic bacteria

Denitrifying
bacteria

'Manganese-
reducing'
bacteria -

'Iron-reducing'
bacteria

'Sulfate-

reducing'

bacteria

Sulfur-
reducing bacteria

Methanogens

Acetogenic
bacteria

Q 2 H 2 Q

NO- 3 N 2

Mn** Mn 2 *

CO-

Fe 3+ Fe 3+

CO-

H 2 + C0 2

Acetate

Fig. 4. The degradation of organic carbon can occur via a number of different metabolic pathways,
characterized by the principal electron acceptor in the carbon oxidation reaction. A range of
significant groundwater compounds are formed or consumed during this process. The degradation
follows a typical redox ladder pattern.

in the deep igneous rock aquifers and the expected microbial activities have been
demonstrated (Pedersen, 2001), so the model is supported by the qualitative data
obtained so far. Quantitative data are currently being obtained at the underground
laboratory (Fig. 2).

ACTIVITY OF INTRATERRESTRIAL LIFE FORMS

How much do intraterrestrial organisms depend on their environment and to what
extent can they influence their environment? The organisms may have little to do with
hard rock formation but they certainly take part in the formation of less violently
formed, sedimentary rock types. Once the rock is formed, there are still many processes
that can be influenced by microbes and their activities. Precipitation and dissolution,
weathering of rocks and cycling of nitrogen, carbon, phosphorus and many metals can
be influenced by microbes to various degrees. Such processes generally require active
microbes, and gaining an understanding of activity generally requires in situ and/or
laboratory cultivation of the organisms of interest. Radiotracers can be used (Pedersen
& Ekendahl, 1992a, b; Ekendahl & Pedersen, 1994), and there is also the possibility of
analysing stable isotope fractionations using substrates added by nature (Des Marais,
1999).

SGM symposium 65

The deep intraterrestrial biosphere 243

Another approach to exploring intraterrestrial micro-organisms is to study and
interpret geochemical data. For example, why is there no oxygen in deep ground-
water? Oxygen is the preferred electron acceptor of most microbes and its source is
photosynthesis, a surface-based process. Microbes tend to degrade organic carbon
down the redox ladder (Fig. 4), and oxygen is the first electron acceptor to be utilized.
As rainwater percolates down through the ground to become groundwater, it carries
organic carbon and oxygen. Microbes degrade the organic substances along the ladder
in an ongoing process (Banwart et ai, 1996). Oxygen thus does not penetrate very far
underground (although there will be spots where this happens), so the intraterrestrial
world is thus anaerobic. The intriguing thing about this situation is that the under-
ground environment is obviously kept anaerobic by microbes. Life has changed the
surface of our planet dramatically by rilling it with oxygen via photosynthesis, and it
seems likely that life is also responsible for keeping oxygen out of the intraterrestrial
biosphere. The intraterrestrial biosphere thus reflects times on our planet when
photosynthesis had not yet had an effect.

REFERENCES

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Palaeogeogr Palaeoclimatol Palaeoecol 219,131-155.
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Subsurface Microbiology (ISSM-96). Davos, Switzerland, 15-21 September 1996.

FEM5 Microbiol Rev 20, 179-638.
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Res '\53 l 1-22.
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A.-C, Wallin, B. & Wikberg, P. (1996). Organic carbon oxidation induced by large-
scale shallow water intrusion into a vertical fracture zone at the Aspo Hard Rock

Laboratory (Sweden). J Contam Hydrol2'\ l 115-125.
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by G. R. Bock & J.A. Goode. Chichester: Wiley.
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yeasts isolated from deep igneous rock aquifers of the Fennoscandian Shield. Microb

Ecol 46, 41 6^28.

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Fredrickson, J. K. & Fletcher, M. (editors) (2001). Subsurface Microbiology and

Biogeochemistry. New York: Wiley-Liss.
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analogues for radioactive waste. J Contam Hydro! 55, 161-1 74.
Haveman, S. A., Pedersen, K. & Routsalainen, P. (1999). Distribution and metabolic

diversity of microorganisms in deep igneous rock aquifers of Finland. Geomicrobiol

716,277-294.
Laaksoharju, M., Skarman, C. & Skarman, E. (1999). Multivariate mixing and mass

balance (M3) calculations, a new tool for decoding hydrogeochemical information.

Appl Geochem 14, 861-871 .
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(1998). Observations pertaining to the origin and ecology of microorganisms

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734-740.
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Microbiol Lett 1 85, 9-1 6.
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compounds by bacterial communities in ground water from Southeastern Sweden

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attached and unattached bacteria in boreholes along the access tunnel to the Aspo

hard rock laboratory, Sweden. FEMS Microbiol Ecol 19, 249-262.
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of subterranean bacteria in groundwater at Oklo in Gabon, Africa, as determined by

1 6S RNA gene sequencing. Mol Ecol 5, 427-436.
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Investigation of the potential for microbial contamination of deep granitic aquifers

during drilling using 16S rRNA gene sequencing and culturing methods. J Microbiol

Methods 30, 179-192.
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diversity and activity of microorganisms in the hyper-alkaline spring waters of

Maqarin in Jordan. ExtremophilesS, 1 51-164.
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Lahermo, P. W. (1993a). Evidence for bacterially generated hydrocarbon gas in

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The deep intraterrestrial biosphere 245

Canadian Shield and Fennoscandian Shield rocks. Geochim Cosmochim Acta 57,

5073-5085.
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Proc Natl Acad Sci USA 95, 6578-6583.

SGM symposium 65

Iron, nitrogen, phosphorus and
zinc cycling and consequences for
primary productivity in the oceans

John A. Raven # 1 Karen Brown # 1 Maggie Mackay, 1

John Beardall # 2 Mario Giordano, 3 Espen Granum, 4

Richard C. Leegood, 4 Kieryn Kilminster 5 and Diana I. Walker 5

1 Plant Research Unit, Division of Environmental and Applied Biology, School of Life Sciences,
University of Dundee at SCRI, Scottish Crop Research Institute, Invergowrie, Dundee DD2 5DA,
Scotland, UK

2 School of Biological Sciences, Monash University, Clayton, VIC 3800, Australia

department of Marine Science, Universita Politecnica delle Marche, 60131 Ancona, Italy

department of Animal and Plant Sciences, University of Sheffield, Sheffield 510 2TN, UK

5 School of Plant Biology, University of Western Australia, M090 35 Stirling Highway, Crawley,
WA 6009, Australia

INTRODUCTION

Primary productivity in the ocean amounts to the net assimilation of C0 2 equivalent to
about 50 Pg (petagram, i.e. 10 15 g) C year 1 , while on land this is approximately 60 Pg C
year 1 (Field et al., 1998). Almost all of this primary productivity involves photo-
synthesis, and in the ocean it occurs only in the top few hundred metres, even in waters
with the smallest light attenuation (Falkowski & Raven, 1997). About 1 Pg C of marine
primary productivity involves benthic organisms, i.e. those growing on the substratum
(Field et al., 1998), in the very small fraction of the ocean which is close enough to the
surface to permit adequate photosynthetically active radiation (PAR) to allow photo-
lithotrophic growth. This depth at which photosynthetic growth is just possible varies
in time and space, and defines the bottom of the euphotic zone (Falkowski & Raven,
1997). The remaining ~49 Pg C is assimilated by phytoplankton in the water column
(Field et al., 1998). This chapter will concentrate on the planktonic realm, while
acknowledging the importance of marine benthic primary producers and their inter-
actions with micro-organisms (e.g. Dudley et al., 2001; Raven et al., 2002; Raven 6c
Taylor, 2003; Cooke et al., 2004; Walker et al., 2004).

The global net primary productivity of the oceans is less than that on land, despite
about 70 % of the Earth being covered in ocean and primary productivity over
considerable areas of land being limited by water supply. In considering the role of
nutrient elements such as iron, nitrogen, phosphorus and zinc in contributing to the
restriction of primary productivity in the ocean, we must consider them in the context
of other limitations on primary productivity

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

248 J. A. Raven and others

These constraints are frequently divided by oceanographers into 'bottom-up' (physical
and chemical) factors and 'top-down' (biotic) factors. This division echoes that of
Grime (1974), who separated the constraints on terrestrial primary productivity into
'stress' (factors which constrain the production of primary producer biomass) and
'disturbance' (factors which remove primary producer biomass). The bottom-up
(stress) factors include insufficient or excessive PAR, excess UV-B radiation, too high or
too low a temperature and insufficient supply of any of the chemical elements essential
for growth and completion of the life cycle in a form which is available to the organisms.
The top-down (disturbance) factors are the biotic factors of grazing and parasitism
(including viral attack) and the physical factor of removal by, in the case of phyto-
plankton, sinking out of the euphotic zone.

The bottom-up factors determining the growth rate of marine primary producers are
intimately related to physical oceanography and to global biogeochemical cycles.
Physical oceanography is significant for phytoplankton photosynthesis through the
depth of the upper, mixed layer of the ocean in which the phytoplankton organisms
are entrained, as modified by movement of organisms relative to the surrounding water
as a result of the motility of flagellate organisms and upward and downward movement
of non-flagellate organisms whose density is, respectively, lesser or greater than that of
the surrounding sea water. When the mixing depth is greater than the so-called 'critical
depth', i.e. the depth at which the integrated water-column photosynthesis by phyto-
plankton equals community respiration, net growth of phytoplankton cannot occur
(Falkowski & Raven, 1997). Physical oceanography also impacts on the supply of
nutrients to photosynthetic organisms with respect to cycling of nutrients within
the ocean. Nutrients are carried to depth by the sinking of live and dead biota and
faecal pellets, generically termed 'marine snow'. Below the euphotic zone, most of the
particulate organic material is acted on by decomposer organisms which regenerate
inorganic nutrients that can then be used by primary producers upon return to the
euphotic zone. This return of nutrients to the euphotic zone can occur by seasonal
and permanent upwellings and by eddy diffusion across the thermocline at the base
of the upper mixed layer (Falkowski & Raven, 1997).

With this background to marine primary productivity and its determinants, we will
address the cycles of four nutrient elements, iron, nitrogen, phosphorus and zinc, whose
supply has been shown at times to restrict marine primary production in some parts of
the ocean. This chapter first addresses the biogeochemical cycles of these elements and
the evidence as to their roles in limiting primary productivity, dealing with the elements
in the order in which they are thought to be significant in constraining primary
productivity. Secondly, we will consider the interactions between these elements

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 249

in determining primary productivity. Thirdly, we consider some broader aspects of
elemental supply and marine primary productivity in a changing world.

BIOGEOCHEMICAL CYCLES OF NITROGEN, IRON, PHOSPHORUS
AND ZINC IN RELATION TO MARINE PRIMARY PRODUCTIVITY

Nitrogen

The abundant N 7 in the atmosphere, and dissolved in the ocean, is only directly avail-
able to diazotrophic organisms, i.e. those that use the enzyme nitrogenase to reduce N 7
to NFL, which can then be used to produce the organic N compounds needed by the
organism. N 2 fixation by marine organisms involves mainly cyanobacteria, of which
the filamentous Trichodesmium predominates, although the role of unicellular
cyano bacterial diazotrophs has recently been recognized (Capone, 2001). Other inputs
of combined N, i.e. nitrogen sources other than N 9 , which are available to non-diazo-
trophic primary producers, occur as atmospheric inputs resulting from lightning
converting N 2 to NO v (NO + N0 2 ) and hence HN0 3 , and increasingly as NH
(NHL + NHJ, NO A ., HNO3 and organic N from natural and anthropogenic terrestrial
sources. Combined N supply also occurs as fluvial inputs of NH4, NO3 and organic
N from terrestrial ecosystems, again with increases as a result of man's activities
(Falkowski 6c Raven, 1997; Tyrrell, 1999). The losses of combined N from the ocean
occur as denitrification, whereby NO3 is reduced to N 9 and N 2 in anoxic and hypoxic
regions of the ocean, and by the incorporation of particulate organic N into marine
sediments (Falkowski & Raven, 1997; Tyrrell, 1999). While it has been suggested that
there is a net loss of combined N from the present ocean (Falkowski & Raven, 1997),
consideration of the cumulative errors in the estimation of the components of the
combined N inputs and outputs for the present ocean suggests that the data are
also consistent with a balance of oceanic combined N inputs and outputs (Tyrrell,
1999). Even if there is no net loss of combined N from the ocean at present, there is
strong evidence that primary productivity in a large fraction of the world ocean
is limited by the supply of combined N (Falkowski, 1997; Falkowski & Raven, 1997;
Cullen, 1999; Tyrrell, 1999). This evidence is derived from enrichment experiments in
which samples of surface ocean water are incubated for the determination of primary
productivity, either unaltered or with the addition of particular nutrients, individually
and in combination. The results of such experiments in many parts of the ocean
suggest that the nutrient element limiting growth is N, with diazotrophs unable to
supply combined N at a sufficiently rapid rate. More generally, the atomic ratio of
inorganic combined N to P in the surface ocean is, when averaged over time and space,
rather less than the 16:1 Redfield ratio of the average phytoplankton cell with the
elemental composition of 106C : 16N : IP (Falkowski & Raven, 1997; Falkowski, 2000).

SGM symposium 65

250 J. A. Raven and others

There is a significant difference between the two scenarios for oceanic combined N
status (Falkowski, 1997; Tyrrell, 1999) with respect to the possible constraints on
biological N 2 fixation in the ocean. We deal first with the possibility that there is not a
shortfall in biological fixation, i.e. it is adequate to maintain the oceanic combined N
pool by supplying as much combined N as is needed to replace the excess of combined
N loss by sedimentation and denitrification over aeolian and riverine inputs (Tyrrell,
1999). If this is the case then it may not be necessary to seek exogenous limitations
specific to diazotrophy to account for the magnitude of N 2 fixation other than optimal
allocation arguments that the extra machinery required by diazotrophy restricts the
maximum specific growth rate of N^-fixing organisms relative to those using combined
N. This would permit the diazotrophs to grow whenever there was a shortfall in
combined N supply which constrained the growth of non-diazotrophs, despite the
lower maximum specific growth rates, with the non-diazotrophs ousting the slower-
growing diazotrophs when combined N is more readily available. However, it is known
that there are additional resource costs of diazotrophic growth relative to growth on
combined N, such as the increased energy, Fe and Mo (or V) requirement and a less
well-defined requirement for additional P (Sanudo-Wilhelmy et al, 2001; Vitousek
et al, 2001; Kustka et al., 2003a, b). These constraints on the growth of photosynthetic
diazotrophs relative to that of photosynthetic organisms living on combined N could
account for any inadequacy of N 2 fixation to maintain the oceanic combined N pool
by not allowing the diazotrophs to fix N 9 until there was enough combined N to allow
non-diazotrophs to out-compete diazotrophs (Falkowski, 1997). The observation is
that Fe is commonly a limiting resource for diazotrophy in the surface ocean, although
P and PAR are also significant constraints in some areas (Falkowski, 1997; Sanudo-
Wilhelmy et al, 2001 ; Mills et al, 2004) .

The combined N sources for phytoplankton growth in the ocean include nitrate, nitrite,
ammonium and organic N (Falkowski 6c Raven, 1997). As we have noted, nitrate (and
nitrite) enters the surface ocean by fluvial and aeolian inputs, as well as in upwellings
and eddy diffusion from the deep ocean, where it is produced by nitrification of reduced
N sedimented from the euphotic zone. This 'new' (to the euphotic zone) combined N
supports what is termed 'new' primary production which, since it can be exported
(sedimented) without altering the pre-existing combined N pool in the surface ocean, is
also termed 'export production'. Ammonium/ammonia and organic N supply to
the surface ocean is predominantly from excretion by higher trophic levels as well as
from parasitism of phytoplankton, with only a minor component entering the surface
ocean from aeolian and fluvial inputs. The primary productivity supported by the use
of this reduced N is termed 'recycled' production, since it depends very largely on
combined N recycled from particulate organic matter within the surface ocean. The
reason that this 'recycled' production uses reduced combined N directly, rather than

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 251

after it has been converted to the thermodynamically more stable (in the oxygenated
surface ocean) nitrate, seems to be the kinetic constraint on the exergonic conversion of
ammonium to nitrate by nitrifying bacteria resulting from light inhibition of growth
of the nitrifiers (Falkowski Sc Raven, 1997). These nitrifiers are chemolithotrophs which
can assimilate about 0-19 Pg C from inorganic C per year. The energy used by these
organisms can be regarded as the additional energy, ultimately derived from PAR,
required for phytoplankton growth on nitrate rather than on reduced N in the euphotic
zone. This energy is used to power inorganic C reduction deeper in the ocean (Raven,
1996). The expected lower growth rate of photosynthetic primary producers under
light-limiting conditions resulting from the use of nitrate rather than reduced N as the
N source is, however, not found routinely (Raven, 1996).

Most marine primary producers examined are able to use all of the combined N forms
discussed above, i.e. ammonium/ammonia, one or more species of organic N, nitrite
and nitrate (Falkowski & Raven, 1997). However, the very abundant picoplanktonic
cyanobacterium Procblorococcus is unable to use nitrate, and some strains are also
unable to use nitrite (Bryant, 2003; Dufresne et al., 2003; Palenik et al., 2003; Rocap
et al., 2003). This relatively recent loss of metabolic capability during the evolution of
Procblorococcus from within the paraphyletic genus Synecbococcus is clearly not
a fatal impediment to Procblorococcus, which is probably the most abundant, in terms
of number of individuals, photosynthetic organism on Earth, as well as being amongst
the smallest (Bryant, 2003). Procblorococcus is relatively most successful in the oligo-
trophic ocean, where it depends mainly on recycled combined N. It also typically occurs
deeper in the water column than the metabolically more versatile Synecbococcus, where
the energy-requiring assimilation of oxidized combined N would consume scarce
energy in this light-limited environment (Ting et al., 2002; cf. Raven, 1987; MacFarlane
& Raven, 1990; Falkowski &C Raven, 1997). As we shall see when considering Fe,
the absence of nitrate assimilation capacity reduces the Fe requirement of the cells.

A final aspect of the N nutrition of marine primary producers is the possibility of
economizing on N by replacing a macromolecule with a high N content (usually a
protein) with a macromolecule with a lower N content that performs a similar meta-
bolic function (Raven et al., 2004). Exact functional substitutions are uncommon,
making the effectiveness of this possible means of economizing on N difficult to
evaluate. An example is in photosynthetic light harvesting, where the quantity of
apoprotein per molecule of chromophore is at least twice as high for phycobiliproteins
as for chlorophyll-protein complexes. However, differences in chromophore absorption
spectra and in chromophore-specific absorption coefficients make a like-for-like
comparison of the ecophysiological impact of such a substitution difficult to evaluate.
Perhaps the best-examined case is that of the almost total replacement of phyco-

SGM symposium 65

252 J. A. Raven and others

biliproteins by chlorophyll-protein complexes in Prochlorococcus relative to the
ancestral Synecbococcus (Ting et al., 2002). Such substitutions can reduce the cell
N requirement by over 10 % (Raven et al., 2004). Similar magnitudes of economy in N
costs of growth can be achieved by replacing N-containing with N-free low relative
molecular mass organic compounds involved in screening UV-B, scavenging and
quenching reactive oxygen species, providing osmotic compensation as compatible
solutes and restricting the attentions of grazers and parasites using anti-biophage
compounds (Raven et al., 2004). Again, there are problems in establishing the
effectiveness of the N-free compound relative to the N-containing compound in
the organism under natural conditions.

Iron

Fe input to the surface ocean occurs as dissolved and particulate fluxes down rivers,
much of which is sedimented in the coastal zone, and as aeolian inputs of dust, either
dry or in rain (Berner & Berner, 1996; Chester, 2003; Jickells et al., 2005). There is also
some input of Fe to the deep ocean via hydro thermal vents, while the major output of Fe
from the ocean is in sinking particles incorporated into marine sediments (Chester,
2003). Although there is some input of Fe into the surface ocean from the deep ocean
by upwelling and eddy diffusion, the major inputs of Fe to the ocean surface are
from rivers and, more importantly, from the atmosphere. The atmosphere supplies
essentially all of the Fe inputs to the non-coastal surface ocean (Chester, 2003).

Fe is a very abundant element on Earth, but the global oxygenation which began over
2 billion years ago has restricted the availability of Fe to organisms by converting it into
the insoluble ferric form. While chemical reduction of ferric to ferrous iron can occur
in acidic atmospheric droplets, this soluble ferrous iron is converted to the ferric form
with a half-time of about 2 min when the droplet equilibrates with sea water at a pH of
about 8 (Chester, 2003). The requirement for large quantities of Fe in, for example,
nitrogenase and nitrogenase reductase, photosystem I-like photosynthetic reaction
centres, NADH dehydrogenase complexes and nitrite reductase (Raven, 1988, 1990;
Raven et al., 1999) presumably evolved when Fe was much more readily available, more
than 2-5 billion years ago. Squaring the large Fe requirements in these catalysts with
the current relative unavailability of Fe is now met by a range of Fe-acquisition
mechanisms, as well as a range of means by which the use of Fe is restricted.

Fe acquisition from sea water involves a variety of mechanisms (Volker 6c Wolf-
Gladrow, 1999). Cyanobacteria use siderophores which are secreted by the organism,
and some of the molecules are taken up again after they have acquired Fe(III) from
colloidal or ligated sources (Chester, 2003). Some eukaryotic marine primary producers
can take up Fe(III)-loaded siderophores, although they do not produce the sidero-

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 253

phores (Chester, 2003). The genome of the diatom Thalassiosira pseudonana gives no
indication of the production or uptake of siderophores, but has components of the
mechanism which reduces external Fe(III) to Fe(II) followed by uptake of the Fe(II) by a
process involving, paradoxically, oxidation to Fe(III) during uptake (Armbrust et al.,
2004). This mechanism is widespread in marine primary producers, using Fe(III) from a
range of ligands (Falkowski &C Raven, 1997). A few marine planktonic primary
producers are also phagotrophic, and so are able to take up colloidal or particulate
Fe (Raven, 1997).

The constraints on Fe availability to primary producers have evoked a number of
ecological and evolutionary responses. One response is the restricted occurrence
of diazotrophy and, in some habitats, of the use of nitrate and nitrite as N sources
(Falkowski, 1997; Flynn & Hipkin, 1999; Sanudo-Wilhelmy et ai, 2001; Kustka et ai,
2003a, b; Mills et al., 2004). Another response is the decrease in the Fe requirement for
photosystem I on a cell basis in genotypes with low contents of this kind of reaction
centre. Examples of low photosystem I contents include the cyanobacterium Prochloro-
coccus relative to its ancestor Synecbococcus, and the open-ocean relative to the coastal
species of the diatom Thalassiosira (Strzepek & Harrison, 2004). In at least some cases,
the decreased content of photosystem I constrains acclimation to environments with
varying fluxes of PAR (Strzepek 6c Harrison, 2004). These economies of N use, by
changing the exogenous N source or by decreasing the content of photosystem I, can
account for several-fold reduction in Fe requirement from the most costly to the least
costly resource-acquisition mechanisms (Sunda et ai, 1991; Sanudo-Wilhelmy et ai,
2001; Kustka et ai, 2003a, b; Strzepek & Harrison, 2004; cf. Falkowski et al, 2004).
Much smaller economies can be effected by substituting an Fe-containing component
which accounts for a relatively small fraction of the total cell N by an Fe-free
component which performs essentially the same metabolic function. Examples here are
the replacement of the Fe-containing photosynthetic cytochrome c 6 with the Cu-
containing plastocyanin and of the Fe-containing ferredoxin with the metal-free
flavodoxin (Raven et al., 1999). Here, near-equality of function has been established
(Raven et al., 1999). Such replacements can be facultative (in response to Fe deficiency)
or obligate, e.g. in the 'red' line of plastid evolution where plastocyanin is absent (Raven
et al., 1999; Falkowski et al., 2004). Although the economy in Fe content of the
organisms by these substitutions is small, the expression of flavodoxin in natural
phytoplankton assemblages can be used as an indicator of actual, or incipient, Fe
limitation of growth.

Despite the theoretical likelihood of Fe limitation of primary productivity in at least a
part of the surface ocean, problems with the analysis of Fe in sea water, and in carrying
out incubations under conditions free of Fe contamination, meant that evidence

SGM symposium 65

254 J. A. Raven and others

consistent with Fe limitation in significant areas of the ocean only became available
a little over a quarter of a century ago with the insightful work of the late John Martin
and his colleagues (Martin & Fitzwater, 1988; Martin et ai, 1990, 1991). Martin's
conclusions were based on the chemical analysis of sea water and on the sort of incu-
bation experiments described above for N, involving measurements of primary
productivity of sea-water samples with and without the addition of a range of
nutrients. This work showed that the Fe limitation of primary production occurred
in the so-called 'high-nutrient, low-chlorophyll' (HNLC) areas of the ocean, i.e. those
parts of the ocean that have relatively high concentrations of the major nutrients
nitrate and phosphate in the surface waters, but a low biomass of primary producers
as indicated by the chlorophyll content (Chester, 2003). The HNLC areas of the ocean
include the north-east sub-Arctic Pacific, the eastern tropical Pacific and the Southern
Ocean (Chester, 2003; Jickells et ai, 2005). There are also smaller coastal areas of
Fe limitation in the eastern Pacific (Hutchins et ai, 1998, 2002; Hutchins & Bruland,
1998; Jickells £?*#/., 2005).

The incubation experiments described above did not rule out the possibility of the
in situ occurrence of deep mixing, leading to light limitation of primary production, as
a bottom-up factor, and/or of the impact of grazers or parasites in removing primary
producer biomass as a top-down factor. Clarification of the situation came from meso-
scale experiments in which Fe (as iron sulfate) was added to several square kilometres of
the surface ocean in HNLC areas and the impact of the addition on the chemistry and
biology of the Fe-enriched patch was followed for as long as was possible; observations
were usually terminated by the subduction or horizontal fragmentation of the patch
(Chester, 2003). These experiments have now been carried out more than once in each
of the three main HNLC areas of the ocean and, in each case, a decrease in nitrate and
phosphate and an increase in primary producer biomass (especially large diatoms)
and in primary productivity was found, as well as a delayed increase in the biomass of
herbivorous zooplankton (Watson et al., 2000; Boyd et ai, 2002, 2004; Boyd, 2002a, b),
although light limitation can also contribute to the HNLC state (Boyd, 2002a; van
Oijen et al., 2004). However, it has not yet been established whether the increased
primary productivity is paralleled by an increased export flux of particulate organic
matter, as might be expected from the additional primary productivity resulting from
the exogenous input of the productivity-limiting resource (Boyd et al., 2004). It is of
interest that the shortage of Fe does not alter the ratio of C to N in phytoplankton cells,
so that there seems to be no preferential allocation of what Fe is available to either the
C assimilation or the N assimilation pathways. This absence of preferential allocation
of Fe could well be a function of homoeostasis of the C to N ratio. Readers are directed
to Cooke et al. (2004) and Raven et al. (2004) for further information.

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 255

For diazotrophy in the surface ocean, we have seen in the consideration of N above
that Fe is a common limiting resource for N 2 fixation, although there are parts of the
ocean in which P or light energy can limit diazotrophy (Falkowski, 1997; Sanudo-
Wilhelmy et al., 2001; Mills et al., 2004). As with phytoplankton assimilating combined
N, the diazotroph Trichodesmium does not exhibit variations in the cellular C : N ratio
as a function of Fe availability (see references in Cooke et al., 2004; Raven et al., 2004).

Phosphorus

Almost all of the biology of P occurs at the most oxidized state of phosphate. Phos-
phate is supplied to the ocean mainly via rivers, although there is some atmospheric
input (Berner 6c Berner, 1996; Chester, 2003). However, atmospheric input of Fe can
sequester some of the P in surface water as complexes with ferric iron (Chester, 2003),
recalling the decrease in Fe availability with the onset of atmospheric oxygenation more
than 2 billion years ago (Bjerrum & Canfield, 2002). P loss from the ocean is in sinking
particles which become incorporated into marine sediments (Bjerrum & Canfield,
2002; Chester, 2003). While there has been much emphasis on inorganic phosphate as
the dominant source for marine primary producers, there is an increasing appreciation
of the role of soluble organic phosphates in the P nutrition of marine primary
producers. Generally, organic P sources are used by uptake of inorganic phosphates
after the action of extracellular phosphatases. Those marine planktonic primary
producers which are also phagotrophic can acquire P by ingestion of particles, with
absorption of digestion products across the food vacuole membrane, as well as by
transporters in the plasmalemma (Raven, 1997).

There are cogent geochemical arguments that the supply of P, rather than of N, is
the long-term determinant of the productivity of Earth, since the supply of P (in the
absence of man's activities) is restricted by the rate of weathering on land and is
balanced by sedimentation in the ocean, whereas N supply can be maintained by
diazotrophy in the face of sedimentation and denitrification (Falkowski 6c Raven,
1997). However, as we have seen, there are factors in addition to the availability of P that
restrict the extent of diazotrophy in the ocean (Falkowski, 1997; Tyrrell, 1999; Sanudo-
Wilhelmy et al., 2001; Mills et al., 2004). Nevertheless, there are regions of the ocean
in which there is evidence from measurements of nutrient concentrations in the surface
ocean, and from nutrient enrichment studies on primary productivity of samples
of surface water, consistent with P limitation of primary productivity (Karl et al.,
1993; Karl & Tien, 1997; Cullen, 1999; Wu et al., 2000; Bemtez-Nelson & Karl, 2002;
Ridame & Guieu, 2002; Bjorkman & Karl, 2003; Heldal et al., 2003; Krom et al., 2004).
Particular attention has been paid to P limitation in parts of the Mediterranean Sea
(Krom etal., 2004).

SGM symposium 65

256 J. A. Raven and others

The possibilities of economy in the use of P in marine primary producers by sub-
stitution of another, more readily available, element for P while maintaining equivalent
metabolic function are smaller for P than for N or, especially, Fe. As for reducing
the cell quota of a P-containing component, much of the functional (not-stored) P in
most marine primary producers is contained in rRNA. In many non-photosynthetic
organisms, there is a positive correlation of growth rate and rRNA content even when
growth of a given genotype is constrained by environmental factors other than P supply,
or when growth stages of a metazoan are compared, or closely related taxa are
compared (Sterner & Elser, 2002; Raven et al., 2004, 2005). This correlation is much less
clear for photosynthetic organisms, including the marine primary producers (Sterner

& Elser, 2002; Elser et al., 2003; Agren, 2004; Klausmeier et al., 2004a, b; Leonardos &
Geider, 2004; Raven et al., 2004, 2005). However, the finite rate of protein synthesis
across each rRNA molecule means that there must be a critical rRNA cell quota below
which growth rate is constrained by the rRNA content, although the discussion above
suggests that the rRNA quota in adequately P-nourished marine primary producers
is significantly above this critical level. This apparent excess of rRNA means that a
potential P economy measure in marine primary producers has not been widely
adopted in these organisms.

The genome is a P-containing cell component that is not generally a large fraction of
the non-storage cell P. However, in the very small marine primary producer Prochloro-
coccus growing under moderate P deficiency, the genome contains over half of the cell
P quota (Bertilsson et al., 2003). This is a function of the non-scalable nature of the
components of the genome: the size of a gene encoding a given protein is very similar in
very small cells and in larger cells. Thus, despite having as few as 1500 protein-coding
genes and a high gene density in the genome, the very small size of Prochlorococcus
means that even its miniaturized genome contains a very significant fraction of the total
cell P. The discussion above on N sources for Prochlorococcus shows that there are
constraints on the nutrition of this organism as a function of gene loss, and some
strains have a restricted ability to acclimate to varying light environments.

To a limited extent, the large fraction of cell P in the genome of Prochlorococcus can be
attributed to the small cell quota of P relative to C and N in well-nourished cells
(Bertilsson et al., 2003; Heldal et al., 2003). Although it is tempting to relate these high
C:P and N:P ratios to the oligotrophic and potentially P-limited habitat of
Prochlorococcus, it must be remembered that the low P content in comparison with the
Redfield ratio of 106C : 16N : IP (by atoms) is not exceptional, granted the variability in
the C :P and N : P ratios of marine phytoplankton cultured under nutrient-replete
conditions. The Redfield ratio applies to temporal and spatial averages over the world
ocean (Falkowski & Raven, 1997; Falkowski, 2000; Geider & La Roche, 2002). A

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 257

further complication in assessing the P quota of marine phytoplankton is the
occurrence, in some organisms, of significant (relative to the intracellular content)
surface-bound P (Sanudo-Wilhelmy et al., 2004).

Zinc

Zn enters the ocean by dry and, especially, wet deposition from the atmosphere, as well
as in rivers, and leaves it by sedimentation (Chester, 2003). Evidence consistent with Zn
limitation in parts of the ocean comes from the work of Crawford et al. (2003) and
Franck et al. (2003) using enrichments with Fe, Zn or Fe + Zn in experiments on
primary productivity, or components thereof, using samples of surface ocean water.
However, no drawdown of Zn was found during the stimulation of primary product-
ivity in HNLC regions by the addition of Fe, suggesting that Zn was not close to
growth-limiting in these surface waters (Frew et al., 2001; Gall et al., 2001). There is
evidence consistent with Zn limitation of phytoplankton growth in some freshwater
habitats (Ellwood et al., 2001 ; Sterner et al., 2004) .

The enzymes that use Zn in marine primary producers include carbonic anhydrase,
alkaline phosphatase and, in seagrasses and in peridinin-containing dinoflagellates,
Cu-Zn superoxide dismutase (Raven et al., 1999). Most marine primary producers have
only the Fe/Mn superoxide dismutases (Raven et al., 1999). A possible rationale for the
lack of Cu-Zn superoxide dismutase in most marine primary producers is the very
widespread occurrence of inorganic carbon-concentrating mechanisms (CCMs) in
these organisms which, in many cases, involve an accumulation of bicarbonate ions to
concentrations in excess of those in sea water. Cu-Zn superoxide dismutases catalyse,
in addition to the eponymous dismutation of superoxide to hydrogen peroxide and
oxygen, a bicarbonate-mediated peroxidation activity which can damage the enzyme
itself or, by reacting with exogenous substrates, protect the enzyme from inactivation
(see Elam et al., 2003). Perhaps the absence of Cu-Zn superoxide dismutases in these
organisms with CCMs is related to the increased extent of self-oxidation and damage
of the enzyme as a function of increasing concentrations of bicarbonate. It is of interest
that CCMs in terrestrial vascular plants, i.e. those in plants with C 4 and 'crassulacean
acid metabolism' (CAM), and with Cu-Zn superoxide dismutases in their cytosol and
stroma, accumulate C0 2 rather than bicarbonate (Giordano et al., 2005).

The quantitatively most significant role of Zn in most marine primary producers is in
carbonic anhydrases, although Co or Cd can substitute for Zn in vitro and, in some
cases, in vivo (Giordano et al., 2005). There is also a Cd-specific carbonic anhydrase
in the diatom Thalassiosira weissflogii (Lane & Morel, 2000b; Giordano et al., 2005).
Carbonic anhydrases are involved, inter alia, in CCMs (Giordano et al., 2003, 2005).
These CCMs are frequently expressed more strongly at low inorganic C concentrations

SGM symposium 65

258 J. A. Raven and others

(Giordano et al., 2005), and Morel et al. (1994) and Lane & Morel (2000a) have shown
additional possibilities of Zn limitation at low inorganic C concentrations where
carbonic anhydrase is expressed at higher levels. Subsequently, Reinfelder et al. (2000,
2004) and Morel et al. (2002) have offered evidence consistent with C 4 -like photo-
synthetic metabolism in the planktonic marine diatom T. weissflogii, especially at low
inorganic C levels, which would decrease the requirement for Zn. The argument here is
that, if bicarbonate is the inorganic C form entering the cytosol, no carbonic anhydrase
activity is needed in that compartment because bicarbonate is the inorganic C substrate
for phosphoenolpyruvate carboxylase, the enzyme which generates the C 4 dicarboxylic
acid oxaloacetate (Reinfelder et al., 2000, 2004; Morel et al., 2002). Decarboxylation of
oxaloacetate by phosphoenolpyruvate carboxykinase in the chloroplast stroma
generates C0 2 , the substrate for the core carboxylase of photosynthesis, ribulose
bisphosphate carboxylase-oxygenase (Rubisco), which also occurs in the stroma
(Reinfelder et al., 2000, 2004; Morel et al., 2002). This scheme for inorganic C
assimilation does not require expression of carbonic anhydrase in the cytosol or the
stroma, or on the cell surface, since inorganic C enters as bicarbonate, the predominant
form of inorganic C in sea water. Although the evidence for C 4 metabolism as an
obligatory component of photosynthesis in T. weissflogiiis incomplete (Johnston et al.,
2001; Granum et al., 2005; Giordano et al., 2005), the case is becoming stronger
(Reinfelder et al., 2004). Such a mechanism does not necessitate the accumulation of
bicarbonate in intracellular compartments, so that the argument suggested above for
the absence of Cu-Zn superoxide dismutases from organisms which accumulate
bicarbonate is inapplicable, as it is for higher plant C 4 and CAM photosynthesis.

INTERACTIONS AMONG THE AVAILABILITY OF IRON #
NITROGEN, PHOSPHORUS AND ZINC AND THAT OF OTHER
RESOURCES IN DETERMINING MARINE PRIMARY
PRODUCTIVITY AND ITS FATE

Interactions among Fe, N, P and Zn

The discussion above of the interaction of the cycles of the four elements with marine
primary producers has perforce included discussion of some of the interactions among
these elements, and with some other resources required for growth. An example is the
increased Fe requirement for growth when nitrate rather than ammonium is the
exogenous N source, with an even greater Fe requirement when diazotrophy supplies
ammonia/ammonium from N 9 . We also saw that Fe deficiency does not specifically
decrease the assimilation of nitrate or N 2 relative to inorganic C, despite the involve-
ment of additional Fe-containing enzymes in the assimilation of these N sources which
are not needed for photosynthetic growth with reduced N as the N source. Another
example is the role of Fe in atmospheric dust deposited in the ocean on the availability

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 259

of Pfrom dissolved inorganic phosphate. Finally, Zn is required through its involvement
as a cofactor in most carbonic anhydrases in the photosynthetic assimilation of
inorganic C in most marine primary producers and through its involvement as a
cofactor in extracellular phosphatases in the assimilation of external organic P.

The elementary composition of marine phytoplankton grown under nutrient-sufficient
conditions has been investigated by Quigg et al. (2003), Ho et al. (2003) and Falkowski
et al. (2004). This work shows that the content of trace elements in the eukaryotes
correlates with the evolutionary origin of the plastids, with the 'green' organisms
relatively enriched in Fe, Cu and Zn, while the 'red' organisms are relatively enriched in
Mn, Co and Cd. These trends in trace metal content cannot all be readily explained by
differences in the use of the metals in catalysis, as indicated by the presence or absence
of catalysis involving a given metal, and the variations in the quantity of the different
metal-containing catalysts, among the taxa (Raven et al., 1999; Quigg et al., 2003; Ho
etal.,2003).

Interactions of Fe, N, P and Zn with the assimilation of
inorganic C

We have already discussed some interactions of Fe, N, P and Zn with photosynthesis by
marine primary producers. Here, we consider some other interactions and comment on
the possible impact of increasing surface ocean C0 2 concentrations resulting from
anthropogenic C0 2 emissions to the atmosphere on these interactions. We concentrate
particularly on the impact on the role of CCMs, since these are the means by which
most marine primary producers have their growth rates saturated with inorganic C
under the current environmental conditions (Beardall & Raven, 2004; Giordano et al.,
2005).

Table 1 shows the effect of the availability of Fe, N, P and Zn on the extent of
engagement of CCMs. The evidence as to the extent to which CCMs are involved
in inorganic C assimilation generally comes from studies of the concentration of
inorganic C required to achieve half of the inorganic-C-saturated rate of photo-
synthesis, a lower half-saturation concentration suggesting a greater involvement of
CCMs. More rarely, the evidence comes from the measured ratio of the inorganic C
concentration in the cells to that in the sea-water medium during steady-state
photosynthesis. Some of the data come from work on freshwater rather than marine
organisms.

The effect of limitation of photosynthesis by the supply of PAR, e.g. in benthic primary
producers growing deep under water or in the shade of other organisms, or planktonic
primary producers in a deeply mixed ocean surface layer, is to decrease the engagement

SGM symposium 65

260 J. A. Raven and others

Table 1. Inorganic C affinity and expression of CCMs in cyanobacteria and algae
with CCMs as a function of resource limitation for growth

Growth-limiting
resource

Effect on inorganic

C affinity/CCM expression

Reference(s)

PAR

Decrease

Beardall (1 991 ); Kubler & Raven (1 994, 1 995);
Young & Beardall (2005); Giordano etal. (2005)

Inorganic C

Increase

Giordano et al. (2005)

Inorganic N

Increase (nitrate as N source)

Beardall etal. (1982, 1991); Young & Beardall
(2005)

Decrease (ammonium as
N source)

Giordano etal. (2003)

Inorganic P

Increase

Beardall etal. (2005)

Decrease

Bozena etal. (2000)

Inorganic Fe

Increase

Young & Beardall (2005)

Inorganic Zn

Decrease

Morel etal. (1994); Buitenhuisef al. (2003)

of CCMs in C assimilation. Limitation of photosynthesis by decreased inorganic C
supply can occur in marine habitats, e.g. in high intertidal rock pools with high
densities of primary producers and infrequent flushing with sea water at neap tides, and
in blooms of plankton (Falkowski & Raven, 1997; Raven et al., 2002). Such inorganic
C limitation increases the engagement of CCMs (Table 1). Conversely, the increase in
sea surface water concentration of C0 2 , and to a smaller relative extent in bicarbonate
and total inorganic C as well as decrease in pH, predicted to occur by 2100 leads to a
decreased CCM engagement and a decreased affinity of photosynthesis for inorganic C
(Table 1; Beardall 6c Raven, 2004). A decreased availability of combined N generally
increases the engagement of CCMs, the exception being for an organism grown on
ammonium rather than on nitrate, as were all of the other organisms tested (Table 1).
There are conflicting data as to the effect of a decreased availability of P on CCM
engagement, while the only data available on the effect of limited Fe availability indicate
increased CCM engagement (Table 1). Finally, decreased Zn availability decreases the
engagement of CCMs (Table 1).

These data (Table 1) are in general agreement with predictions from theoretical
considerations on the effect of variations in resource supply. The effects of light
availability can be rationalized in terms of the relatively greater impact of leakage from
the accumulated inorganic C pool as the energy supply for (re-)accumulation is
decreased, and there is an increased 'affinity' for PAR in driving photosynthesis (higher
rate at a given low photon flux density) paralleling the decreased CCM expression
for organisms acclimated to low light, and the reverse for organisms acclimated to
low inorganic C and high light (Kubler 6c Raven, 1994, 1995). The theoretical rationale

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 261

for the commonly observed increase in CCM expression when N availability is restrict-
ed relates to the N costs of the additional protein components related to the CCM
relative to the N costs of the alternative of Rubisco oxygenase and photorespiratory
metabolism (Beardall etal., 1982). Similarly, the increased CCM engagement under low
Fe conditions can be related to Fe requirements for thylakoid components involved in
CCMs and in Rubisco oxygenase activity and photorespiration (Raven 6c Johnston,
1991). The decreased CCM engagement in organisms not thought to use a C 4 -like
metabolism when Zn availability is restricted can be related to the Zn in the increased
levels of carbonic anhydrases attendant on increased CCM activity

While these rationalizations are useful in interpreting the data, further mechanistic
studies are needed, especially related to the interaction of the various resource supply
factors that determine CCM engagement. Also required is work in which the possibility
of long-term genetic adaptation, as well as of short-term acclimation, is addressed
for increased CO ? supply (Collins & Bell, 2004) in relation to other resource supply
factors.

Interaction of Fe, N, P and Zn supply with the fate of marine
primary producers

So far we have concentrated on the impact of the availability of resources on the rate of
primary production, i.e. a consideration of bottom-up or stress factors determining
the accumulation of biomass. We now consider the implications of resource availability
for the fate of the primary producer biomass. This includes the biotic interactions of
grazing and parasitism, i.e. the classic top-down factors, as well as the abiotic factor
of sinking in the case of phytoplankton; all of these would be subsumed under Grime's
concept of disturbance.

Dealing first with the effect of nutrient limitation on grazing, restricted availability of
a nutrient can alter the chemical composition of primary producers and hence their
nutritional value to herbivores (grazers). Most of the work on the impact of the
chemical composition of aquatic primary producers on their nutritional value has
involved fresh waters (Sterner & Elser, 2002). While there are influences on the ingestion
rate of primary producers by grazers as a function of the resource supply conditions
during the growth of the photosynthetic organisms, it is not clear whether these
changes routinely alter the impact of grazers on primary producer populations (Sterner
& Elser, 2002). The situation is made more complex by the general effect of nutrient
deficiency not just on the ratio of major nutritional components such as proteins,
carbohydrates and lipids, and the content of essential trace elements, but by additional
responses of primary producers. The further responses to nutrient deficiency include

SGM symposium 65

262 J. A. Raven and others

the production of additional quantities of organic chemicals which deter or poison
grazers or render the ingested material indigestible. These responses by the primary
producers are a matter of some controversy One debate concerns the extent to which
the production of 'defence compounds' is selected for because of their effect in
restricting removal of biomass when the rate of biomass production is already
constrained by limited nutrient availability (Sterner 6c Elser, 2002). What some workers
perceive as an alternative explanation is that the additional production of at least the
N-free 'defence compounds' is related to the excess of photosynthate over the capacity
of the organism to use the photosynthate in growth when nutrient deficiency constrains
photosynthate use in growth more than it does photosynthate production (Sterner &
Elser, 2002). This contrast may be more apparent than real, with 'excess' organic C in
nutrient-limited primary producers used to make compounds which limit the loss to
grazers of a population whose capacity for replacement is resource-limited. The argu-
ments become especially complex when very small organisms, such as phytoplankton,
are considered (Wolfe et al., 1997; Ianora et al., 2004), as the effectiveness of defence
compounds in restricting grazing in terms of natural selection is a function of the
ability of the grazers to distinguish among prey, and then to restrict the ingestion of
the better-defended cells. Resolution of these problems involves not just consideration
of the sensory and manipulative abilities of the grazers in relation to their learned or
innate behaviour in detecting and avoiding better-defended cells, but also the nature
of the phytoplankton populations (e.g. clonal or outbreeding) and the unit of natural
selection in these organisms (Thornton, 2002; Raven & Waite, 2004). The implications
of these 'defence compounds' in avoiding parasitism, including viral infection, are even
less clear than their effects on grazers.

Additional effects of nutrient deficiency on the palatability and nutritional value
of marine primary producers include the impact of nutrient deficiency on the extent of
mineralization. Tables 2 and 3 show the effects of resource deficiency on the ratio
of mineral material to organic matter in two groups of quantitatively very important
marine primary producers, the silicified diatoms (planktonic and benthic) and the
calcified coccolithophores (planktonic, with benthic stages in the life cycle of a few
coastal species). The data in Table 2 indicate that restricted availability of Fe, N, P and
Zn, as well as of PAR and of inorganic C, increases the content of silica relative to that
of organic matter. Indeed, every factor which increases the length of the vegetative life
cycle of diatoms increases the time of the G2 phase of the life cycle. The G2 phase
is when silicification occurs, so that slowing cell growth increases the relative amount
of silica per cell (Martin-Jezequel et al., 2000). The situation is rather less clear
for calcification in coccolithophores (Table 3). Here, restricted light availability
decreases calcification relative to organic C production, as does restricted bicarbonate

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 263

Table 2. Silicification by diatoms as a function of resource limitation for growth

Growth limitation Effect on particulate Reference(s)
by supply of the silica/particulate

named resource organic C

Martin-Jezequel etal. (2000); Claquin etal. (2002)

Milliganefa/. (2004)

Martin-Jezequel etal. (2000); Claquin etal. (2002)

Hutchins & Bruland (1 998); Takeda (1 998);
De La Rocha etal. (2000); Franckef a/. (2003);
Leynaert etal. (2004)

De La Rocha etal. (2000); Franck etal. (2003)

* Increased cell Si quota at low C0 2 is a result of decreased loss of Si by efflux and dissolution rather
than increased influx of Si(OH) 4 .

PAR

Increase

Inorganic C

Increase*

Inorganic N

Increase

Inorganic P

Increase

Inorganic Fe

Increase

Inorganic Zn

Increase

Table 3. Calcification by coccolithophores as a function of resource limitation for
growth

Growth limitation
by supply of the
named resource

Effect on particulate
inorganic C/particulate
organic C

Reference(s)

PAR

Decrease

Inorganic C

Increase with decreased CO

Decrease with decreased

bicarbonate

Inorganic N

Increase

Inorganic P

Increase

Inorganic Fe

No effect

Inorganic Zn

Increase

Paasche (1964, 1999,2001)

Riebesellefa/. (2000)

Sekino & Shiraiwa (1 994); Shiraiwa (2003)

Paasche (1998, 2001)

Paasche & Bruback (1 994); Paasche (1 998, 2001 )
Schulzef a/. (2004); cf. Crawford etal. (2003)
Schulzef a/. (2004); cf. Crawford etal. (2003)

supply, while restricted C0 2 supply increases calcification. Restricted supply of N,
P and Zn increase calcification, while there is no significant effect of restricted supply
of Fe.

Does the increased mineralization as a result of nutrient deficiency restrict mortality
due to grazing and parasitism via mechanical effects or, in the case of calcification, via
challenges to the maintenance of low pH in any acid parts of the grazer's digestive
system? There is little direct evidence even on how effective the 'nutrient-sufficient'
quantities of silica and of calcite in diatoms and coccolithophores are in decreasing
losses due to grazing and parasitism relative to notionally comparable organisms
lacking silicification and calcification (Raven & Waite, 2004).

SGM symposium 65

264 J. A. Raven and others

Increased mineralization must increase the density of the organisms, since silica and
calcite are each two to three times as dense as the cell protoplast, thus increasing
the sinking rate of planktonic organisms. Raven &C Waite (2004) consider the role of the
increased rate of sinking resulting from the additional 'nutrient-deficient' quantities of
silica and calcite in diatoms and coccolithophores in the context of movement down the
water column to where there might be higher concentrations of nutrients. However,
the vertical water movements in the upper mixed layer typical of habitats favoured by
planktonic diatoms mean that such effects are essentially limited to an increased chance
of falling out of the upper, mixed layer into the more nutrient-rich thermocline, with
the possibility of the less-dense cells resulting from (light-limited but nutrient-suffi-
cient) growth in the thermocline becoming reincorporated into the upper mixed layer
(see Rodriguez et al., 2001; Ptacnik et al., 2003). The variations in density which permit
very large-celled planktonic diatoms such as Ethmo discus and some species of
Rhizosolenia to make vertical migrations taking days within one cell division cycle in
oligotrophic waters, gaining photosynthate near the surface and nutrients at depth,
must depend on changes in protoplast density, since silicification-related changes in
density are not manifest within a cell cycle covering a range of light and nutrient-supply
conditions (Raven & Waite, 2004).

Aside from the effects of mineralization and its variation with resource supply in the
life of diatoms and coccolithophores, the increased ballast per unit organic matter
in nutrient-limited cells increases the potential for sinking of dead cells. In the case
of diatoms, this effect is exacerbated by the absence of the mechanisms regulating
(typically by lowering) the density of the protoplast in living cells (Raven & Waite,
2004). This effect could be multiplicative with the effect of nutrient deficiency in
enhancing production of extracellular polysaccharides by algal cells (Fogg & Westlake,
1955; Hellebust, 1974; Wetz & Wheeler, 2003; but see Engel, 2002; Engel et al, 2004).
These exopolysaccharides are important in the flocculation of particles in sea water
(Engel et al., 2004), which increases their sinking rate for a given density by decreasing
the surface area per unit volume (Stokes' law). While terrigenous mineral, e.g. clay,
particles are also important as mineral ballast in the sinking of particulate organic
matter in marine snow, the double effect of nutrient deficiency, by increasing the
amount of biomineral ballast and the amount of floe-forming extracellular poly-
saccharides, could enhance not only the sinking of particulate organic matter in the
'biological pump', but also the removal of Fe, N, P and Zn from already nutrient-
deficient surface waters. A further feedback on nutrient removal as a result of low
nutrient availability is the capacity of the extracellular polysaccharides in the floes
to bind Fe and Zn from sea water (Engel et al., 2004). While Passow (2004) suggests
that the organic C fluxes may drive mineral particle fluxes rather than vice versa, it

SGM symposium 65

Fe, N, P and Zn cycling in the ocean 265

seems inescapable that mineral particles increase the density, and sinking rate, of
organic particles (see Klaas & Anchor, 2002; Boyd et al., 2004; Jickells et ai, 2005).

CONCLUSIONS

Fe, N and P have important roles in constraining marine primary productivity; the role
of Zn in limiting the growth of primary producers is less clearly established. There are
very substantial ecological, physiological and biochemical interactions among these
nutrients, and also interactions with the supply of other resources. Decreased
availability of these nutrient elements also potentially increases the sedimentation of
particulate material, including these four nutrient elements; further work is needed
to establish the significance of the potential feedbacks.

ACKNOWLEDGEMENTS

The work of K. B. and E. G. on inorganic C assimilation by diatoms is funded by a grant to
R. C. L. and J.A.R. from the Natural Environment Research Council UK, and that of M. M. on
mechanisms of inorganic C assimilation in coccolithophores is funded by the Biotechnology and
Biological Sciences Research Council UK. J. B. s work on algal photosynthesis is funded by the
Australian Research Council. Discussion with Professor Geoff Codd was very helpful.

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SGM symposium 65

Mechanisms and environmental
impact of microbial metal
reduction

Jonathan R. Lloyd

The Williamson Research Centre for Molecular Environmental Studies and the School of Earth,
Atmospheric and Environmental Sciences, University of Manchester, Manchester M1 3 9PL, UK

INTRODUCTION

Although it has been known for over a century that micro-organisms have the potential
to reduce metals, more recent observations showing that a diversity of specialist
bacteria and archaea can use such activities to conserve energy for growth under
anaerobic conditions have opened up new and fascinating areas of research with
potentially exciting practical applications (Lloyd, 2003). Micro-organisms have
also evolved metal-resistance processes that often incorporate changes in the oxida-
tion state of toxic metals. Several such resistance mechanisms, which do not support
anaerobic growth, have been studied in detail by using the tools of molecular biology.
Three obvious examples include resistance to Hg(II), As(V) and Ag(II) (Bruins et aL,
2000). The molecular bases of respiratory metal-reduction processes have not, however,
been studied in such fine detail, although rapid advances are expected in this area
with the imminent availability of complete genome sequences for key metal-reducing
bacteria, in combination with genomic, proteomic and metabolomic tools. This
research is being driven forward both by the need to understand the fundamental
basis of a range of biogeochemical cycles, and also by the possibility of harnessing
such activities for a range of biotechnological applications. These include the bio-
remediation of metal-contaminated land and water (Lloyd 6c Lovley, 2001), the
oxidation of xenobiotics under anaerobic conditions (Lovley 6c~ Anderson, 2000),
metal recovery in combination with the formation of novel biocatalysts (Yong et aL,
2002a) and even the generation of electricity from sediments (Bond et aL, 2002). The
aim of this review is to give an overview of the range of metals (and metalloids) reduced
by micro-organisms, the mechanisms involved and the environmental impact of

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

274 J. R. Lloyd

such transformations. Where appropriate, possible applications for these processes will
also be discussed.

REDUCTION OF Fe(lll) AND Mn(IV)

A wide range of archaea and bacteria are able to conserve energy though the reduction
of Fe(III) (ferric iron) to Fe(II) (ferrous iron). Many of these organisms are also able to
grow through the reduction of Mn(IV) to Mn(II). The environmental relevance of
Fe(III) and Mn(IV) reduction has been well-documented (Thamdrup, 2000), whilst
geochemical and microbiological evidence suggests that the reduction of Fe(III) may
have been an early form of respiration on Earth (Vargas et al., 1998). Some workers
have even proposed Fe(III) reduction as a candidate process for the basis of life on
other planets (Nealson 6c Cox, 2002). On modern Earth, Fe(III) can be the dominant
electron acceptor for microbial respiration in many subsurface environments (Lovley
& Chapelle, 1995). As such, Fe(III)-reducing communities can be responsible for
the majority of organic matter oxidized in such environments. Recent studies have
shown that a range of important xenobiotics that contaminate aquifers can also be
degraded under anaerobic conditions by Fe(III)-reducing micro-organisms (Lovley,
1997; Anderson et ai, 1998; Lovley &; Anderson, 2000). Transformations of inorganic
contaminants are also possible, and Fe(III)-reducing micro-organisms can also have an
impact on the fate of other high-valency contaminant metals through direct enzymic
reduction or via indirect reduction catalysed by biogenic Fe(II).

Focusing on the enzymic transformations of Fe(III) and Mn(IV), these organisms
can also influence the mineralogy of sediments through the reductive dissolution of
insoluble Fe(III) and Mn(IV) oxides (Fig. la). These processes can result in the release
of potentially toxic levels of reduced Fe(II) and Mn(II), and also trace metals that were
bound by the host Fe(III) or Mn(IV) minerals. Depending on the chemistry of the water,
a range of reduced minerals can also be formed, including, for Fe(III) reduction,
magnetite (Fe 3 4 ), siderite (FeC0 3 ) and vivianite [Fe 3 (P0 4 ) 2 . 8H 2 0] (see Fig. lb-d),
resulting in a change in structure of the sediments. There is a considerable level of
interest in these end products of Fe(III) reduction, as they are nanoscale and have
properties that may make them useful for a range of biotechnological processes. For
example, large quantities of regular-shaped, nanosized (5-10 nm) crystals of the
ferromagnetic mineral magnetite (Fe 3 4 ) are formed when Fe(II)-reducing bacteria
respire by using Fe(III) oxides in laboratory culture (Fig. 2). By controlling and
manipulating this process of biomineral production, a low-cost, low-energy, environ-
mentally friendly method of manufacture of nanoparticles could be developed to
replace existing practices. Of particular appeal is the potential to convert natural and
waste Fe(III) oxides (e.g. from mining/water industries), which are bulky and difficult
to handle, to a high-value product that is easy to process. Indeed, ferrite spinels such as

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Microbial metal reduction 275

Acetate

Fig. 1. Mechanism of Fe(lll) reduction (a) and the biogenic Fe(ll)-bearing minerals magnetite (b),
vivianite (c) and siderite (d). Bars, 50 nm (b); 10 pm (c); 20 pm (d). Figures were generously provided by
M. Wilkins, V. Coker, F. Islam and L. Adams (University of Manchester, UK).

magnetite have low coercivity (i.e. low power needed to magnetize or demagnetize
them), high permeability, high magnetic saturation and low conductivity. These
properties make them ideal for use in ultrahigh-density data-storage media, frequency-
selective circuits, radio-receiver antennae, microwave waveguides and other high-
frequency devices. Indeed, recent work from our laboratory has also demonstrated that
'designer' magnets can be made by Fe(III) -reducing bacteria by incorporating other
transition metals into the spinel structure in place of Fe, in some cases enhancing
the magnetic properties of the biomineral (Coker et al., 2004).

Diversity of Fe(lll)-reducing organisms

The first organisms were shown to conserve energy for growth through the reduction of
Fe(III) [and Mn(IV)] in the 1980s. These model organisms were Shewanella oneidensis
(formerly Alteromonas putrefaciens and then Shewanella putrefaciens) and Geobacter
metallireducens (formerly strain GS-15) (Lovley et al., 1987, 1989a; Myers & Nealson,
1988). Earlier studies had focused on organisms that grow predominantly via ferment-
ation of sugars such as glucose, with metals utilized as minor electron acceptors
(Roberts, 1947); typically, <5 % of the reducing equivalents is used for metal reduction
(Lovley, 1991).

Over the last 20 years, numerous organisms have been isolated that can grow by using
Fe(III) as an electron acceptor; more than 90 are listed in a recent review (Lovley et al.,
2004). In many freshwater subsurface environments, the most abundant seem to be

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276 J. R. Lloyd

Oxygen

Octahedral sites

Tetrahedral sites

Fig. 2. Production (a) and structure (b) of magnetite produced by Geobacter sulfurreducens.
Magnetite exhibits a classical spinel crystalline structure, with the oxygen ions forming a compact,
face-centred cubic assembly and the iron cations occupying octahedral and tetrahedral sites.
Figures were generously provided by V. Coker and R. Pattrick (University of Manchester, UK).

relatives of Geobacter metallireducens and fall within the family Geobacteraceae, in
the delta subdivision of the Proteobacteria (e.g. Rooney-Varga etaL, 1999; Snoeyenbos-
West et ai, 2000; Roling et al., 2001; Stein et al., 2001; Holmes et al., 2002; Islam et al.,
2004). This group comprises the genera Geobacter, Desulfuromonas, Desulfuromusa
and Pelobacter (Lovley et al., 2004). These organisms, with the exception of Pelobacter
species, are able to oxidize a wide range of organic compounds completely, including
acetate, when respiring by using Fe(III); Pelobacter species are more restricted in the
range of electron donors utilized, although they can couple hydrogen oxidation to
metal reduction. Some members of the family Geobacteraceae are also able to use

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Microbial metal reduction 277

aromatic compounds, including toluene, phenol and benzoate, as electron donors
for metal reduction. This is in contrast to Shewanella oneidensis and close relatives
in the gamma subdivision of the Proteobacteria (a range of Shewanella, Ferrimonas
and Aeromonas species) that are generally able to use only a restricted range of small
organic acids and hydrogen as electron donors for Fe(III) and Mn(IV) reduction. The
full range of other prokaryotes able to reduce Fe(III) is extensive and increasing
steadily, and crosses a wide range of environments (including extremes of pH and
temperature) and, indeed, taxonomic groupings. Although Shewanella and Geobacter
species have been the most intensively studied model Fe(III)-reducing bacteria, the
reader is referred to an excellent overview of the wide range of currently identified
Fe(III) -reducing bacteria (Lovley et al., 2004). Perhaps one of the most interesting new
isolates that can grow through the reduction of Fe(III) is archaeal 'strain 121', which
has pushed the upper temperature limit for life to 121 °C (Kashefi & Lovley, 2003).

Mechanisms of Fe(lll) and Mn(IV) reduction: electron transfer
to insoluble minerals

The mechanisms of Fe(III) reduction and, to a lesser degree, Mn(IV) reduction have
been studied in most detail in Shewanella oneidensis and Geobacter sulfur reducens.
Indeed, research on these organisms has been given added impetus through the
availability of their genome sequences (available at http://www.tigr.org) and suitable
genetic systems for the generation of deletion mutants for both of these organisms
(Myers & Myers, 2000; Coppi et al., 2001). Although the terminal reductase has yet to
be identified unequivocally in either organism, the involvement of c-type cytochromes
has been implicated in electron transport to Fe(III) and Mn(IV) by several studies
(Myers & Myers, 1993, 1997; Gaspard et al., 1998; Magnuson et al., 2000; Beliaev et
al., 2001; Lloyd et al., 2003). In some examples, activities have also been localized to
the outer membrane or surface of the cell, consistent with a role in direct transfer of
electrons to Fe(III) and Mn(IV) oxides that are highly insoluble at circumneutral pH
(Myers & Myers, 1992, 2001; Gaspard et al., 1998; DiChnstina et al., 2002; Lloyd et al.,
2002). In addition to the proposed direct transfer of electrons to Fe(III) and Mn(IV)
minerals, soluble 'electron shuttles' are also able to transfer electrons between metal-
reducing prokaryotes and the mineral surface. This mechanism alleviates the
requirement for direct contact between the micro-organism and mineral. For example,
humics and other extracellular quinones are utilized as electron acceptors by Fe(III)-
reducing bacteria (Lovley et al., 1996) and the reduced hydroquinone moieties are able
to transfer electrons abiotically to Fe(III) minerals. The oxidized humic is then avail-
able for reduction by the micro-organism, leading to further rounds of electron
shuttling to the insoluble mineral (Nevin 6c Lovley, 2002). Very low concentrations of
an electron shuttle, e.g. 100 nM of the humic analogue anthraquinone-2,6-disulfonate
(AQDS), can rapidly accelerate the reduction of Fe(III) oxides (Lloyd et al., 1999a) and

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278 J. R. Lloyd

possibly other insoluble metal oxides, such as Mn(IV). The environmental significance
of such processes, however, remains to be confirmed. The secretion of soluble electron
shuttles by actively respiring Fe(III) and Mn(IV) reducers has also been proposed for
both Shewanella oneidensis and Geobacter sulfurreducens, and remains hotly debated
in Geobacter species. Early studies suggested the release of a small, soluble, otype
cytochrome by Geobacter sulfurreducens (Seeliger et al., 1998), but more recent studies
have suggested that this protein is not an effective electron shuttle (Lloyd et al., 1999a).
Studies have also suggested that a small, quinone-containing, extracellular electron
shuttle is released by Shewanella oneidensis and may also promote electron transfer to
Fe(III) and Mn(IV) minerals (Newman 6c Kolter, 2000). Finally, an important new
discovery was made recently when it was shown that Geobacter metallireducens
synthesized pili and flagella when grown on insoluble Fe(III) or Mn(IV) minerals, but
not soluble forms of the metals (Childers et al., 2002). These results suggest that
Geobacter species sense when soluble electron acceptors are depleted and synthesize the
appropriate appendages that allow movement to Fe(III) and Mn(IV) minerals and
subsequent attachment. Pili may also play a direct role in electron transfer to the
extracellular electron acceptor (Reguera et al., 2005).

REDUCTION OF OTHER TRANSITION METALS

In addition to Fe(III) and Mn(IV), dissimilatory metal-reducing prokaryotes are able
to respire by using a wide range of transition metals, including high-valency ions
of vanadium, chromium, molybdenum, cobalt, palladium, silver, gold and mercury. In
many cases, reduction leads to a dramatic change in solubility, can potentially lead
to the precipitation of metal-containing ores and may also offer routes to the bio-
remediation of metal-contaminated water.

Vanadium reduction

Early studies showed V(V) reduction by 'Micrococcus lactilyticus', Desulfovibrio
desulfuricans and Clostridium pasteurianum (Woolfolk & Whiteley, 1962), followed by
the observation that the ability to reduce V(V) was widespread amongst soil bacteria
and fungi (Bautista & Alexander, 1972). More recent work has focused on two
pseudomonads: 'Pseudomonas vanadiumreductans' and 'Pseudomonas isachenkovii\
isolated from a waste stream from a ferrovanadium factory and sea water, respectively
(Yurkova 6c Lyalikova, 1991). Anaerobic cells were able to utilize a wide range of
electron donors, including hydrogen, sugars and amino acids. V(V) was reduced to
blue-coloured V(IV) and possibly further to V(III), the latter indicated by the formation
of a black precipitate and by its reaction with the reagent Tairon (Yurkova 6c Lyalikova,
1991). Geobacter metallireducens also reduces V(V) and this form of metabolism has
been suggested as a mechanism for remediating vanadium-contaminated water (Ortiz-
Bernad <?£#/., 2004a).

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Microbial metal reduction 279

Chromium reduction

The widespread use of chromium in the metal industries and subsequent contamin-
ation problems have led to a lot of interest in this metal. Although trace quantities are
required for some metabolic activities, e.g. glucose and lipid metabolism, chromium
is considered toxic and is designated a priority pollutant in many countries. Two
oxidation states dominate: Cr(VI) is the most toxic and mobile form encountered
commonly, with Cr(III) being less soluble and less toxic. Indeed, Cr(III) is considered
1000 times less mutagenic than Cr(VI) (Wang, 2000). Current treatment involves
reduction of Cr(VI) to Cr(III) by using chemical reductants at low pH, followed by
adjustment to near-neutral pH and subsequent precipitation of Cr(III). Recent studies,
however, have shown that micro-organisms can also reduce Cr(VI) efficiently at circum-
neutral pH, and could be used to treat Cr(VI)-contaminated water. In most cases,
Cr(VI) reduction does not support anaerobic growth, although a few publications have
suggested that conservation of energy is possible through this form of anaerobic
metabolism (e.g. Tebo 6c Obraztsova, 1998).

A wide range of facultative anaerobes are able to reduce Cr(VI) to Cr(III), including
Escherichia coli, pseudomonads, Shewanella oneidensis and Aeromonas species
[see Wang (2000) for a more exhaustive list]. Anaerobic conditions are generally
required to induce maximum activity against Cr(VI), but some enzyme systems operate
under aerobic conditions, e.g. the soluble NAD(P)H-dependent reductases of l Pseudo-
monas ambigucf G-l (Suzuki et al., 1992) and Pseudomonas putida (Park et al., 2000).
Obligate anaerobes are also able to reduce Cr(VI) enzymically and anaerobic growth
coupled to Cr(VI) reduction has been reported for a sulfate-reducing bacterium (Tebo
& Obraztsova, 1998). The reduction of Cr(VI) by sulfate-reducing bacteria is parti-
cularly well-studied (e.g. Lloyd et al., 2001) and has been shown to be catalysed by
cytochrome c 3 (Lovley & Phillips, 1994). Other studies have also implicated the involve-
ment of cytochromes in Cr(VI) reduction by bacteria: cytochrome c in Enterobacter
cloacae (Wang et al., 1989) and cytochromes b and d in Escherichia coli (Shen & Wang,
1993). Environmental factors that affect Cr(VI) reduction include competing electron
acceptors, pH, temperature, redox potential and the presence of other metals (Wang,
2000). A recent study has also demonstrated that the presence of complexing agents
can promote Cr(VI) reduction, possibly through protection of the metal reductase by
chelation of Cr(III) or intermediates formed (Mabbett et al., 2002). The type of electron
donor supplied can also have an effect on the rate and extent of Cr(VI) reduction.
Optimal electron donors, in keeping with other dissimilatory metal-reduction pro-
cesses described in this review, are low-molecular-mass carbohydrates, amino acids and
fatty acids. Finally, indirect mechanisms that also promote Cr(VI) reduction in con-
taminated sediments are catalysed by biogenic sulfide (Smillie et al., 1981; Fude et al.,
1994) and Fe(II) (Fendorf & Li, 1996). Experiments using contaminated sediments from

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280 J. R. Lloyd

Norman, OK, USA, have, however, confirmed that indirect mechanisms may not always
be the critical control on chromium solubility, with direct enzymic Cr(VI) reduction by
a consortium of methanogens being implicated (Marsh et al., 2000).

Molybdenum reduction

Although microbial Mo (VI) reduction could play a role in the molybdenum cycle, for
example leading to the concentration of insoluble molybdenum in anaerobic marine
sediments and reduction spots in rocks (Lovley, 1993), comparatively few studies have
addressed this process. Early work suggested that 'Pseudomonas guillermondi? and a
Micrococcus species could reduce Mo (VI) to molybdenum blue (Bautista 6c Alexander,
1972). More recently, similar activities have been identified in cultures of a molyb-
denum-resistant Enterobacter species (Ghani et al., 1993). The organism was grown
under anaerobic conditions in glucose-containing medium supplemented with 200 mM
Mo (VI). Reduction of Mo (VI) was accompanied by a change in colour, as Mo(V)
formed and complexed with phosphate in the medium to form methylene blue (Ghani
et al., 1993). The use of metabolic inhibitors suggested that the electrons for Mo (VI)
reduction were derived from the glycolytic pathway, whilst NADH functioned as an
electron donor in broken cells in vitro (Ghani etal., 1993). The ability to reduce Mo (VI)
has also been identified in pre-grown cells of the sulfate-reducing bacterium Desulfo-
vibrio desulfuricans, both through direct enzymic mechanisms and indirectly via sulfide
under sulfate-reducing conditions (Tucker et al., 1997). Cells of Desulfovibrio desulfur-
icans immobilized in a bioreactor have also been used to remove Mo (VI) from solution
at high efficiency (Tucker et al., 1998). The organism was unable to grow using Mo (VI)
as an electron acceptor (Tucker et al., 1997) and it is unlikely that actively growing
cultures of sulfate-reducing bacteria would play a direct role in reducing high con-
centrations of Mo (VI) in the environment, given the toxicity of molybdate to these
organisms (Oremland & Capone, 1988).

Cobalt reduction

The reduction of Co (III) has received recent attention because radioactive 60 Co can be a
problematic contaminant at sites where radioactive waste has been stored. Co (III) is
especially mobile when complexed with EDTA and several studies have focused on the
ability of Fe (III) -reducing bacteria to retard the mobility of the metal through reduction
to Co(II) (Caccavo et al., 1994; Gorby et al., 1998). The Co(II) formed does not asso-
ciate strongly with EDTA [it is over 25 orders of magnitude less thermodynamically
stable than Co (III) EDTA] and absorbs to soils, offering the potential for in situ immo-
bilization of the metal in contaminated soils. The precise mechanisms of dissimilatory
Co (III) reduction remain to be investigated.

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Microbial metal reduction 281

Palladium reduction

The reduction of soluble Pd(II) to insoluble Pd(0) has also attracted interest, as this
enzymic process may be used to recover palladium from industrial catalysts (Lloyd et
al., 1998a) and may also be used to synthesize nanoscale bioinorganic catalysts of
considerable commercial potential (Yong et al., 2002a). Interest in this area is driven
by the widespread use of platinum-group metals (PGMs), including palladium, in
automotive catalytic converters required to reduce gaseous emissions, and problems
associated with their recycling. With approximately 5 g PGM per catalytic converter,
the consumption of PGMs was altogether 70312-5 kg in 1994, with only 11 250 kg
recovered (Lloyd et al., 1998a). The lifetime of a catalytic converter is only approx-
imately 80000 km (50000 miles), although many fail sooner, and future shortages and
higher prices may be predicted. Chemical and electrochemical treatments are made
difficult by complex solution chemistry.

Initial experiments aimed at reducing and recovering palladium were based on the use
of Desulfovibrio desulfuricans because it is active against a wide range of metals,
including Fe (III), Mn(IV), U(VI), Cr(VI) and Tc (VII), via hydrogenase or cytochrome c 3
(Lloyd et al., 1998a). Cells were able to reduce 0-5 mM Pd(II) [as Pd(NH 3 ) 4 Cl] with a
range of electron donors, including pyruvate, formate and H 9 . Although the enzyme
responsible for Pd(II) reduction has not been identified, the involvement of a peri-
plasmic hydrogenase is implicated by the use of hydrogen as electron donor and
inhibition by treatment with 0-5 mM Cu 2+ (Lloyd et ai, 1998a). Transmission electron
microscopic studies, in combination with energy-dispersive X-ray microanalysis,
confirmed precipitation in the periplasm, with X-ray diffraction studies confirming
reduction to Pd(0). More recent studies have focused on the recovery of Pd(II) at a range
of pH values and also from acid (aqua regia) leachates from spent automotive catalysts
(Yong et al., 2002b). Inhibition by chloride ions was reported and this may therefore
necessitate leaching from catalysts to minimize the formation of PdCl^ - prior to
bioreduction and recovery. Delivery of reducing power to an immobilized biocatalyst
has also been studied in a novel electrobioreactor (Yong et ai, 2002b) containing a
biofilm of Desulfovibrio desulfuricans immobilized on a Pd-Ag membrane that trans-
ported atomic hydrogen to the cells, minimizing loss of gaseous hydrogen. Pd(0)
recovered in the electrobioreactor proved a better catalyst that its chemical counterpart,
as determined by hydrogen liberation from hypophosphite (Yong et al., 2002a) and
reduction of several target organics (L. E. Macaskie, personal communication).

Gold and silver reduction

It has been argued that Fe(III)-reducing bacteria may play a role in the deposition
of gold ores, as these organisms are present in high- and moderate-temperature
sedimentary environments where gold deposits have been recovered (Kashefi et al.,

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282 J. R. Lloyd

2001). Indeed, several dissimilatory Fe(III)-reducing bacteria and archaea, including
the hyper thermophilic archaea Pyrobaculum islandicum and Pyro coccus furiosus, the
hyperthermophilic bacterium Thermotoga maritima and the mesophilic bacteria
Shewanella algae and Geovibrio ferrireducens, were shown to reduce Au(III) (as gold
chloride) to insoluble Au(0) (Kashefi et aL, 2001). The ability to reduce Au(III) seems
to be species-specific, and closely related organisms with similar activities against a
range of other metals have differing activities against Au(III) (Kashefi et aL, 2001). For
example, unlike Pyrobaculum islandicum, a close relative, Pyrobaculum aeropbilum, is
unable to reduce Au(III). Also, there is an obligate requirement for hydrogen as an
electron donor in organisms that can reduce Au(III), suggesting the involvement of a
hydrogenase. Given the direct reduction of other metals by hydrogenase (Lloyd et aL,
1997), it is tempting to hypothesize that hydrogenases may play a direct role in Au(III)
reduction.

Microbial reduction of Ag(I) has also been studied, but in little detail. Early reports
noted that the reduction of Ag(I) may account for resistance to silver in some micro-
organisms (Belly 6c Kydd, 1982), but more recent studies have uncovered alternative
strategies for resistance to Ag(I) in organisms isolated from hospital burns wards, where
silver may be used as a biocide (Gupta et aL, 1999), with a recent review of this new area
presented by Silver (2003). Several studies, including that by Fu et al. (2000), have shown
biosorption of Ag(I) to the surface of cells (in this case, a Lactobacillus species),
followed by reduction to Ag(0). Here, the mechanism for Ag(I) reduction remains
unknown.

Mercury reduction

A well-studied metal-resistance system is encoded by genes of the mer or mercury-
resistance operon, which relies upon the reduction of Hg(II) (mercuric ions). Here,
Hg(II) is transported into the cell via the MerT transporter protein and detoxified by
reduction to relatively non-toxic, volatile elemental mercury by an intracellular
mercuric reductase (MerA) (Hobman & Brown, 1997). The biotechnological potential
of this process has been described recently, focusing on the use of mercury-resistant
bacteria and the proteins that they encode (Lloyd et aL, 2004). Applications include the
bioremediation of mercury-contaminated water and the development of Hg(II)-
detecting biosensors. Finally, in addition to the MerA-mediated mechanism of mercury
reduction, other enzymes are also able to reduce Hg(II). A novel, Fe 2+ -dependent
mechanism for mercury reduction has been characterized in the membrane fraction of
Tbiobacillus ferrooxidans, which may involve cytochrome c oxidase (Iwahori et aL,
2000), and c-type cytochromes of Geobacter metallireducens also reduce Hg(II) (Lovley

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Microbial metal reduction 283

et al., 1993). Results from our laboratory have also shown that whole cells of Geobacter
sulfur re due ens are able to reduce Hg(II) without the involvement of mer mercury-
resistance genes (N. Law & J. R. Lloyd, unpublished data).

REDUCTION OF METALLOIDS

Arsenic reduction: a role in mass poisoning worldwide?

Contamination of groundwaters, abstracted for drinking and irrigation, by sediment-
derived arsenic threatens the health of tens of millions of people worldwide, most
notably in Bangladesh and West Bengal (Chakraborti et al., 2002; Smedley &C Kinni-
burgh, 2002). Despite the calamitous human-health impacts arising from the extensive
use of arsenic-enriched groundwaters in these regions, the mechanisms of arsenic
release from sediments remain poorly characterized and have been a topic of intense
debate (Nickson et al., 1998; Chowdhury et al., 1999; Oremland & Stolz, 2003).
However, recent results have suggested that metal-reducing bacteria may be a major
cause of this humanitarian catastrophe.

The immediate source of arsenic in these groundwaters in Bengal is widely considered
to be the host sediments, which are transported by the rivers Ganges, Brahmaputra and
Meghna and derived from the weathering of the Himalayas within the l-5x 10 6 km 2
catchment area of these three great river systems. Depending upon the sediment type,
arsenic typically occurs at concentrations of 2-100 parts per million and is found in,
and adsorbed onto, a variety of mineralogical hosts, including hydrated ferric oxides,
phyllosilicates and sulfide minerals (Smedley & Kinniburgh, 2002). The mechanism of
arsenic release from contaminated sediments remains controversial, but microbially
mediated release of arsenic from hydrated ferric oxides is gaining the consensus as the
dominant mechanism for the mobilization of arsenic into groundwater systems of the
Ganges delta. For example, we recently provided direct evidence for the role of
indigenous metal-reducing bacteria in the formation of toxic, mobile As(III) in sedi-
ments from the Ganges delta (Islam et al., 2004). In this study, sediment samples were
collected at a depth of 13 m from a site in West Bengal known to have relatively high
concentrations of arsenic in the groundwater. Arsenic mobilization was optimal under
anaerobic conditions, and addition of acetate to anaerobic sediments, as a proxy for
organic matter and a potential electron donor for metal reduction, resulted in stimu-
lation of microbial reduction of Fe(III), followed by As(V) reduction and release
of mobile As (III). Microbial communities responsible for metal reduction and arsenic
mobilization in the stimulated anaerobic sediments were analysed by using molecular
(PCR-based) and cultivation-dependent techniques. Both approaches confirmed an
increase in numbers of Fe (III) -reducing bacteria and suggested a vital role for metal-

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284 J. R. Lloyd

reducing bacteria in mediating arsenic release. Indeed, the PCR studies showed that
the microbial communities in these sediments were dominated by Geobacter species.
We have also recently obtained similar results using Cambodian sediments liberating
arsenic into groundwater (H. A. L. Rowland, R. L. Pederick, D. A. Polya, R. D. Pancost,
B. E. van Dongen, A. G. Gault, J. M. Charnock, D. J. Vaughan & J. R. Lloyd, unpubl-
ished data). However, we have also noted recently that Geobacter species (e.g.
Geobacter sulfurreducens) cannot reduce As(V) enzymically and may actually promote
the retention of arsenic in sediments, through the formation of Fe (II) minerals that also
sorb the metalloid (F. S. Islam, R. L. Pederick, A. G. Gault, L. K. Adams, D. A. Polya,
J. M. Charnock & J. R. Lloyd, unpublished data). These results also show that the
reduction of Fe(III) alone is not sufficient to mobilize sorbed arsenic, suggesting that
specialized As(V)-respiring bacteria play a very important role in the reductive
mobilization of arsenic as mobile, toxic As (III).

Several organisms capable of growing through dissimilatory reduction of As(V) have
now been isolated, although none from south-east Asian aquifers. Chrysiogenes arsena-
tes is a strict anaerobe that was isolated from wastewater from a gold mine (Macy et aL,
1996). This organism contains a periplasmic arsenate reductase consisting of two sub-
units (masses 87 and 29 kDa) that contain molybdenum, iron, sulfur and zinc co-factors
(Krafft & Macy, 1998). Sulfuro spirillum arsenophilum (previously strain MIT 13 T ) is
a Gram-negative, vibrioid, microaerobic, sulfur-reducing bacterium isolated from
arsenic-contaminated watershed sediments in eastern Massachusetts, USA. It is a very
close relative of Sulfuro spirillum barnesii, which can also reduce As(V), and has a
broad activity against metals, including Fe(III) and Se(VI) (see sections on these metals/
metalloids). A recent review has mentioned the preliminary purification and characteri-
zation of an arsenate reductase from this organism, which constituted a trimeric
complex of 120 kDa, consisting of subunits of 65, 31 and 22 kDa (Oremland & Stolz,
2000). A Gram-positive, sulfate-reducing bacterium, Desulfotomaculum auripig-
mentum, has also been described, which reduces As(V) followed by sulfate, resulting in
the formation of orpiment (As 2 S 3 ) (Newman et aL, 1997).

Reduction of As(V) to As(III) by the ArsC reductase also forms the basis of a well-
studied microbial arsenic-resistance mechanism, preceding efflux of As (III) from the
cell (Mukhopadhyay et aL, 2002). It should be noted, however, that ArsC-mediated
As(V) reduction does not support microbial growth and there is currently little evidence
linking this mechanism of As(V) reduction to the biogeochemical cycling of arsenic.
Finally, on this note, it is worth mentioning that a novel organism has recently been
isolated that is able to use As (III) as an electron donor for aerobic growth (Santini et aL,
2000), thus closing the biological arsenic cycle.

SGM symposium 65

Microbial metal reduction 285

Finally, As(V) is actually a minor (but important) electron acceptor for anaerobic
respiration in aquifers. However, in some environments, such as Mono Lake, CA, USA,
As(V) can be the dominant electron acceptor for carbon oxidation (Oremland et aL,
2000) and here, the role of microbial metabolism in controlling arsenic speciation is far
clearer. This fascinating environment and the organisms that it supports have been
reviewed recently (Oremland et al., 2004).

Reduction of Se(VI) and Se(IV) and other group VIB elements

In contrast to As(V) reduction, the biotransformation of Se(VI) and Se(IV) to relatively
unreactive Se(0) results in its removal from water. Several studies have demonstrated
that these transformations can be catalysed by microbes (Oremland & Stolz, 2000). For
example, the ability to reduce Se(VI) is widespread in sediments, with biological
reduction demonstrated unequivocally in 10 out of 11 sediment types (Steinberg &
Oremland, 1990) . Also, Se(VI) is not reduced chemically under physiological conditions
of pH and temperature, and Se(VI) reduction is inhibited by autoclaving of sediments.

Organisms that are known to reduce Se(VI) enzymically include Wolinella succinogenes
(Tomei et aL, 1992), Desulfovibrio desulfuricans (Tomei et aL, 1995), Pseudomonas
stutzeri (Lortie et aL, 1992), Enterobacter cloacae (Losi & Frankenberger, 1997) and
Escherichia coli (Avazeri et aL, 1997). In these examples, Se(VI) reduction does not
support growth and seems to be incidental to the physiology of the organism. In at least
one organism (Escherichia coli), the involvement of broad-specificity nitrate reductases
is implicated by biochemical studies (Avazeri et aL, 1997). In addition to these rather
non-specific reactions, specialist organisms are known to conserve energy through
Se(VI) reduction, including Thauera selenatis (Macy & Lawson, 1993), Sulfuro-
spirillum barnesii [originally 'Geo spirillum barnesif strain SES-3 (Stolz et aL, 1997)]
and two bacilli (Bacillus arseniciselenatis and Bacillus selenitireducens) , both isolated
from Mono Lake, CA, USA (Switzer Blum et aL, 1998). Of these four model organisms,
the mechanism of Se(VI) reduction is best understood in Thauera selenatis (Schroder
et aL, 1997). A periplasmic complex of approximately 180 kDa (with subunits of
masses 96, 40 and 23 kDa) has been characterized and shown to contain molybdenum,
iron and acid-labile sulfur. Specificity for Se(VI) is high, with a K m of 16 jiM. The
enzyme was unable to reduce nitrate, nitrite, chlorate, chlorite or sulfate. Biochemical
studies are not so advanced in Sulfuro spirillum barnesii, although the enzyme activity
contrasts with that characterized in Thauera selenatis, as it is localized in the membrane
fraction and may have a wider substrate specificity (Stolz et aL, 1997).

Tellurite (TeOj~) reduction has also been studied in several organisms, mainly in the
context of resistance to this toxic oxyanion. Indeed, the antibacterial properties of
Te(IV) have been known for more than 70 years: in the pre-antibiotic era, Te(IV) was

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286 J. R. Lloyd

used to treat a range of bacterial infections and Te(IV) remains an ingredient of several
selective media (e.g. for verocytotoxigenic Escherichia coli 0157). Basal levels of
resistance to toxic Te(IV) have been attributed to the activity of a membrane-bound
nitrate reductase in Escherichia coli (Avazeri et aL, 1997). An additional Te(IV)
reductase was detected in the soluble fraction of anaerobically grown cells. Growth by
using Te(IV) as an electron acceptor was also reported in an engineered strain over-
expressing nitrate reductase, but was not thought to be physiologically relevant in wild-
type cells of Escherichia coli (Avazeri et aL, 1997). Rhodobacter sphaeroides has also
been reported to reduce Te(IV) (as well as oxyanions of selenium), with an absolute
requirement for a functional photosynthetic electron-transfer chain under photo-
synthetic (anaerobic) growth conditions, or functional cytochromes hc x and c 2 under
aerobic growth conditions (Moore 6c Kaplan, 1992). Again, metal reduction was
discussed in the context of resistance to the metalloids. Finally, plasmids are known to
encode several distinct resistance determinants for Te(IV) and, again, Te(IV) reduction
is implicated as the resistance mechanism, as elemental tellurium is deposited within
tellurite-resistant bacteria (Taylor, 1999). However, other mechanisms of resistance,
involving cysteine-metabolizing enzymes and methyl transferases, may be important
(Taylor, 1999). Finally, Te(IV) [and Se(VI)/Se(IV)] reduction and precipitation by
sulfate-reducing bacteria has also been reported, in the order Te(IV) >Se(VI) >Se(IV),
which is in contrast to that predicted by the redox potentials alone (Lloyd et aL, 2001).
To date, there have been no reports of microbial growth coupled to the reduction
of Te(IV) by non-genetically engineered micro-organisms.

REDUCTION OF ACTINIDES AND FISSION PRODUCTS AND THE
BIOREMEDIATION OF RADIOACTIVE WASTE

The release of radionuclides from nuclear sites and their subsequent mobility in the
environment is a subject of intense public concern and has prompted much recent
research on the environmental fate of key radionuclides (Lloyd & Renshaw, 2005a). The
major burden of anthropogenic environmental radioactivity is from the controlled
discharge of process effluents produced by industrial activities allied to the generation
of nuclear power, although significant quantities of natural and artificial radionuclides
were also released as a consequence of nuclear weapons testing in the 1950s and 1960s
via accidental release, e.g. from Chernobyl in 1986, and from the ongoing storage of
nuclear materials amassed over the last 60 years of nuclear activities. Indeed, the scale
of our nuclear legacy is enormous, including 120 Department of Energy sites in the
USA alone, and other facilities in Europe and the former USSR (Lloyd & Renshaw,
2005a). In several cases, storage has been compromised, leading to contamination of
trillions of litres of groundwater and millions of cubic metres of contaminated soil and
debris. The costs of cleaning up these sites are estimated to be in excess of a trillion US
dollars in the USA alone, and 50 billion pounds sterling in the UK. Given these high

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Microbial metal reduction 287

costs and the technical limitations of current chemical-based approaches, there has
been an unprecedented interest in the interactions of micro-organisms with key
radionuclides, in the hope of developing cost-effective bioremediation approaches
for decontamination of sediments and waters affected by nuclear waste (Lloyd et aL,
2004).

Because many radionuclides of concern are both redox-active and less soluble when
reduced, bioreduction offers much promise for controlling the solubility and mobility
of target radionuclides in contaminated sediments, e.g. the reduction of U(VI) (the
uranyl ion; UO^ + ) to U(IV) (uraninite; U0 2 ) (Lovley et aL, 1991; Lovley & Phillips,
1992b) or the reduction of the fission product Tc(VII) (the pertechnetate ion; TCO4) to
Tc(IV) (Tc0 2 ) (Lloyd et aL, 2000b). Several studies have also addressed the colonization
of radioactive environments (see Lloyd 6c Renshaw, 2005b) and it would seem that the
radioactive burden of several nuclear-waste types is not necessarily inhibitory to all
microbial life. For example, a recent study using pure cultures of bacteria proposed for
application in bioremediation programmes has addressed the toxicity of actinides,
metals and chelators. The model organisms tested include Deinococcus radiodurans,
Pseudomonas putida and Shewanella putrefaciens CN32 (Ruggiero et aL, 2005).
Actinides, including chelated Pu(IV), U(VI) and Np(V), inhibited growth at millimolar
concentrations, suggesting that actinide toxicity is primarily chemical (not radio-
logical) and that radiation resistance (e.g. in Deinococcus species) does not necessarily
ensure radionuclide tolerance. The author proposes that actinide toxicity will not
impede bioremediation using naturally occurring bacteria, although the toxicity of
these key radionuclides remains to be determined in sedimentary environments under
field conditions.

Uranium reduction

The first demonstration of dissimilatory U(VI) reduction was by Lovley et aL (1991),
who reported that the Fe (III) -reducing bacteria Geobacter metallireducens (previously
designated strain GS-15 T ) and Shewanella oneidensis (formerly Alteromonas putre-
faciens and then Shewanella putrefaciens) can conserve energy for anaerobic growth via
the reduction of LI (VI). It should be noted, however, that the ability to reduce U(VI)
enzymically is not restricted to Fe (III) -reducing bacteria. Other organisms, including a
Clostridium species (Francis, 1994) and the sulfate-reducing bacteria Desulfovibrio
desulfuricans (Lovley & Phillips, 1992a) and Desulfovibrio vulgaris (Lovley & Phillips,
1994), also reduce uranium, but are unable to conserve energy for growth via this
transformation. To date, the enzyme system responsible for U(VI) reduction has been
best studied in Desulfovibrio vulgaris. Purified tetrahaem cytochrome c 3 was shown to
function as a U(VI) reductase in vitro, in combination with hydrogenase, its physio-
logical electron donor (Lovley & Phillips, 1994). In vivo studies using a cytochrome c 3

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288 J. R. Lloyd

r«

ft?

'^Extracellular
'/' -U(IV)

*0i#

Periplasmic U(IV)

Acetate

CO-

U(VI)

soluble

(b)

U(V) "=<^ Disproportionation

unstable

iO.»

OMfi u(iv)

p _ ■ * * insoluble

Fig. 3. Reduction of U(VI) to insoluble U(IV) (electron-dense deposits) in thin sections of Geobacter
sulfurreducens, viewed by using transmission electron microscopy (a). U(VI) reduction by this organism
is via unstable U(V), which disproportionates to give insoluble U(IV) [and U(VI) for further reduction]
(b). Transmission electron microscopy by S. Glasauer (University of Guelph, Canada). Bar, 0-5 \xm.

mutant of the close relative Desulfovibrio desulfuricans strain G20 confirmed a role for
cytochrome c 3 in hydrogen-dependent U(VI) reduction, but suggested additional
pathways from organic electron donors to U(VI) that bypassed the cytochrome (Payne
et al., 2002). Similar cytochrome-mediated mechanisms have been proposed in
Geobacter species, whilst U(VI) reduction in a Shewanella putrefaciens strain shares
components of the nitrite-reducing pathway in this organism (Wade ck: DiChristina,
2000). The mechanism of U(VI) reduction by Geobacter sulfurreducens has also been
studied recently in detail by using X-ray absorbance spectroscopy, which showed the
formation of an unstable U(V) intermediate (Renshaw et al., 2005). This organism was
unable to reduce the stable analogue Np(V), suggesting that the further reduction of
U(V) is by disproportionation to U(IV) [with U(VI) generated available for further
reduction] (Fig. 3). This surprising level of specificity of hexavalent actinides illustrates

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Microbial metal reduction 289

a need for detailed investigations on the impact of micro-organisms on complex,
actinide-containing wastes.

Field studies on uranium bioreduction in situ

There has been a considerable level of interest in harnessing the metabolism of U(VI)-
reducing bacteria for the bioremediation of uranium-contaminated aquifers. For
example, recent studies focused on biostimulation of U(VI) -reducing bacteria at a
'Uranium Mill Tailings Remedial Action' (UMTRA) site in Rifle, CO, USA, through
the injection of an electron donor (acetate) into the subsurface (Anderson et al., 2003).
The decrease in soluble U(VI) was coincident with an increase in Fe(II) in the
groundwater and a significant enrichment of Geobacter species. However, after 39 days,
the composition of the microbial community began to change as sulfate-reducing
organisms dominated, and soluble U(VI) increased with a decrease in Fe(II), acetate and
sulfate and an accumulation of sulfite. Thus, the precise constituents of the microbial
communities present in the sediments clearly exert control on U(VI) speciation and
require careful optimization. The geochemistry of the groundwaters should also not
be overlooked. For example, Ca~ + cations at millimolar concentrations cause a signifi-
cant decrease in the rate and extent of bacterial U(VI) reduction by a range of metal-
reducing bacteria, suggesting that U(VI) is a less effective electron acceptor when
present as the Ca 2 U02(C0 3 )3 complex (Brooks et al., 2003). High nitrate concen-
trations can also inhibit U(VI) reduction by acting as a competing electron acceptor,
and may even promote reoxidation of reduced U(IV) (Istok et al., 2004). Mineralogical
constraints can also be important in controlling the end points for uranium bio-
remediation. For example, although much work has focused on the reduction of soluble
U(VI), the fate of sorbed U(VI), which can be appreciable in sediments, is also
potentially important. For example, although soluble U(VI) was reduced in a slurry
prepared from sediments from Rifle, CO, USA, sorbed U(VI) was not reduced and was
deemed 'not bioavailable' for microbial reduction (Ortiz-Bernad et al., 2004b).

Reduction of other actinides (plutonium and neptunium)

Although 2o8 U remains the priority pollutant in most medium- and low-level radio-
active wastes, other actinides, including 230 Th, 237 Np, 241 Pu and 241 Am, can also be
present (Macaskie, 1991; Lloyd & Macaskie, 2000). Th(IV) and Am (III) are stable
across most E h values encountered in radionuclide-contaminated waters, but the
potentials for Pu(V)/Pu(IV) and Np(V)/Np(IV), in common with that of U(VI)/U(IV),
are more electropositive than the standard redox potential of ferrihydrite/Fe~ + (approx.
V; Thamdrup, 2000). Thus, Fe(III)-reducing bacteria have the metabolic potential
to reduce these radionuclides enzymically or via Fe(II) produced from the reduction of
Fe(III) oxides. This is significant because the tetravalent actinides are amenable to
bioremediation, due to their high ligand-complexing abilities (Lloyd & Macaskie,

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290 J. R. Lloyd

2000), and are also immobilized in sediments containing active biomass (Peretrukhin
et al, 1996). Thus, although it is possible for Fe(III)-reducing bacteria to reduce and
precipitate actinides directly, e.g. the reduction of soluble U(VI) to insoluble U(IV) (see
above), some transformations do not result in formation of an insoluble mineral phase,
but in the formation of a cation more amenable to bioprecipitation. This is illustrated
when considering highly soluble Np(V) (NpO?), which was reduced to soluble Np(IV)
by Shewanella putrefaciens, with the Np(IV) removed as an insoluble phosphate
biomineral by a phosphate-liberating Citrobacter species (Lloyd et al, 2000a). This is in
sharp contrast to the case in Geobacter sulfur re due ens, which is unable to reduce
Np(V) (see above). Also, some studies have suggested that the reduction of Pu(IV) to
Pu(III) can be achieved by Fe (III) -reducing bacteria, although the Pu(III) was reported
to reoxidize spontaneously (Rusin et al, 1994). Although this may lead to solublization
of sediment-bound Pu(IV), it will yield a trivalent actinide that is also amenable to
bioremediation by using a range of microbially produced ligands (Lloyd & Macaskie,
2000). The biochemical basis of these transformations remains uncharacterized.

Technetium reduction

The fission product technetium is another long-lived radionuclide that is present in
nuclear waste and has attracted considerable recent interest. This is due to a combin-
ation of its mobility as the soluble pertechnetate ion [Tc(VII); TCO4], bioavailability as
an analogue of sulfate and long half-life (2-13 x 10 5 years) (Wildung et al, 1979). Like
Np(V), Tc(VII) has weak ligand-complexing capabilities and is difficult to remove from
solution by using conventional 'chemical' approaches. Several reduced forms of the
radionuclide are insoluble, however, and metal-reducing micro-organisms can reduce
Tc(VII) and precipitate the radionuclide as a low-valency oxide [Tc(IV); Tc0 2 ].

In an early study on Tc(VII) bioreduction, a novel phosphorimaging technique was used
to show reduction of the radionuclide by Shewanella putrefaciens and Geobacter
metallireducens, with similar activities subsequently detected in laboratory cultures of
Rbodobacter sphaeroides, Paracoccus denitrificans, some pseudomonads (Lloyd et al,
2002), Escherichia coli (Lloyd et al., 1997) and a range of sulfate-reducing bacteria
(Lloyd et al., 1998b, 1999b, 2001). Other workers have used this technique to show that
Thiobacillus ferrooxidans and Thiobacillus thiooxidans (Lyalikova & Khizhnyak,
1996) and the hyperthermophile Pyrobaculum islandicum (Kashefi & Lovley, 2000) are
also able to reduce Tc(VII). It should be stressed that Tc(VII) reduction has not been
shown to support growth in any of these studies, and seems to be a fortuitous
biochemical side reaction in the organisms studied to date. Recent work has also shown
that Tc(VII) can be reduced through indirect microbial processes via, for example,
biogenic sulfide (Lloyd et al, 1998b), Fe(II) (Lloyd et al, 2000b) or U(IV) (Lloyd et al,
2002). Tc(VII) reduction and precipitation by biogenic Fe(II) are particularly efficient

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Microbial metal reduction 291

and may offer a potentially useful mechanism for the remediation of technetium-
contaminated sediments containing active Fe (III) -reducing bacteria (Lloyd et al.,
2000b).

This latter point has been confirmed by several recent studies using a range of sediment
materials. In one study, sediments from the Humber Estuary, UK, were left to age and
exhibited a clear progression of terminal electron-accepting processes (Burke et al.,
2005). The reduction and precipitation of Tc(VII) were associated with the formation
of biogenic Fe(II) and were catalysed by pure cultures of Fe(III)-reducing prokaryotes
inoculated in sterilized microcosms. Technetium solubility has also been studied by
using core samples from a shallow, sandy aquifer located on the US Atlantic Coastal
Plain (Wildung et al., 2004). The dominant electron donor in the sediments was Fe(II)
(0-5 M HCl-extractable), with Tc(IV) hydrous oxide being the major solid-phase reduc-
tion product. The authors noted presumptive evidence for direct enzymic reduction
in only a few key sand samples. The potential for biogenic Fe (II) -mediated reduction of
Tc(VII) has also been assessed in detail in other studies, e.g. by using sediments from the
US Department of Energy's Hanford and Oak Ridge sites (Fredrickson et al., 2004).

The biochemical basis of Tc(VII) reduction has been best studied in Escherichia coli.
Initial studies demonstrated that anaerobic, but not aerobic, cultures of Escherichia
coli reduced Tc(VII), with the reduced radionuclide precipitated within the cell (Lloyd
et al., 1997). Results obtained from studies conducted with wild-type cells and 34
defined mutants defective in the synthesis of regulatory or electron-transfer proteins
were used to construct a model for Tc(VII) reduction by Escherichia coli. The central
tenet of this model is that the hydrogenase 3 component of formate hydrogenlyase
catalyses the transfer of electrons from dihydrogen to Tc(VII) (Fig. 4). According to this
model, the formate dehydrogenase component (FdhH) is required only if formate, or a
precursor, is supplied as an electron donor for Tc(VII) reduction in place of hydrogen.
This model has been validated by the observations that a mutant unable to synthesize
hydrogenase 3 was unable to reduce Tc(VII) when either hydrogen or formate was
supplied as an electron donor (Lloyd et al., 1997). Hydrogenase-mediated Tc(VII)
reduction has also been noted in sulfate-reducing bacteria (Lloyd et al., 1999b; De Luca
et al., 2001) and Geobacter sulfur re ducens (J. C. Renshaw 6c J. R. Lloyd, unpublished
observations).

DEGRADATION OFXENOBIOTICS BY METAL-REDUCING
BACTERIA

Some metal-reducing bacteria, most notably Geobacter species, have the ability to
couple Fe(III) reduction to the complete oxidation of aromatic contaminants (Lovley et
al., 1989b; Lovley 6c Lonergan, 1990; Coates et al., 2001). It is possible to stimulate

SGM symposium 65

292 J. R. Lloyd

(a)

Formate

FdhH

FNR

Hydrogenase 3

Tc(VII)

Tc(IV)

25-

(c)

"D 20

c
o

Tc cl

"> 15-

£ 10

Cu P

5-
0..

i — <—

u I

— i — i — i — i — «

i — . — , — . — , — , — .-

Energy (keV)

Fig. 4. Mechanism of Tc(VII) reduction by E. coli. Hydrogenase 3 of the formate hydrogenlyase
complex is able to catalyse the reduction of soluble Tc(VII) to insoluble Tc(IV) (a). Hydrogen or formate
are suitable electron donors; with the latter a formate dehydrogenase (FdhH) is also required for
reduction of Tc(VII). The insoluble Tc(IV) is precipitated within the cell, visible as an electron dense
deposit in TEM images of thin sections of the cells (b) and confirmed by EDS analysis (c). Bar, 1 jim.

these activities by increasing the bioavailability of Fe(III) oxides in subsurface sediments
by, for example, using Fe(III) chelators that solublize Fe(III) (Lovley et al., 1994) or
humic acids that can act as electron shuttles between the Fe(III) -reducing species and
Fe(III) oxides (Lovley et al., 1996). Both approaches eliminate the need for the organism
to contact the insoluble Fe(III) oxide directly to reduce it.

SGM symposium 65

Microbial metal reduction 293

(a)

Xenobiotic hydrazone and azo bonds are part of the chromophore

H0 3 SOCH 2 CH 2 -S— t ^-N— N y J" 2 ^=N— I Vs-CH.CH.OSO^

H0 3 S

SO3H

Xenobiotic aromatic sulfonic acid groups make the dye highly soluble

(b)

Fig. 5. Reduction and decoloration of the azo dye remazol black B (a) by anaerobically grown cells of
Shewanella sp. J 1 81 43 (b). Figures were generously provided by C. Pearce (University of Manchester,
UK).

In addition to coupling the oxidation of aromatics to metal reduction, Fe (III) -reducing
bacteria can also reduce redox-active organic xenobiotics in lieu of Fe(III). A good
example is the reduction of recalcitrant, coloured azo dyes, the disposal of which poses
a considerable problem to textile-dyeing industries worldwide. Due to the relatively low
levels of dye-fibre fixation in current reactive dyeing processes, as much as 50 % of the
dye that is present in the original dyebath is lost to the wastewater. Physical and/or
chemical processes are available, but are costly and can generate problematic
concentrated sludges for disposal. An alternative procedure utilizes anaerobically
grown cultures of a Shewanella species (designated strain J18143), isolated from soil
contaminated with textile dyes, to reduce and decolorize textile wastewaters (Nelson et
al., 2000). This highly efficient biocatalyst has been incorporated into the recently
developed BIOCOL commercial process for the treatment of azo dyes (Conlon &
Khraisheh, 2002). In this process, the bacterial cells are immobilized on an activated

SGM symposium 65

294 J. R. Lloyd

carbon support that adsorbs the target dye molecules and the potentially toxic amine
breakdown products for further biodegradation. Recent studies from our laboratory
have studied the underlying physiology of this organism, which is able to reduce a
wide range of azo compounds (e.g. remazol black B; see Fig. 5) and is compatible with
realistic process conditions, including moderate temperatures and alkaline pH
(C. Pearce, J. Guthrie & J. R. Lloyd, unpublished data).

CONCLUSIONS

Although the environmental relevance of microbial metal-reduction processes has only
recently become apparent, rapid advances in the understanding of these important
biotransformations have been made. However, we still have much to learn about the
precise mechanisms involved and the full impact of such reactions on a range of
biogeochemical cycles. Given the availability of genomic sequences for key metal-
reducing micro-organisms, new post-genomic approaches and the possibility of
combining these tools with advanced techniques from other branches of science and
technology (e.g. isotopic, spectroscopic and computational tools), rapid advances in
these areas are predicted.

ACKNOWLEDGEMENTS

The author thanks the UK Natural Environment Research Council, Biotechnology and
Biological Sciences Research Council and Engineering and Physical Sciences Research Council
and the Natural and Accelerated Bioremediation Research (NABIR) programme of the US
Department of Energy for financial support.

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SGM symposium 65

New insights into the physiology
and regulation of the anaerobic
oxidation of methane

Martin Kruger 1 - 2 and Tina Treude 2

1 Federal Institute for Geosciences and Resources (BGR), Stilleweg 2, D-30655 Hannover, Germany
2 Max-Planck-lnstitute for Marine Microbiology, Celsiusstrasse 1, D-28359 Bremen, Germany

INTRODUCTION

Methane is an important link within the global carbon cycle and has become a major
focus for scientific investigations over the last decades, especially since the discovery of
large deposits of methane hydrates in continental margins. The majority of recent
methane production is biogenic, i.e. produced either by thermogenic transformation of
organic material or by methanogenesis as the final step in fermentation of organic
matter carried out by methanogenic archaea in anoxic habitats (Reeburgh, 1996). There
are also abiotic sources of methane, e.g. at mid-oceanic ridges, where serpentinization
takes place. In marine environments, the bulk of the methane is produced in shelf and
upper continental-margin sediments, which receive large amounts of organic matter
from deposition (Reeburgh, 1996). As methane builds up, it migrates upwards and may
reach the sediment surface. Here, its ebullition and oxidation can lead to the formation
of complex geostructures, such as pockmarks or carbonate chimneys and platforms, as
well as large-scale topographies, such as mud volcanoes and carbonate mounds (Ivanov
et al., 1991; Milkov, 2000). In most of the deeper continental margin and the abyssal
plain sediments, methane production is low, as only 1-5 % of the surface primary
production reaches the bathyal and abyssal seabed, due to degradation processes in the
water column (Gage & Tyler, 1996).

Despite the high rates of methane production in shallow marine regions, the contri-
bution of the oceans, with around 3-5 % to the global methane emission into the
atmosphere, is extremely low compared with major methane sources, such as wetlands,
rice fields or ruminants (IPCC, 1994; Reeburgh, 1996). The reason for this is the

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

304 M. Kruger and T. Treude

SOJ (tnmol m 3 )

o 10 20 ao

Depth
(cm)

^sol

i
100

\CH 4

150

i i

Sulfate-dependent
AOM

2 4

CH 4 (mmoidm 3 )

Fig. 1. Scheme depicting typical profiles of sulfate and methane concentrations in porewater of an
anoxic marine sediment, indicating a distinct zone of AOM activity.

presence of sulfate in marine systems, which is an electron acceptor used for microbial
sulfate reduction (Jorgensen & Fenchel, 1974; Jorgensen, 1982). As long as sulfate is
present in the sediment, methanogenesis is shifted to deeper sediment layers due to
substrate competition between sulfate-reducing bacteria and methanogens (Zehnder,
1988). When migrating towards the sediment surface, methane is consumed by two
microbial pathways: anaerobic oxidation of methane (AOM) and aerobic oxidation
of methane. Reeburgh (1996) proposed that, by these processes, up to 80 % of the
methane produced in the marine sediments is oxidized prior to its release into
the hydrosphere.

Unlike freshwater and terrestrial habitats, aerobic oxidation of methane is less import-
ant in marine ecosystems, as oxygen availability in the sediments is low compared to
sulfate, the electron acceptor for AOM (D'Hondt et al., 2002). However, information
about rates and micro-organisms involved in aerobic oxidation of methane in the ocean
is scarce. So far, only a few studies have been published (e.g. Lidstrom, 1988; Holmes et
al., 1996; Valentine et al., 2001; Kruger et aL, 2005), of which the majority deal mainly
with pelagic processes.

In marine sediments, the bulk of upward-migrating methane is consumed during AOM
using sulfate instead of oxygen as electron acceptor (equation 1; Zehnder & Brock,
1980; Hoehler et al., 1994; Reeburgh, 1996; Valentine & Reeburgh, 2000; Hinrichs &

SGM symposium 65

Physiology and regulation of AOM 305

Table 1. Characteristics of habitats investigated for AOM (see also Fig. 1).

Characteristic

Hydrate Ridge

Eckernforde

Bay

Chilean
continental margin

Type of seep

Gas seep/gas
hydrates

Gassy coastal

sediment

Diffusive system

Methane transport

Advective

Diffusive/advective

Diffusive

Sulfate penetration depth (cm)

150-350

Sediment depth at which
methane is completely
consumed (cm)

Methane reaches
hydrosphere

0-5

110-360

Thickness of AOM zone (cm)

0-10

20-25

4-40

Measured areal AOM rate
(mmol m -2 day -1 )

56-100

0-4-1-5*

0-2-5-7

Calculated methane flux
(mmol rrr 2 day -1 )

56-200t

0-6-1-3*

0-07-0-13

Retention of methane in
sediment (%)

50-100

1 00 (except gas)
bubbles

100

* Rates determined by Treude etal. (2005b).

tCalculated from rate measurements (Treude etal., 2003) and methane effluxes (Torres etal., 2002).

t- Estimated from the methane profiles of Abegg & Anderson (1 997).

Boetius, 2002). The high availability of sulfate enables methane consumption far below
the sediment-water interface (Fig. 1; Table 1).

CH 4 +S0 4

HCOI+H.O + HS"

(equation 1)

It has been proposed that the only major escape route for methane into the hydrosphere
or atmosphere is ebullition of free gas or gas hydrates floating from the seabed.
Wherever sulfate is available, AOM communities control the release of dissolved
methane from the sediment (Treude, 2003). Without this mechanism, the contribution
of the world's oceans to global methane emission would approximately equal the
amount of methane emanating from ruminants, one of the biggest sources of today's
methane emission into the atmosphere (Reeburgh, 1996).

After a historical introduction, this review focuses on recent findings concerning
the micro-organisms involved in AOM and its environmental regulation and new
information concerning the mechanism of this still enigmatic process.

HISTORY OF AOM

The first published investigation of AOM was by Martens & Berner (1974), who
described conspicuous methane and sulfate profiles in organic-rich sediments (Fig. 1),

SGM symposium 65

306 M. Kruger and T. Treude

in which methane did not accumulate before sulfate depletion from porewater. From
the decrease of methane concentrations in the sulfate-reducing zone, they concluded
that methane must be consumed with sulfate instead of oxygen as terminal electron
acceptor.

Since then, more biogeochemical evidence has been published confirming that the
process of methane consumption in marine sediments takes place at the base of
the sulfate zone, linked directly or indirectly to the activity of sulfate-reducing bacteria
(Reeburgh, 1980; Iversen & Jorgensen, 1985; Hoehler et al., 1994; Hansen et al., 1998;
Niewohner et al., 1998; Borowski et al., 2000). These findings were based on depth
profiles of methane and sulfate/sulfide (Fig. 1), 13 C : 12 C ratios in carbon dioxide and
methane in sediment profiles, and labelling studies with sediment samples (e. g. Iversen
& Blackburn, 1981; Alperin & Reeburgh, 1985; Iversen &C Jorgensen, 1985; Hoehler
et al., 1994). However, Zehnder &C Brock (1979, 1980) were the first to demonstrate
methane oxidation under anoxic conditions by methanogenic archaea and hypo-
thesized a coupled two-step mechanism of AOM. They postulated that methane is first
activated by methanogenic archaea working in reverse, leading to the formation of
intermediates, e.g. acetate or methanol. In a second step, the intermediates are oxidized
to C0 2 under concurrent sulfate reduction by other non-methanogenic members of the
microbial community

Since these pioneering studies, knowledge of AOM has increased substantially,
involving biogeochemical, microbiological and molecular methods. Radiotracer
measurements enabled the first direct quantification of AOM and concurrent sulfate
reduction in anoxic marine sediments (Fig. 2; Reeburgh, 1976; Iversen & Blackburn,
1981; Devol, 1983). Iversen & Blackburn (1981) measured a 1 : 1 ratio of AOM and
sulfate reduction in the sulfate-methane transition zone of Danish sediments,
demonstrating the close coupling between these processes.

Hoehler et al. (1994) confirmed by thermodynamic modelling that a consortium of
methanogenic archaea and sulfate-reducing bacteria could gain energy from AOM
(Fig. 3). In situ and inhibitor studies (Hoehler et al., 1994), as well as laboratory
experiments with growing methanogens converting [ 14 C]methane to 14 C0 9 during
methanogenesis (Harder, 1997; Zehnder & Brock, 1979, 1980), further stimulated
the discussion about AOM being a reverse process of methanogenesis.

MICRO-ORGANISMS INVOLVED IN AOM

It has only been during the last 5-10 years that the identification of the micro-
organisms responsible for AOM has been possible. It was advanced by investigations of
lipid biomarkers in sediments from methane seeps. In the search for these organisms,

SGM symposium 65

Physiology and regulation of AOM 307

(TJ

E
u

10 000

1 000

100

Methane seeps

Gassy
sediment

Diffusive system

Hydrate Ridge

Black Sea
reef

Eckernf6rde

Bay

Chile 800 m Chile 1160 m

Fig. 2. Rates of AOM at methane hot spots in selected habitats differing in methane availability and
flux rates (mean ±sem, n = 3-1 8). Data are from Treude eta/. (2003) (Hydrate Ridge), Michaelisefa/.
(2002) (Black Sea), Treude era/. (2005b) (Eckernforde Bay) andTreudeef al. (2005a) (Chile).

CH 4

o »

o n - methanotroph'

Sulfate reducer

Gas hydrates/seeps

Fig. 3. Interactions between methanotrophic archaea and sulfate-reducing bacteria according to
the hypothesis of Hoehleref al. (1994). Scheme by K. Nauhaus.

scientists found methanogen-associated lipids, named crocetane, archaeol and hydroxy-
archaeol, in active methane seeps, revealing extremely light d u C values down to
-110 %o, giving evidence for an involvement of archaea in methane consumption (Elvert
et d. y 1999; Hinrichs et al, 1999; Pancost et al., 2000; Thiel et al., 2001; Schouten et al.,
2003). These archaeal lipids were also found in association with isotopically light
bacterial lipids, commonly found in sulfate-reducing bacteria (Hinrichs et al., 2000;
Hinrichs & Boetius, 2002; Elvert et al., 2003). Similar to archaeal lipids, this relative
enrichment in 12 C indicated the incorporation of methane-derived carbon into
bacterial cells.

SGM symposium 65

308 M. Kruger and T. Treude

Boetius et al. (2000) presented the first microscopic pictures of an AOM consortium
visualized by fluorescence in situ hybridization (FISH), showing aggregates of archaeal
cells surrounded by a shell of sulfate-reducing bacteria. The aggregates grow to a size
of about 6-10 ^m before they break apart into subaggregates, implicating the need to
keep short distances between cells during substrate exchange. These consortia were
discovered in surface sediment overlying methane hydrates at Hydrate Ridge, where
they represented >90 % of the microbial biomass.

After revealing the conspicuous morphology of the AOM consortium, further methods
were used to obtain direct evidence for methanotrophy of the AOM consortium. A
combination of FISH and secondary-ion mass spectrometry allowed the measurement
of (5 13 C profiles of the biomass of single aggregates (Orphan et al., 2001a). A high
depletion in 13 C, with values down to -96 and -62 %o, was detected in archaeal and
bacterial cells, respectively. These results confirmed the assimilation of isotopically
light methane by the consortium.

Molecular studies showed that the anaerobic methanotrophs (ANME) were affiliated
most closely with methanogenic archaea of the order Methanosarcinales (Hinrichs
et al., 1999; Orphan et al., 2001b) and have frequently been associated with sulfate-
reducing bacteria of the genera Desulfosarcina and Desulfococcus (Boetius et al., 2000;
Michaelis et al., 2002; Knittel et al., 2003). Today, three major groups of methane-
oxidizing archaea have been identified: ANME-1, ANME-2 and ANME-3 (Hinrichs
et al, 1999; Boetius et al., 2000; Orphan et al., 2001b; Knittel et al., 2005) . ANME-2 and
ANME-3 belong to the order Methanosarcinales. The ANME-2 group has recently
been divided into three subgroups, ANME-2a to -2c (Knittel et al., 2005), which seem to
exhibit differences in environmental preferences and the structure of their aggregation
with sulfate reducers. ANME-1 is distinct from, but related to, methanogenic archaea
of the orders Metbanomicrobiales and Methanosarcinales.

The bacterial diversity at methane seeps is high, especially within the d-proteo bacteria
(Knittel et al., 2003), including members of the genera Desulfosarcina, Desulfo-
rhopalus, Desulfocapsa and Desulfobulbus. Comprehensive overviews on the diversity
and phylogeny of archaea and sulfate reducers involved in AOM and associated with
methane seeps have been published recently by Knittel et al. (2003, 2005), respectively

HOT SPOTS FOR THE STUDY OF AOM

In general, AOM can be expected wherever methane and sulfate coexist in anoxic
environments. One main factor determining the magnitude of AOM is the methane
supply, because methane-turnover rates were found to increase with methane concen-
tration and methane flux (Nauhaus et al., 2002; Treude, 2003). Hot spots for AOM have

SGM symposium 65

Physiology and regulation of AOM 309

been found in diverse habitats, characterized by a wide range of environmental
characteristics. Hinrichs & Boetius (2002), Treude (2003) and Kriiger et al. (2005) have
reviewed AOM rates in marine sediments of different water depths, as well as methane
seeps. These first compilations of AOM field measurements and modelling have
suggested a direct coupling between methane supply and methane consumption in the
habitat. At methane seeps of ancient reservoirs or gas-hydrate locations, AOM rates
were found to be 10-100 times higher than those in non-seep regions. However, the
data of environmental AOM rates are still fragmentary. Below, two methane-rich
marine environments, which have so far played a major role during the investigation
of AOM, are briefly introduced: they are located at Hydrate Ridge, off the coast of
Oregon, USA (Boetius & Suess, 2004 and references therein) and on the north-western
shelf of the Black Sea (Michaelis et al., 2002).

Hydrate Ridge

At Hydrate Ridge, gas-hydrate deposits are located a few centimetres below the sedi-
ment surface in a water depth of 600-800 m, corresponding to the hydrate-stability
zone (Suess et al., 1999). These layers lead to very high methane fluxes (up to 200 mmol
m~ 2 day _1 ; Table 1), which fuel a diverse, seep-associated community (Sahling et al.,
2002; Treude et al., 2003), including zones covered by thick mats of sulfide-oxidizing
members of the genus Beggiatoa or inhabited by different dwelling clams, such as Cal-
yptogena and Acharax species. In these surface sediments, aggregates of methano-
trophic ANME-2 and sulfate-reducing DesulfococcuslDesulfosarcina cells dominate
the microbial biomass (Boetius et al., 2000). Their maximum abundance is located
within the upper 10 cm below the sea floor, where methane and sulfate meet (Treude
et al., 2003). In this zone, some of the highest densities of ANME cells and methane-
turnover rates known from marine environments have been found (Boetius et al., 2000;
Boetius & Suess, 2004). Elevated HCOT concentrations caused by this high AOM
activity result in an increase in alkalinity and support carbonate precipitation, forming
large carbonate landscapes at Hydrate Ridge.

Black Sea

In the north-western Black Sea, hundreds of active gas seeps occur along the shelf edge
west of the Crimea peninsula, at water depths between 35 and 800 m (Ivanov et al.,
1991). Within the anoxic zone, massive carbonate accumulations up to 4 m high and
1 m in diameter have been found associated with these seeps (Pimenov et al., 1997; Thiel
et al., 2001; Lein et al., 2002; Michaelis et al., 2002; Blumenberg et al., 2004). These
build-ups are covered by up to 10 cm thick microbial methanotrophic mats. From holes
in these structures, streams of gas bubbles emanate into the water column. Strong 13 C
depletions indicate an incorporation of methane carbon into carbonates, bulk micro-
bial biomass and specific lipids. The main matrix of the microbial mats consists of

SGM symposium 65

310 M. Kruger and T. Treude

densely aggregated cells of methanotrophic ANME-1 and sulfate-reducing Desulfo-
coccuslDesulfosarcina, but many other bacteria of unknown diversity and function co-
occur in the mats (Michaelis et aL, 2002; Blumenberg et aL, 2004; Knittel et aL, 2005).
The physiology of the methanotrophic mats has been studied in greater detail (Pimenov
et aL, 1997; Michaelis et aL, 2002; Treude, 2003; Nauhaus et aL, 2005), as described
below.

ENVIRONMENTAL REGULATION OF AOM

The most important questions regarding the functioning of AOM in the ocean concern
its regulation and the growth and environmental adaptation of the communities
mediating AOM. According to thermodynamic calculations, whether free energy is
available from AOM depends on the environmental settings (Zehnder 6c Brock, 1980;
Iversen & Blackburn, 1981). So far, only limited experimental data are available to
investigate the effect of variable environmental factors on the efficiency of AOM
(Valentine & Reeburgh, 2000; Nauhaus et aL, 2002, 2005). For example, the balance
between the concentrations of sulfate and sulfide might be an important factor
regulating microbial methane consumption (Treude, 2003; Treude et aL, 2003).

So far, no methanotrophic archaea are available for cultivation, probably because of
their extremely slow growth rate (Girguis et aL, 2003). Hence, the investigation of their
physiological capabilities and adaptations is only possible by in vitro studies with
naturally enriched samples from the environment. At many sites studied extensively for
AOM, including Hydrate Ridge, the Gulf of Mexico and the Black Sea, ANME-1 and
ANME-2 have been found to co-occur. However, the dominance of either group varies;
for example, in Hydrate Ridge samples, ANME-2 populations dominated the com-
munity (Boetius et aL, 2000) whereas, in the Black Sea mats, ANME-1 far outnumbered
ANME-2 (Knittel et aL, 2005). These differences in the community composition might
be due to differences in environmental parameters, such as temperature or the avail-
ability of methane and sulfate.

The stoichiometry of AOM, which has been estimated from porewater studies and
rate measurements in the field (Iversen & Blackburn, 1981), has been confirmed by in
vitro studies with environmental samples. Simultaneous measurements of methane
oxidation and sulfide production in samples from Hydrate Ridge and the Black Sea
have revealed a molar ratio of 1 : 1 between the two processes (Nauhaus et aL, 2002,
2005). Interestingly, the comparison of methane-driven sulfate reduction per cell
revealed that ANME-2 communities (Hydrate Ridge) were up to 20 times more active
than the ANME-1 communities in the microbial mats from the Black Sea. However, it is
not known whether all methanotrophic cells within a cell aggregate or within a mat are
equally active, as cells with no contact to the bacterial partner might be inactive if not

SGM symposium 65

Physiology and regulation of AOM 31 1

situated within an optimal substrate-concentration range required for sufficient energy
conservation (Sorensen et al., 2001). Therefore, it remains an interesting question for
future research, ideally with pure cultures, whether this difference in cell-specific
activity between ANME-1 and -2 can be attributed to specific substrate kinetics or
enzymic mechanisms of the ANME groups.

It can be assumed that the efficiency of AOM in mitigating methane emissions is
influenced by environmental parameters, such as pH, temperature and methane and
sulfate fluxes (Joye et al., 2004). Methane availability in situ depends on the methane
flux from subsurface reservoirs, as well as methane solubility, which is in turn influenced
by hydrostatic pressure and temperature (Yamamoto et al., 1976). Consequently, it
is essential to investigate the response of the AOM organisms to changes in these
parameters, to be able to estimate the effects of environmental or climatic changes
on AOM efficiency. An increase of methane partial pressure from 0-1 to 14 MPa
resulted in a fivefold increase of AOM rates in ANME-2-dominated samples (Nauhaus
et al., 2002) and a twofold increase in ANME-1-dominated samples (Nauhaus et al.,
2005). A similar stimulation was also detected in samples from shallow water depths
(Kriiger et al., 2005), which generally do not encounter such high concentrations of
methane. It is remarkable that it was also possible to induce AOM in formerly inactive
sediments by increasing the methane availability (Girguis et al., 2003; Kriiger et al.,
2005).

The free gas ebullition observed at seeps such as the microbial reefs in the Black Sea and
at Hydrate Ridge indicates methane saturation, with theoretical values of 2-3 MPa
(40 mM) and 8 MPa (140 mM) for the Black Sea and Hydrate Ridge, respectively.
Consequently, the rates observed in vitro at only 14 MPa must still represent sub-
stantial underestimations of rates occurring under in situ conditions. This inability in
reaching environmental methane concentrations might also inhibit attempts to culture
these organisms in the laboratory (see below).

The pH of the environment might also influence activities, as well as the composition
of microbial communities. For example, Nauhaus et al. (2005) showed that ANME-2
from Hydrate Ridge sediments had a distinct maximum of AOM rates at pH 7-4
(pH 7-7-5), whilst the pH optimum was broader for the ANME-1 community from the
Black Sea, ranging from pH 6-8 to 84. These small differences in environmental
preferences might contribute to the development of either ANME-1- or -2-dominated
communities.

Besides the pH, temperature is another important environmental factor influencing
micro-organisms. Despite a difference of only 4°C between in situ temperatures of

SGM symposium 65

312 M. Kruger and T. Treude

habitats investigated in the Black Sea and at Hydrate Ridge, the ANME-2 community
of Hydrate Ridge was more active at low temperatures (8-12 °C), whereas the ANME-1
community in the Black Sea was mesophilic, with highest AOM activities between 16
and 24 °C (Nauhaus et al., 2005). These distinct temperature optima for AOM of
environmental samples dominated by ANME-1 or ANME-2 may indicate a selective
advantage for either population. Further temperature optima for AOM ranged from
4°C for sediment from the Haakon Mosby Mud Volcano (-1-5 °C in situ) to 25 °C
in the Baltic Sea (between 4 and 16 °C in situ) (Kruger et al., 2005). So far, the different
temperature optima of AOM reflected the different in situ temperatures of the habitat.
Especially at shallow sites, seasonal temperature changes might also cause changes in
AOM activity. Indeed, such seasonality of AOM activity has been reported from studies
in Eckernforde Bay, Baltic Sea (Treude et al., 2005b) and Cape Lookout Bight, USA
(Hoehler *?**/., 1994).

So far, reports on the discovery of AOM have been restricted to habitats with sufficient
concentrations of sulfate available for the microbial partners of the methanotrophic
archaea (Hinrichs & Boetius, 2002 and references therein), predominantly in marine
habitats. However, the question is still pending whether, under specific environmental
conditions, AOM might also proceed with other alternative electron acceptors. In a
recent study with samples from Hydrate Ridge and the Black Sea, Nauhaus et al. (2005)
observed no AOM activity without sulfate. Instead, both ANME-1 and -2 communities
oxidized methane with similar rates at sulfate concentrations ranging from 10 to
100 mM. Other electron acceptors for AOM, such as nitrate, sulfur, ferric iron and
manganese oxide, were also reduced in Hydrate Ridge sediments by the indigenous
microbial population. However, this reduction was not coupled to AOM.

In summary, the ecological niches occupied more frequently by ANME-1 or ANME-2
seem to be defined mainly by temperature and the simultaneous availability of methane
and sulfate. Nevertheless, the factor(s) leading to the dominance of either group remain
to be identified.

MECHANISM OF AOM

The invesigation of mechanistic details of AOM is still hindered by the lack of pure
cultures of anaerobic methanothophs. Only a few studies are available, which have been
conducted with environmental samples naturally enriched in methanotrophic biomass
(Hoehler et al., 1994; Nauhaus et al., 2002, 2005). An important question regarding the
mechanism of AOM is whether the two reactions involved, i.e. methane oxidation and
sulfate reduction, are indeed carried out by a consortium of methanotrophic archaea
and associated bacteria (Fig. 3), as indicated by the striking structural features revealed
by microscopic analysis (Boetius et al., 2000; Orphan et al., 2001b; Michaelis et al.,

SGM symposium 65

Physiology and regulation of AOM 313

2002), or whether the entire process is mediated by a single organism. The latter is
indicated by repeated findings of ANME (mainly ANME-1) without contact to sulfate-
reducing bacteria (Orphan et al., 2001a; Michaelis et al., 2002; Joye et al., 2004; Treude
etai, 2005b).

Besides the microscopic evidence for a consortium of methanotrophic archaea and
sulfate-reducing bacteria mediating AOM, further evidence was gained by inhibition
experiments. In these experiments on environmental regulation of AOM, molybdate
and bromoethanesulfonate (BES) as inhibitors for sulfate reduction and methano-
genesis, respectively, have been used (Alperin 6c Reeburgh, 1985; Hoehler et al., 1994;
Hansen et al., 1998; Nauhaus et al., 2005). In both ANME-1- and -2-dominated sam-
ples, AOM was inhibited completely by BES. This inhibition was reversible and AOM
activity was resumed after removal of BES. Molybdate inhibited AOM completely in
ANME-2-dominated samples, but only partially in ANME-1-dominated samples
(Nauhaus et al., 2005), which was explained by strong adsorption of molybdate to
extracellular polysaccharide. These results are in good agreement with previous studies,
in which a partial to complete inhibition of AOM by these compounds was observed
(Alperin & Reeburgh, 1985; Hoehler et al., 1994). In conclusion, the application of a
specific inhibitor, i.e. BES for methanogens and molybdate for sulfate-reducing bacteria,
inhibited AOM. However, this inhibition is not absolute proof for syntrophy or a
symbiotic association between ANME and sulfate-reducing bacteria. A complete inhi-
bition would also be observed if methane oxidation and sulfate reduction were carried
out by a single organism.

Nevertheless, the inhibition of AOM by a methanogen-specific inhibitor provides
important evidence for the hypothesis that AOM is taking place via a reversal of
methanogenesis using similar enzymes (Zehnder 6c Brock, 1980; Hoehler et al., 1994).
Recently, an abundant protein closely resembling methyl coenzyme M reductase, the
terminal enzyme in methanogenesis (Thauer, 1998), was detected in ANME-1 cells
from Black Sea samples (Kriiger et al., 2003). This protein represents a likely candidate
for the initial step in AOM (Fig. 4). Furthermore, Hallam et al. (2003) found gene
sequences of the same enzyme in sediments with AOM activity, which they assigned to
ANME-2. In a subsequent paper, even the almost-complete methanogenic enzymic
system was attributed to ANME, based on the analysis of a metagenomic library from
an AOM site (Fig. 4) (Hallam et al., 2004).

If, indeed, a syntrophic mechanism is necessary, the question arises as to what the exact
mechanism of this process might be and which intermediates are exchanged within the
consortium (Fig. 3). The nature of the cooperation between the archaeal and bacterial
partners in AOM has not yet been elucidated (Sorensen et al., 2001; Nauhaus et al.,
2002).

SGM symposium 65

314 M. Kruger and T. Treude

CO*

formyl MF

formyl H 4 HPMT

methenyl-H 4 HMPT

26 ">2e-

methylene H 4 MPT

methyl H 4 MPT

methyl Co M

Formylmethanofuran dehydrogenase*
(fwd/fmd)

Formylmethanofuran H 4 MPT
W-formy transferase* (ftr)

Methenyl H 4 MPTcyclohydrolase* (mch)

Methylene H 4 MPT dehydrogenase*
(mtd)

Methylene H 4 MPT reductase (mer)
H 4 MPT-S-methy (transferase (mtr)

I coenzyme M reductase* (mcr)

CH (

Fig. 4. Activities (marked with an asterisk) and genes for methanogenic enzymes found in
methanotrophic archaea (AN ME). Activities are after Kruger eta/. (2003) and genes are after Hallam
eta/. (2004).

The addition of exogenous electron donors in the form of methanogenic substrates or
other C 1 -C 3 compounds, including acetate, formate, hydrogen and methanol, did not
stimulate sulfate reduction in the absence of methane (Nauhaus et al., 2002, 2005). In
theory, if the sulfate-reducing bacteria are adapted to the substrate supplied by the
methane-oxidizing partner, they should respond immediately to the addition of
potential AOM intermediates, with higher activity

As an alternative to hydrogen or carbon compounds, the possibility of electron transfer
between the archaeal and sulfate-reducing partners has been discussed for AOM
consortia (Sorensen et al., 2001; Nauhaus et al., 2002). Such an electron transfer to an
external acceptor has been reported previously for different anaerobic reactions
(Seeliger et al., 1998; Schink 6c Stams, 2001). However, the addition of several com-
pounds able to capture electrons - phenazines, AQDS (anthroquinone disulfonate) and
humic acids - did not replace the function of the sulfate-reducing bacterium. Neither of
the added compounds induced AOM (Nauhaus et al., 2005). One explanation for this
might be that, besides their important role in membrane-bound electron transport,
phenazines might also have toxic effects on micro-organisms (Ingram & Blackwood,
1970; Abken et al., 1998). Even though Straub & Schink (2003) showed that AQDS
may serve as an electron shuttle for iron-reducing bacteria, the redox conditions in
the incubations might change upon the addition of AQDS or humic acids, perhaps
suppressing AOM (Hernandez &C Newman, 2001). In summary, these experiments did

SGM symposium 65

Physiology and regulation of AOM 315

not provide direct evidence for a methanogenic or sulfidogenic substrate or electrons
as an intermediate in AOM. Consequently, despite the conspicuous aggregation of
archaea and sulfate-reducing bacteria in the environment, it might be possible that
AOM is carried out by one organism alone (see below).

GROWTH OF AOM MICRO-ORGANISMS IN THE LABORATORY

Pure cultures of anaerobic methanotrophic archaea have not yet been isolated.
Consequently, physiological studies have only been possible on environmental samples
naturally enriched in methanotrophic archaea and their sulfate-reducing partners
(Hoehler et al., 1994; Blumenberg et ai, 2004, 2005; Nauhaus et al., 2002, 2005).
However, sample availability and quality have limited the scope of these experiments.
For major questions regarding the mechanism and regulation of AOM, it seems -
despite substantial progress - inevitable to work with pure cultures. Nevertheless,
recent studies on microbial communities with a limited diversity have shown that there
is the possibility to work on these aspects of AOM by using metagenomic (Hallam et
al., 2004) or biochemical (Kriiger et ai, 2003) approaches.

Recently, Girguis et al. (2003) developed a novel continuous-flow system to study AOM,
which simulates the in situ conditions and could support the growth of anaerobic,
methanotrophic archaea. The major limitation of this system was that it only operated
under 0-1 MPa methane pressure and did not resemble in situ conditions on the sea
floor. Consequently, thermodynamic calculations showed that the Gibbs free-energy
yield for AOM at 1 atm methane pressure was low. In contrast, high in situ methane
concentrations and a close physical association between AOM partners can produce
energy yields sufficient to support biosynthesis (Sorensen et al., 2001; Nauhaus et ai,
2002). Therefore, it would be advantageous to use high-pressure systems, as described
by Nauhaus et al. (2002), for future research on the growth of AOM micro-organisms,
thus allowing the application of elevated methane partial pressures.

The means by which methanotrophic archaea are capable of growing at atmospheric
methane pressures are difficult to understand. High rates of AOM observed in the deep
oceans require high dissolved-methane concentrations (Nauhaus et al., 2002), which
can only be sustained at high pressures. Nevertheless, there is evidence for AOM at
atmospheric methane pressure in different shallow-water ecosystems (Iversen &
Jorgensen, 1985; Martens et al., 1999; Kriiger et al., 2005).

CONCLUSIONS

Microbially mediated AOM significantly influences biological and biogeochemical
processes on local to global scales. The process reduces methane flux into the water
column, stimulates subsurface microbial metabolism and also supports rich deep-sea

SGM symposium 65

316 M. Kruger and T. Treude

chemosynthetic communities that derive energy from one of its by-products, hydrogen
sulfide. However, advective systems such as methane seeps might still represent a
significant source of methane emission from the ocean to the hydrosphere, as recent
data have shown that AOM is only able to inhibit methane emission into the water
column completely in diffusive systems (Table 1; Treude, 2003).

The range of datasets available today (Orphan et al., 2001b; Hinrichs & Boetius, 2002;
Teske et al., 2002; Knittel et al., 2003; Treude et al., 2003; Boetius 6c Suess, 2004;
Kallmeyer & Boetius, 2004) suggests that the co-occurrence of methane and sulfate is
the major environmental factor defining the ecological niche occupied by AOM
communities. So far, no environment is known in which only one of the ANME groups
occurs. The significant dominance of either group points to the presence of defined
environmental niches within AOM zones, which have not been distinguished so far.
Further experimental studies of ANME-enriched samples from different environments
with a range of habitat characteristics are needed to answer the question of niche
selection.

Finally, it seems that, although a lot is known about AOM and the respective microbial
consortia, there are still many questions to be answered. For example, the intermediate
that is exchanged between archaea and sulfate-reducing bacteria is still unknown.
Thermodynamic considerations show that the archaea alone might not gain free energy
by this process, thus depending on cooperation with the sulfate-reducing bacteria.
Despite recent progress in the fields of biochemistry and metagenomics, this and other
questions will probably not be answered until pure cultures become available.

ACKNOWLEDGEMENTS

We especially thank Dr Gundula Eller for critical reading of the manuscript. This is publication
no. GEOTECH-129 of the programme GEOTECHNOLOGIEN of the Bundesministerium fiir
Bildung und Forschung and Deutsche Forschungsgemeinsch aft . Within this programme, the
study is part of projects MUMM-1, -2 and GHOSTDABS. Further support came from the Max
Blanch. Society.

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SGM symposium 65

Biogeochemical roles of fungi in
marine and estuarine habitats

Nicholas Clipson, 1 Eleanor Landy 2 and Marinus Otte 3

1 ' 3 Department of Industrial Microbiology 1 and Department of Botany 3 , University College Dublin,
Belfield, Dublin 4, Ireland

2 School of Biomedical and Molecular Sciences, University of Surrey, Guildford GU2 7XH, UK

INTRODUCTION

A fungal component of the marine biota was only recognized as recently as 1944
(Barghoorn & Linder, 1944), and it was not until the 1960s that studies commenced to
assess the extent and diversity of fungi in marine systems. Since this time, considerable
effort has been exerted to uncover marine fungal diversity, with high decadal discovery
indices in the 1970s and 80s (Hawksworth, 1991), resulting in around 1000 fungal
species known today from marine environments. Nevertheless, it is hardly surprising
that, with the extent of marine environments globally, we probably have a very in-
complete view of fungal diversity, together with their frequency and function in these
ecosystems. The objective of this review is to assess the extent of our present knowledge
and to highlight future directions to further elucidate their biology and ecology

THE NATURE OF MARINE ENVIRONMENTS

Marine ecosystems are globally extensive, and account for around 70 % of global
surface area. They can be defined generally as aquatic systems influenced by substantial
concentrations of salts, particularly sodium chloride, from existing oceanic systems.
Seas and oceans divide between regions bordering and influenced by terrestrial regions
and the open ocean, which is strongly zoned through the water column. These broad
boundaries are illustrated in Fig. 1, which also details linkages between marine com-
partments. At the sea surface, light is an important factor driving primary productivity
in the photic zone, to a depth of around 200 m (Dring, 1992), with carbon and other
nutrients moving downwards through the water column to stimulate food webs within
benthic zones, finally adding to sediments.

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

322 N. Clipson, E. Landy and M. Otte

V-i

.<*>

Terrestrial

Estuary <^

~^r>

Sediments

Atmosphere
/\

Coastal waters

=> <

Open ocean, pelagic

Deep ocean layer, -_

sub-pelagic * J

Near ocean floor,
benthic

Hydrothermal
vents

Sediments

Fig. 1. Interactions between and within marine ecosystems. While most interactions are bi-directional,
interactions between rivers and estuaries and between hydrothermal vents and their surroundings are
unidirectional. Circular arrows within oceanic compartments indicate cycles that are largely confined
within those compartments (after Wangersky, 1 980; Salomons & Forstner, 1 984).

Salinity concentrations within open ocean systems are characteristically very stable, at
around 500 mM NaCl, although these are lower in enclosed seas such as the Baltic and
Black Seas, where NaCl concentrations are diluted by the effects of freshwater inputs
from riverine discharge and ice melt and may be in the region of 50-200 mM. Typical
open ocean sea water ion concentrations are given in Table 1. The open ocean water
column changes with depth, with productivity declining dramatically as light and
temperature decrease. The ocean floor is an area of low primary productivity and low
biomass density (Nybakken, 1997) except for deep-sea hydrothermal vent and cold seep
communities, which are characterized by temperatures in the range 8-16 °C (as
opposed to the normal 2 °C of most of the deep sea) and high pressures exerted by the
overlying water column (Ballard, 1977). These environments are generally found at
the edges of tectonic plates and are typically very rich in reduced sulfur compounds.
Primary productivity is driven by prokaryal chemolithoautotrophs, which drive

SGM symposium 65

Roles of fungi in marine and estuarine habitats 323
Table 1. Concentrations of major and selected minor constituents of sea water

Data were compiled from Martin (1 970) and Dring (1 992).

Constituent

Concentration

Constituent

Concentration

(g kg 1 )

(mM)

(Mg I" 1 )

Chloride

19-35

548

Silicon

0-4900

Sodium

10-76

470

Nitrogen (combined)

0-560

Sulfate

2-71

Phosphorus

0-90

Magnesium

1-29

Aluminium

0-10

Calcium

0-413

Iron

0-1-62

Potassium

0-387

Zinc

1-48

Bicarbonate

0-142

Iodine

48-80

Bromide

0-067

0-8

Copper

0-5-27

Strontium

0-008

1-5

Manganese

0-2-8-6

Boron

00045

0-4

Cobalt

0-005-4-1

Fluoride

0-001

0-07

relatively rich food webs with high productivity (Nybakken, 1997). Cold seeps are often
associated with hypersaline brines or hydrocarbon seeps.

Near-shore marine ecosystems are much more variable and dynamic, particularly being
affected by terrestrial and meteorological influences. The most extensive and productive
of these environments are saline wetlands, which characteristically form as salt marshes
in polar and temperate zones dominated by low-growing herbs and mangrove
swamps in tropical regions dominated by salt-tolerant trees (Packham 6c Willis, 1997).
Both systems are affected by diurnal tidal inundation (much less so in enclosed seas,
where diurnal variation in tidal height is very small), which makes soil and sediment
environmental parameters highly variable, particularly for soil salinity, oxygenation
and drainage (Jefferies et aL, 1979; Armstrong et al., 1985). Terrestrial influences can
also be important, resulting in inputs of nutrients and freshwater as a result of riverine
deposition and run-off. Other salt-affected coastal environments might include dune
and shingle systems and rocky foreshores, which are influenced as spray communities
from sea water (Packham &C Willis, 1997). Adjacent to coastal regions, and where
continental shelves are shallow, coastal sea communities form, including coral reefs,
which are found in both tropical and cold seas. Marine communities can also be asso-
ciated with former oceanic activity, generally where seas have become separated from
present oceanic areas. These can either form highly saline inland seas such as the Dead
Sea, where saline concentrations can reach 4-5 M (Buchalo et al., 1998), or evaporate
completely to form salt pans and salt deserts. Many of these highly saline environments
are dominated by ions other than sodium and chloride, including magnesium, calcium

SGM symposium 65

324 N. Clipson, E. Landy and M. Otte

and potassium, with the counter ions carbonate, bicarbonate and sulfate (Flowers
et ai, 1986) . Perhaps 10 % of global land surface area is dominated by these hypersaline
environments.

For a detailed view of the elemental contents of the major environments associated
with marine systems, see Hutzinger (1980) and Salomons & Forstner (1984). In general
terms, open oceans are more dilute than areas above continental shelves, adjacent to
land masses or estuarine. Terrestrial and near-shore systems are usually well connected
via rivers, leading to rapid elemental transfer from terrestrial to marine ecosystems.
Transport to the open ocean and its underlying layers is much slower. In both cases,
many elements are immobilized in sediments, which act as elemental sinks after
deposition. These can be remobilized back slowly to the water column, particularly for
phosphorus and metals, where transfer to the atmosphere is negligible (Elmsley, 1980;
Salomons & Forstner, 1984). Some elements have a large atmospheric component, such
as carbon, nitrogen and sulfur, which are much more rapidly cycled in these systems.
This may also include some metals and metalloids such as mercury, arsenic and
selenium, which can be methylated and lost to the atmosphere.

Although open ocean systems are more stable than marine systems at terrestrial
boundaries, important dynamics do occur through the water column, largely reflecting
oceanic mixing by circulatory systems. Through the oceans there are local and depth-
delineated changes in factors such as temperature, salinity, oxygen and nutrient levels.
Examples are water-column transect studies across the Atlantic Ocean (Lavin et al.,
2003) and near-shore studies across the Strait of Georgia (British Columbia) (Masson,
2002). In the open ocean North Atlantic, Lavin et al. (2003) found a mean salinity of
around 3-5 %, varying between 3-45 and 3-7 % between 6000 m depth and the surface.
Mean surface oxygen was 200 ^mol kg -1 , declining to around 150 |imol kg -1 at 1000 m
but increasing to around 250 ^imol kg -1 at 6000 m. Silicates, nitrates and phosphates
were all depleted in the surface oceanic layers, increasing to average levels at around
1000 m and below. Oxygen and nutrient levels are also heavily influenced by the
circulation of currents through these systems, leading to localized variations.

MARINE FUNGAL DIVERSITY IN MARINE ECOSYSTEMS

The presence of fungi in marine ecosystems was first recognized by Barghoorn &
Linder (1944), who proposed the term 'marine fungi' for fungal species living in these
systems. Although there has been considerable effort since the early 1960s to elucidate
marine fungal biodiversity, present estimates probably grossly underrepresent its true
extent. This is for a number of reasons. Firstly, the oceans are exceptionally extensive,
both on the basis of their global surface area and as they are compartmentalized

SGM symposium 65

Roles of fungi in marine and estuarine habitats 325

through the water column. Many of these environments are inaccessible and a
challenge to meaningful sampling. Secondly, total fungal diversity is itself grossly
underestimated, with many geographical areas undersampled, many species lying
hidden in interactions with other organisms such as insects or plants and the problems
of isolating and culturing most species. Overall, Hawksworth (2001) estimates that
around 80000 species have been identified and classified in a background of possibly in
excess of 1-5 million species. It is within this background that marine fungal diversity is
probably heavily underestimated. Generally, two approaches are used to assess marine
fungal diversity. Firstly, direct observation of fungal structures on natural substrates or
baited substrates such as wood blocks or litter bags can be made. Secondly, isolations
can be made onto traditional agar-based cultural media. Certainly, in the latter case,
culturability of environmental isolates of fungi probably falls into the region of that for
bacteria, in the region of 1-5 %. There have been very few studies (e.g. Buchan et aL,
2003) to test whether conventionally made diversity estimates in a given location corres-
pond to molecular estimates. In the case of fungal diversity during the degradation of
the salt-marsh halophyte Spartina alterniflora, Buchan et al. (2003) found Phaeo-
sphaeria spartinicola, Phaeosphaeria halima and a Mycosphaerella sp. to be dominant
using terminal restriction fragment length polymorphism.

There are also problems in defining what a marine fungus is. Kohlmeyer 6c Kohlmeyer
(1979) advanced three ecological groupings of fungi occurring in marine environments,
including obligate marine fungi, facultative marine fungi and terrestrial fungi (see also
Kohlmeyer 6c Volkmann-Kohlmeyer, 2003). Obligate marine fungi were defined by
their inability to complete their life cycle outside the marine environment, whereas
facultative species could complete life cycles in either marine or terrestrial systems but
were of terrestrial origin. Terrestrial fungi do not have the ability to complete life cycles
in marine environments and, although commonly isolated, might just exist in a dor-
mant state until more favourable conditions prevail for spore germination (Kohlmeyer
6c Kohlmeyer, 1979). On a physiological basis, it is unclear what confers the marine
lifestyle of fungal organisms. The principal adaptation required by marine fungi is
probably salt-tolerance. Many terrestrial species, for example most aspergilli or
penicillia (e.g. Beever 6c Laracy, 1986), are exceptionally salt-tolerant both vegetatively
and for sporulation, but are not generally found to be active in marine ecosystems.
Some terrestrial Aspergillus species have been found in marine ecosystems, such as
Aspergillus sydowii and Aspergillus fumigatus, which cause disease in the Caribbean
sea fan Gorgonia ventalina. Disease-causing populations are probably maintained by
continual influx of vegetative material from neighbouring terrestrial environments
(Smith et al., 1996; Geiser et al., 1998). The question concerning the nature of the
marine fungal lifestyle remains. It might relate to the substrate conditions existing
in marine ecosystems, sensitivity at a certain lifestyle stage (for example sporulation or

SGM symposium 65

326 N. Clipson, E. Landy and M. Otte

spore germination), the oligotrophic nature of the marine environment, the selection of
appropriate media for physiological studies or the ability to attach to substrate.

Biogeographically, and in terms of marine ecosystem-type, fungi appear to be ubiqui-
tous. Nevertheless, some caution must be expressed with this view, as there are sub-
stantial areas of the globe which have not been surveyed for this group. Marine fungi
appear to be found worldwide, including polar regions (Pugh & Jones, 1986; Grasso
et al.j 1997), although probably the highest levels of diversity so far have been found in
the tropics associated with coastal ecosystems, especially mangroves. This may reflect
that most effort has been exerted within these regions. Additionally, coastal and near-
shore habitats are much better studied than oceanic habitats because of the expense of
boat time. Benthic and deep-sea fungal communities have only been sampled on a very
few occasions (e.g. Alongi, 1987; Nagahama et al., 2001). Another important issue
leading to diversity underestimates is the issue of culturability. Traditionally, marine
fungal diversity has been assessed almost exclusively by culture-based isolation
approaches or direct observation of natural substrates.

The most comprehensive assessment of known marine fungal diversity to date was
made by Jones &C Mitchell (1996), taking into account those that had been described
and estimating species waiting for full description and publication, including thrausto-
chytrids and lower fungi. They advanced a total of around 1400 species. Another
approach to assessing marine fungal diversity has been to compile checklists of species
occurring either on particular substrates or in known areas. Schmit & Shearer (2003)
compiled a checklist from available literature of mangrove-associated fungi, recording a
total of 625 species, with both marine and terrestrial species recorded. Advances in
producing checklists of fungal species illustrate how knowledge of their diversity has
been extended. Kohlmeyer (1969) could list 76 mangrove-fungi species, Hyde 6c Jones
(1988) listed 90 species, with Jones & Alias (1997) extending this to 268 species. These
latter two studies speculated that mangrove-fungal diversity is most extensive in South-
East Asia, reflecting greater substrate diversity, although this probably also reflects
greater numbers of isolation studies.

Two studies have aimed to produce checklists on an area basis. Gonzalez et al. (2001)
have produced a checklist of higher marine fungi of Mexico, listing 62 species. They
believed that this number significantly underestimates the true diversity of Mexican
marine fungi. In a literature search of marine fungi identified within the seas and
coastal regions of the European Union (EU), Clipson et al. (2001) found 318 species
described. Interestingly, all three checklists, together with that of Jones & Mitchell
(1996), found that marine fungi were predominantly either ascomycetes (particularly
the Halosphaeriales) or mitosporic fungi, with details given in Table 2. Many of the

SGM symposium 65

Roles of fungi in marine and estuarine habitats 327

Table 2. Marine fungal diversity, based upon a global estimate and literature
searches for the European Union area, Mexico and on mangrove substrates

Data were taken from Jones & Mitchell (1 996) (global), Clipson etal. (2001 ) (EU marine), Gonzalez ef
a/. (2001 ) (Mexico marine) and Schmit & Shearer (2003) (mangrove). Species awaiting description and
publication are indicated in parentheses.

Fungal group

Global

EU marine

Mexico marine

Mangrove

Lower fungi

100 (+32)

Thraustochytrids

Ascomycetes

214

278

Hemiascomycetes

Euascomycetes

305 (+143)

Deuteromycotina

79 (+200)

277

Basidiomycotina

7 (+7)

Trichomycetes

Lichens

18 (+410)

genera found have common characteristics (reproduction in aquatic habitats; thin-
walled, unitunicate, deliquescing asci; central pseudoparenchyma in immature asci;
hyaline, bicelled ascospores with polar or equatorial appendages) which Jones (1995)
considers as adaptations to aquatic habitats. The large number of mitosporic fungi
found is problematic because little is known about their relationship with their sexual
stages. Currently, the EU list is being updated, with 338 marine species being recognized
by 2005 (E. Landy, unpublished).

Schmit &C Shearer (2003) also explored substrate preferences of mangrove fungi, with
marine ascomycetes being intertidal or submerged species and mitosporic species being
associated with sediments and possibly of terrestrial origin. Most marine ascomycetes
possess appendages or gelatinous sheaths to aid attachment, making them particularly
adapted to the forces operating in intertidal zones. Also, there is a further problem in
that marine fungi are defined only imprecisely, with single mangrove trees possessing
populations of both marine and terrestrial mangrove fungi. At what position on a
single tree, and exposed to what outside environmental conditions, is a fungal species
marine? At best, this must represent a rather arbitrary division. As Schmit & Shearer
(2003) point out, we need to know much more about the evolution and ecology (related
to their physiology and genetics) to understand the determinants of distribution of
mangrove fungi. The same is true for marine fungal distribution in all other marine
habitats. To date, only coastal habitats have been explored in any detail, leaving fungal
diversity in the vast bulk of the open oceans and their depths largely undetermined. It is
known that yeasts are frequent in open seas (e.g. Van Uden & Fell, 1968). In a study
based in the Atlantic south of Portugal, Gadanho et al. (2003) identified 31 yeast taxa

SGM symposium 65

328 N. Clipson, E. Landy and M. Otte

using microsatellite-primed PCR. Yeast cell densities typically decline with increased
depth and distance from land (Gadanho et al., 2003), presumably reflecting substrate
availability and favourable environmental conditions. Yeasts are also known in the deep
sea, with Nagahama et al. (2002) reporting the presence of red yeasts from abyssal
zones. There are also some reports of fungi from hypersaline habitats. From the Dead
Sea, Kritzman (1973) reported an osmophilic yeast, and Buchalo et al. (1998) made the
first reports of filamentous fungi, including species of Gytnnascella, Ulocladium and
Penicillium, for this hypersaline habitat. A number of species of melanized yeast-like
fungi have been reported from hypersaline salterns (Gunde-Cimermann et al., 2000).

FUNGAL ADAPTATION TO MARINE ENVIRONMENTS

That a specialized mycoflora has become representative of marine ecosystems would
suggest that specialized adaptations are necessary for the marine fungal habit. Clearly,
adaptation to a combination of factors such as salinity, redox potential, pressure,
substrate availability and quality and temperature, throughout the fungal life cycle,
determines fungal diversity and activity in individual marine ecosystems.

The best-studied fungal response to such factors is that to salinity, which is ubiquitous
in marine ecosystems. Much of our understanding of the mechanisms by which marine
fungi adapt to salt comes from the physiological studies on the marine hyphomycete
Dendryphiella salina by Jennings and co-workers in the 1980s and 1990s and from
subsequent genetic studies on this and other non-marine fungi (e.g. Clement et al.,
1999; and see Bohnert et al., 2001; Hooley et al., 2003). Integrating information
between these marine species and other osmotolerant but non-marine species has led to
a model of fungal cellular adaptation to salinity Central to osmotic adaptation by fungi
is maintenance of osmotic gradients across the hyphal membrane to maintain inwardly
directed water fluxes. Water (=osmotic) potential of sea water is around -24 MPa,
meaning that cellular water potentials have to be maintained more negative. Fungal
cells have rigid cell walls and possess turgor pressure generated as the difference
between cellular water potential and cellular osmotic potential (generated by osmo-
tically active solutes), turgor pressure being an important component of apical growth
and expansion processes in fungi (Money, 1997). To generate hyphal water potentials,
fungal cells have to maintain substantial concentrations of intracellular solutes.
The fungal cell has two potential sources of osmotically active solutes, ions taken up
from the external environment and organic solutes either synthesized from metabolism
or taken up externally if available. Both these strategies are potentially problematic
to fungal growth; many ions, particularly sodium and chloride, are toxic to cellular
metabolic processes, and organic solutes for osmotic purposes can divert carbon from
growth. In vitro studies have shown that many enzymes involved in fungal primary
metabolism are inhibited substantially by sodium chloride levels well below (approx.

SGM symposium 65

Roles of fungi in marine and estuarine habitats 329

100 mM) those of sea water (Pa ton & Jennings, 1988), although this may be altered in
vivo by either upregulation of protein systems or by effects of the counter-anion
environment (Nilsson & Adler, 1990; Hooley et al., 2003).

In Dendryphiella salina, a balance of these two approaches has been found to generate
hyphal osmotic potential (Clipson &C Jennings, 1992). X-ray microanalysis studies
of hyphal cells growing at sea-water concentrations (500 mM NaCl) indicated
cytoplasmic concentrations of around 170 mM sodium (Clipson et al., 1990). This is
probably still compatible with cytosolic enzyme activity, albeit with some inhibition of
more sensitive enzymes. There was also no evidence of preferential accumulation
of ions in vacuoles, which are much less involved in metabolism than is the cytoplasm
(Clipson et al., 1990). Many algal and plant halophiles have large vacuoles which are
used as ion stores for osmotic adjustment, in conjunction with low cytoplasmic ion
levels (Flowers et al., 1986). Although Dendryphiella salina does not have large vacu-
oles, vacuolar localization of ions may be important in those fungal species that do.

In Dendryphiella salina, the balance of osmotic potential is generated through
compatible solutes, either accumulated or synthesized. These solutes are 'compatible'
in that they do not interfere with cellular metabolism. In fungi, the principal compatible
solutes are polyols, including glycerol, mannitol, arabinitol and erythritol (Blomberg &
Adler, 1993). In most fungi, it is mannitol which is the primary osmoresponsive polyol,
with others generally increasing as cultures age. For species such as Dendryphiella
salina and the osmophilic yeasts Debaryomyces hansenii and Zygosaccharomyces
rouxii, where polyol concentrations have been measured under saline conditions
in vitro, polyols generated between 36 and 75 % of cellular osmotic potential (Hooley
et ai, 2003) . Other compatible solutes may also be important in marine fungi, including
proline and trehalose (Jennings & Burke, 1990; Davis et ai, 2000).

Clearly, these osmotic mechanisms are under genetic control, although precise details
are not understood at present due to the difficulty of genetic manipulation in marine
fungal species. The genetics of osmotolerance in yeast is quite well understood and has
been reviewed recently by Yale & Bohnert (2001) and Hooley et al. (2003). Saccharo-
myces cerevisiae is not a good model for halophilic fungi as it is relatively salt-sensitive,
with the halophilic yeast Debaryomyces hansenii being more representative. A number
of genes involved in stress responses by the marine yeast Debaryomyces hansenii have
been characterized, including the cloning of the genes for superoxide dismutase
(Hernandez-Saavedro & Romero-Geraldo, 2001), a homologue of the high-osmolarity
glycerol response gene HOG1 encoding a MAP kinase (Bansal & Mondal, 2000)
and the identification of several ion-transport-related genes including DhENAl and
DhENA2 for sodium ATPases (Almagro et al., 2001). Nevertheless, there is still a very

SGM symposium 65

330 N. Clipson, E. Landy and M. Otte

incomplete view of the genetics of salt tolerance in halophilic fungi in general and
filamentous marine fungi in particular.

So far, an overview of fungal salt tolerance has been presented. In most marine environ-
ments, particularly the open ocean, salt concentrations do not fluctuate appreciably or
rapidly, and adaptation is probably a continuous and gradual process presenting little
problem to fungal organisms. In coastal regions, the influence of tidal effects, evapo-
transpiration and freshwater inputs makes salt concentrations much more unstable
(Armstrong et al., 1985). Although the mechanisms are not known, presumably fungal
species living in these environments are adapted to adjust osmotically rather quickly in
response to such fluctuations. Such an ability might differentiate these species from
terrestrial fungi. Another important issue is tolerance to marine environmental factors
through the full fungal life cycle, particularly spore germination and early hyphal
development, and sporulation processes. These have not been examined to any extent,
although it is well known that many non-marine osmophilic fungi, particularly
aspergilli and penicillia, are tolerant through these stages and certainly in the labora-
tory can complete life cycles on highly saline media (Hooley et al., 2003).

Marine fungi, dependent upon the habitat they occupy, may have to be adapted to other
adverse environmental factors. On salt marshes, sediments and soils tend to be anoxic,
with highly negative redox potentials. Most fungal species are aerobic (Carlile et al.,
2001) and salt-marsh soils probably do not form an important habitat for fungi.
Although fungal species can be cultured from marine soils, these are probably transient
species only present as spores and not vegetationally active.

Open oceans tend to be relatively oligotrophic except where ocean circulatory systems
cause upwelling of nutrients, for example the Benguela System off south-west Africa.
This may lead to planktonic blooms in the photic zone, giving localized regions with
higher potential substrate concentrations for degrading organisms such as fungi and
bacteria. Within the water column, it might have been thought that low oxygen contents
might be inhibitory to fungal activity, although recent oceanological studies would
indicate that deep-sea water oxygen concentrations are probably close to saturated
(Lavin et al., 2003). Lorenz & Molitoris (1997) tested the effects of simulated deep-sea
conditions on a number of marine yeasts, finding that many grew under pressures
equivalent to a depth of 4000 m.

MARINE BIOGEOCHEMICAL PROCESSES

Our view of the role of fungi in biogeochemical processes remains rather limited,
perhaps with the exception of some coastal ecosystems. This mirrors diversity esti-
mates, particularly with regard to fungal distribution in the open ocean away from

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Roles of fungi in marine and estuarine habitats 331

coastal regions. In terms of substrate availability, there are radical differences between
the marine-terrestrial interface, perhaps one of the most productive ecosystems on
Earth (Newell, 1996), and the open ocean, which is relatively oligotrophic. Fungi are
heterotrophs and their abundance is likely to reflect availability of carbon-based
substrates. Knowledge concerning the fungal contribution to marine cycling processes
also mirrors research effort and ease of study Coastal environments are generally close
to research centres and require little specialist equipment to sample. In contrast, studies
performed at sea require expensive boat time, detailed planning and, particularly for
studies in the benthic and abyssal zones, sophisticated sampling equipment.
Nevertheless, this has been carried out widely and successfully for marine bacteria (e.g.
Eilersrt*/.,2000).

Biogeochemical cycling

Biogeochemical cycling refers to all processes and pathways involved in the cycling
and turnover of elements. This includes not only chemical transformations, but
also physical processes such as adsorption/desorption and biological transfer within
(compartmentation) and between organisms. The best studied and most relevant
elements important in marine biocycling are carbon, nitrogen, phosphorus and sulfur,
for which overviews are given in Fig. 2. For all four elements, the fluxes between and
within the compartments are very small compared with the sizes of the terrestrial
and oceanic pools. The fluxes of carbon and nitrogen between the oceans and the
atmosphere are bi-directional and of equal size, while the fluxes of phosphorus and
sulfur show a net deposition. The oceanic pools of nitrogen and phosphorus are of the
same order of magnitude as the terrestrial pools. In contrast, the oceanic pools for
carbon and sulfur are several orders of magnitude larger than the terrestrial pools.
While data available for carbon indicated that there is a distinct difference in the sizes of
the carbon pools between the surface layer and deep ocean, no such data are available
for the other elements. From a global perspective, all compartments are connected,
but this does not apply at the organismal (fungal) level (Hutzinger, 1980), with little
connectedness between open ocean near-surface ecosystems and the ocean floor.
Within the marine environment, a number of more or less distinct systems can be
identified that each have their own biogeochemical cycles (see Fig. 1). The focus of the
ensuing sections will be on biogeochemical cycling of individual elements, particularly
carbon, nitrogen, phosphorus and sulfur. Although these will be considered separately
for clarity, it must be noted that they are intrinsically linked within biogeochemical
cycling processes, all being basic constituents of organic matter.

Carbon, nitrogen and phosphorus. The quality and availability of organic
compounds based upon carbon skeletons is central to fungal distribution and function
within marine ecosystems. Fungi develop a range of responses to organic matter within

SGM symposium 65

332 N. Clipson, E. Landy and M. Otte

Terrestrial
2500

face 725
40

Open ocean
Deep 38 000

0-47 x10 6

0-16x10 6

1000

0-13 x10 6

0-27 x10 6

210

1-3x10 9

140

Fig. 2. Fluxes and cycling of carbon, nitrogen, phosphorus and sulfur through the oceans. Pools
(boxes) are expressed in 10 15 g for carbon and 10 12 g for nitrogen, phosphorus and sulfur; fluxes
(arrows) are in 10 15 gyear 1 for carbon and 10 12 g year 1 for nitrogen, phosphorus and sulfur. Circular
arrows within oceanic compartments indicate cycling through biota, but no such data were available
for sulfur (after Pierrou, 1976; Richey, 1983; Rosswall, 1983; Freneyefa/., 1983).

ecosystems, dependent upon the condition of organic matter, and are capable of
deriving nutrition saprotrophically from non-living organic matter or either biotrophi-
cally (symbiosis) or pathogenically from other living organisms. Certainly, we know
most about how fungi derive nutrition saprotrophically; nevertheless, there are many
examples of fish, invertebrate and plant pathogens in both open ocean and coastal
systems (see Porter, 1986, for overview) and examples of symbionts, particularly as
mycorrhizas and endophytes of marine plants (Cooke et al., 1993; Cornick et aL, 2005).
Availability of hosts will be important in determining the distribution of such fungal
interactions, although this has not been studied to any great extent.

SGM symposium 65

Roles of fungi in marine and estuarine habitats 333

Table 3. Production of exoenzymes for the breakdown of substrate polymers
demonstrated in isolates of filamentous marine fungi

Data were taken from Torzilli (1982), Schaumann ef a/. (1986), Rohrmann ef a/. (1992) and Bucheref
a/. (2004).

Marine fungal species

Exoenzyme produced

Lulworthia spp.

Halocyphina villosa, Lulworthia spp.

Halocyphina villosa

Halocyphina villosa, Lulworthia spp., Corollospora maritima,
Digitatispora marina, Cirrenalia pygmea, Varicosporina ramulosa

Lulworthia spp.

Pleospora pelagica, Pleospora vagans, Phaeosphaeria typharum

Halocyphina villosa, Pleospora pelagica, Pleospora vagans,
Phaeosphaeria typharum

Lulworthia spp.

Corollospora maritima, Cirrenalia pygmea, Varicosporina ramulosa,
Pleospora pelagica, Pleospora vagans, Phaeosphaeria typharum

Lulworthia spp.

Ascocratera manglicola, Astrosphaeriella striatispora,
Cryptovalsa halosarceicola, Linocarpon bipolaris, Rhizophila marina

Carrageenase
Laminarinase
Gelatinase
Caseinase

Alginase

Pectinase

Xylanase

Agarase
Lipase
Amylase
Chitinase
Cellulase complex

Laccase

Tyrosinase

Ligninase

Central to fungal nutrition is the production and extracellular secretion of degradatory
enzymes, making fungi particularly effective in the breakdown of complex polymeric
compounds. A number of studies has examined the range of exoenzymes produced
by marine fungal saprotrophs in vitro (Torzilli, 1982; Torzilli & Andrykovitch, 1986;
Molitoris & Schaumann, 1986; Schaumann et al., 1986; Lorenz & Molitoris, 1992;
Rohrmann et al., 1992), which are summarized in Table 3.

Although there is little quantitative information except for salt-marsh systems, fungal
biomass will reflect substrate availability This differs markedly between substrate-rich
coastal ecosystems and the generally oligotrophic open ocean. In the open ocean,
fungal communities are largely yeast-dominated, as opposed to comprising filamentous
species (Sieburth, 1979). In a study in the Pacific, Van Uden & Fell (1968) found viable
yeast cell numbers varying between 13 and 274 cells 1 _1 . In general, yeast cell densities
decline with increased depth and distance from land (Gadanho et al., 2003), pre-
sumably reflecting substrate availability and favourable environmental conditions.

SGM symposium 65

334 N. Clipson, E. Landy and M. Otte

Substrate differences do occur in the open ocean, particularly where the upwelling of
colder waters creates regions of high productivity Presumably, the activity of recycling
organisms such as fungi will be higher in these regions. Nevertheless, we know very
little about the role of yeasts in biogeochemical cycling in such systems.

Coastal zones are highly productive and rich in vegetation, which provides complex
polymeric material for recycling by fungi and other degraders. Fungi are highly effective
degraders of lignocellulosic material and may be central in commencing mixed degra-
dative processes for this material (see Newell, 1996). Many fungal isolates from marine
systems have been classified as white rot (lignin-degrading) organisms (Molitoris &
Schaumann, 1986; Bucher et al., 2004), particularly where isolates have been made from
woody material. Overall, it is clear that fungi of coastal zones are able to produce a very
extensive battery of exoenzymes, making them important components of recycling
activity

Nitrogen is essential to fungal growth and its availability is central to fungal activity in
all environments, including the marine environment (Sguros 6c Simms, 1963; Jennings,
1989; Feeney et ai, 1992; Edwards et ai, 1998; Ahammed & Prema, 2002). In general,
fungi can utilize both inorganic and organic forms of nitrogen, depending upon species
and environment (Carlile et ai, 2001). Little is known about how marine fungi derive
nitrogen from their environments or contribute to nitrogen cycling, although both
generalist and specialist utilizers will be present. In marine systems, total inorganic
nitrogen (as combined nitrogen) ranges between and 560 jig 1 _1 (Dring, 1992). This is
likely to limit fungal growth very much in the same way that marine primary produc-
tivity is limited by oceanic nitrogen and phosphorus concentrations (Dring, 1992).
Fungi probably derive the bulk of their nitrogen requirements from organic substrate
resulting from primary production, probably as reduced nitrogen, with close linkage
between factors affecting primary productivity and fungal biomass.

There are some examples of fungal species in marine habitats deriving nitrogen from
quite specific organic sources. For example, the trichomycete fungus Entromyces
callianassae occurs exclusively in the foregut lining of the callianassid shrimp
Nihonotrypaea harmandi and appears to be involved in the hydrolysis of certain
nitrogen-containing compounds (Kimura et al., 2002). Like their terrestrial counter-
parts, some marine fungi have been found to degrade nitrogen-rich materials such as
chitin which are not easily utilized by other organisms. Kirchner (1995) found that the
yeast-like fungus Aureobasidium pullulans was involved in the degradation of chitin in
the moults and carcasses of the marine copepod Tisbe holothuriae. In general, marine
fungi may play a role in the degradation of more recalcitrant nitrogen-containing
compounds, which may be only slowly degraded by other recyclers. Although quanti-

SGM symposium 65

Roles of fungi in marine and estuarine habitats 335

fying the contribution of marine fungi to nitrogen cycling is extremely difficult, Newell
(1996) considered that, on salt marshes, fungi could account for nearly all the nitrogen
present in decaying standing biomass of salt-marsh plants. Certainly, the role of fungi
in the marine nitrogen cycle may have been grossly underestimated, particularly at the
marine-terrestrial interface; our knowledge of the fungal contribution in the open
ocean is almost entirely absent.

Open ocean phosphorus concentrations are very low, ranging from to 90 ^ig l -1 (Dring,
1992). Similarly to nitrogen, these are likely to be too low to support appreciable fungal
growth, with most phosphorus derived from organic substrates related to primary
productivity However, very few reports exist on marine fungi in relation to phosphorus.
Bongiorni & Dini (2002) described the habitat, seasonality and species distribution of
thraustochytrids, marine fungoid protists, and found that, among the nutrients assessed
in the water column, only total phosphorus was related to variations in thraustochytrid
densities.

On salt marshes, fungi may be involved in phosphorus cycling as mycorrhizas, with
such associations reported in mangroves (Sengupta & Chaudhuri, 2002) and some
salt-marsh plant species (Carvalho et ai, 2001). Mycorrhizal colonization of salt-marsh
halophytes tends to be quite poorly distributed, with some of the most widespread salt-
marsh plant species such as the Chenopodiaceae and many of the Graminae not being
strongly mycorrhizal. In contrast, halophytic members of the Asteraceae do tend to be
functionally mycorrhizal (Rozema et ai, 1986; Carvalho et ai, 2001) . In the open ocean,
a much more oligotrophic environment than coastal waters, phosphorus availability
is known to limit bacterial productivity (Van Wambeke et ai, 2002), although effects
on yeast populations are not known.

Sulfur. While sulfur plays an important role in global biogeochemical cycling in
general, and in marine environments in particular, little is known about fungal
involvement in sulfur cycling. Sulfate is a major constituent of sea water, which typically
contains around 2-6 gl _1 (Dring, 1992). Also, sediments and hydro thermal vent regions
tend to be rich in sulfides (Dring, 1992). Fungi are known to be able to oxidize inorganic
sulfur (Wainwright, 1989) and to produce compounds such as dimethyl sulfide
(Slaughter, 1989), which is an important compound in the cycling of sulfur in marine
environments. Fungi are therefore likely to be involved in biogeochemical cycling of
sulfur, particularly in marine sediments.

A few studies report on fungal sulfur metabolism of inorganic and organic sources.
Phae 6c Shoda (1991) reported a fungal species able to degrade hydrogen sulfide,
methanethiol, dimethyl sulfide and dimethyl disulfide, while Faison et al. (1991)

SGM symposium 65

336 N. Clipson, E. Landy and M. Otte

reported on a coal-solubilizing Paecilomyces sp. able to degrade organic sulfur com-
pounds such as ethyl phenyl sulfide, diphenyl sulfide and dibenzyl sulfide. Reports
on the involvement of fungi associated with marine organisms in relation to sulfur are
scant. Zande (1999) described an ascomycete on the gills of the gastropod Bathynerita
naticoidea, which the author suggested could be involved in detoxification of the
abundant sulfide compounds in its habitat. Also associated with sulfide-rich marine
environments is the fungus Fusarium lateritium, which is able to degrade dimethyl-
sulfoniopropionate derived from algae and the salt-marsh grass Spartina alterniflora
(Bacic&Yoch,1998).

FUNGAL CONTRIBUTION TO CYCLING IN MARINE
ECOSYSTEMS

For most marine ecosystems, the contribution by fungi to biogeochemical cycling can at
best only be speculated upon. As considered above in the discussion of the effect of
carbon availability on marine fungal processes, fungal biomass appears to be maximal
(and presumably most active) where organic substrates are at high levels. This gives a
fundamental difference between the generally oligotrophic open ocean systems and the
land-ocean interface, where some of the most globally productive ecosystems exist. For
the latter, salt-marsh ecosystems are well studied and much is known about the role
of fungal activity in these systems (see Newell, 1993a, 1996).

Salt-marsh ecosystems have high rates of primary productivity resulting from nutrient
recycling, riverine external loading, tidally mediated net export of organic matter,
balancing of nutrient inputs and outputs as well as aerobic and anaerobic microbial
activity (Adam, 1990). The most intensively studied of these systems are those of the
eastern seaboard of the United States, where a sophisticated overview of the role of
fungal activity is now available (Gessner & Kohlmeyer, 1976; Gessner, 1977; Newell
et al., 2000), particularly for the salt-marsh cordgrass Spartina alterniflora Loisel.
Spartina alterniflora is an abundant and highly productive plant species in this system
and has frequently been used as the model for understanding fungal-plant interactions
in this unusual environment (Newell, 1993a, b).

Spartina alterniflora is highly lignocellulosic (>70 %) and is extensively colonized by
fungi during decomposition (Newell et al., 1996). Newell (1996) considered that there
are four main strategies by which marine microbes capture metabolizable organic
carbon: (i) maximized use of surface area coupled with high substrate affinity and
suitable enzymes that are capable of diffusing into solid particles; (ii) penetration by
tunnelling or surface erosion; (iii) penetration by absorptive ectoplasmic nets or
rhizoids; and (iv) pervasion via networks of self-extending tubular flowing cytoplasm
within rigid microfibrillar tubes of chitin laminaran or cellulose laminaran. Marine

SGM symposium 65

Roles of fungi in marine and estuarine habitats 337

fungi contribute to the degradation of Spartina alterniflora through the production of
lignocellulosic enzymes, which commence the breakdown of this material as plants
senesce leading to a subsequent mixed degradation. The life cycle of Spartina alterni-
flora involves shoot growth from perennial rhizomes from April to July, maturation and
seed production from July to October, frost-mediated death in October and an ensuing
decomposition period aided by weathering, shearing and tidal action. Fungi appear
to be central in commencing the decomposition process (Newell, 1993a), efficiently
breaking down lignin and cellulose into smaller, low-molecular-mass fragments
available either for conversion to fungal biomass or for further decomposition by
bacteria over the ensuing months. A number of distinct laccase sequence types have
been identified from a group of fungi isolated from this salt-marsh decay system (Lyons
etaL, 2003).

Some of the fungi reported to be involved in this process include Dreschslera halodes,
Halospbaeria hamata, Pbaeospbaeria typbarum, Stagonospora sp., Buergenerula
spartinae, Alternaria alternata, Epicoccum nigrum, Claviceps purpurea, Leptospbaeria
obiones, Leptospbaeria pelagica, Pleospora pelagica, Pleospora vegans and Lulwortbia
sp. (Gessner, 1977). More recently, ascomycetous fungi have been reported as the major
secondary producers within standing decaying leaves of S. alterniflora L. throughout its
geographical range (Newell et al., 2000; Newell, 2001). This has been further supported
by work on the fungal internal transcribed spacer sequences of these species by Buchan
et al. (2003). The four ascomycetes that appear to be most involved in this system are
Pbaeospbaeria spartinicola Leuchtmann, Mycospbaerella sp. 2 (of Kohlmeyer &
Kohlmeyer, 1979), Pbaeospbaeria halima Johnson and Buergenerula spartinae Kohlm.
et Gessner (Newell, 2001).

The fungal signal molecule ergosterol has also been used extensively to detail fungal
biomass in Spartina alterniflora-dom'mated salt-marsh systems. Significant differences
in ergosterol content of decaying leaves occurred, which were interpreted to indicate
differences in nitrogen (but not phosphorus) availability (Newell, 2001). Nitrogen is
thought to be a limiting nutrient for both Spartina alterniflora (Mendelssohn 6c
Morris, 2000) and its associated fungal populations (Newell et al., 1996). Looking for
seasonal differences, Newell's group also found that the ratio between fungal produc-
tivity and mean ergosterol contents was significantly higher in winter/spring than in
summer/autumn.

The situation in open oceans differs greatly from that in coastal regions. Organic
matter is rarely abundant, except where planktonic blooms occur, often resulting from
nutrient upwelling. Low substrate availability generally leads to very low fungal bio-
mass, of yeasts rather than the filamentous fungi characteristic of coastal ecosystems.

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338 N. Clipson, E. Landy and M. Otte

There are at present no quantitative data linking primary production to fungal numbers
in these systems. Beneath the photic zones in the open ocean, degraders presumably
become more important as organic material sediments through the water column. It is
known that bacterial numbers reduce with depth, with numbers at 5000 m being
around 10 % of those at 75 m (Dring, 1992), together with much lower bacterial growth
rates. It might be expected that, at some depth, particularly below the photic zone,
micro-organisms and fungi are the most important organisms (relative to total bio-
mass), growing on the detritus precipitating down from the ocean surface layers. At
great depths, a surprising abundance of life can be found near 'smokers' (deep-ocean
hydrothermal vents) and it is very likely that fungi form a significant component of
these systems (Duhig et al., 1992; Burgath & Von Stackelberg, 1995; Nagahama et aL,
2001).

CONCLUSIONS

In comparison with terrestrial environments, we have a rather restricted view of the role
of fungi in marine ecosystems. This partially reflects the extent of marine environments
on a global basis, together with their remoteness and difficulty for sampling. There is a
better understanding of land-sea interfaces, particularly from the studies of Newell
and co-workers in the eastern USA, although these studies need to be taken up for other
salt marsh/mangrove systems in other parts of the world. Although there has been
considerable effort to uncover fungal diversity, marine mycologists have been slow to
apply molecular methods to explore diversity. Whereas these would be routine for
marine bacteriologists, fungal molecular studies are few (e.g. Buchan et al., 2003). Our
knowledge of fungal diversity in open ocean systems is even more restricted, but could
more frequently be incorporated into programmes examining marine bacterial
populations, particularly as the predominant fungi in these systems are yeasts. We
urgently need to know more about fungi existing in benthic and deep-sea (hydro-
thermal) zones. Although we know something about marine fungal diversity in general,
little is known about their function except in salt marshes. We need to know to what
extent fungi contribute to food webs and fluxes of organic matter in marine systems
other than salt marshes, in order to give fuller views of these systems in general.
Exploration of new marine environments and determining the role of fungi in these
environments is the next challenge for marine mycology.

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SGM symposium 65

Role of micro-organisms in
karstification

Philip C. Bennett 1 and Annette Summers Engel 2

department of Geological Sciences, The University of Texas at Austin, Austin, TX78712, USA

department of Geology and Geophysics, Louisiana State University, Baton Rouge, LA 70803,
USA

INTRODUCTION

Whilst chemolithoautotrophic micro-organisms are found in nearly every environment
on Earth, they are more abundant in dark habitats where competition by photo-
synthetic organisms is eliminated. Caves, particularly, represent dark but accessible
subsurface habitats where the importance of microbial chemolithoautotrophy to
biogeochemical and geological processes can be examined directly At Lower Kane
Cave, WY, USA, hydrogen sulfide-rich springs provide a rich energy source for chemo-
lithoautotrophic micro-organisms, supporting a surprisingly complex consortium of
micro-organisms, dominated by sulfur-oxidizing bacteria. Several evolutionary lineages
within the class 'Epsilonproteobacteria' dominate the biovolume of subaqueous
microbial mats, and these microbes support the cave ecosystem through chemo-
lithoautotrophic carbon fixation. The anaerobic interior of the cave microbial mats is a
habitat for anaerobic metabolic guilds, dominated by sulfate-reducing and -fermenting
bacteria. Biological controls of speleogenesis had not been considered previously and it
was found that cycling of carbon and sulfur through the different microbial groups
directly affects sulfuric acid speleogenesis and accelerates limestone dissolution. This
new recognition of the contribution of microbial processes to geological processes
provides a better understanding of the causal factors for porosity development in
sulfidic groundwater systems.

Karst landscapes form where soluble carbonate rocks dissolve by chemical solution
(karstification), resulting in numerous geomorphic features, including caves and
subterranean-conduit drainage systems (e.g. White, 1988; Ford & Williams, 1989). This

SGM symposium 65: Micro-organisms and Earth systems - advances in geomicrobiology.
Editors G. M. Gadd, K. T. Semple & H. M. Lappin-Scott. Cambridge University Press. ISBN 521 86222 1 ©SGM 2005

346 P. C. Bennett and A. S. Engel

has traditionally been viewed as an abiotic, chemical process that occurs near the water
table, with biologically produced C0 2 as the principal reactive component.

This model fits well with the classic view of terrestrial subsurface microbial environ-
ments as a 'top-downward' model, where photosynthetically fixed, reduced carbon (C)
supplies the subsurface community with substrates for cell mass and energy (Kinkle &
Kane, 2000), because microbial processes occurring in the absence of light energy have
generally been considered insufficient to support ecosystem-level processes. In the past,
the subsurface community has been portrayed as a mixed community of heterotrophs,
with only hydrogen-oxidizing methanogens typically recognized as a significant auto-
trophic population (Stevens & McKinely, 1995; Chapelle et aL, 2002). However, the
absence of light energy does not preclude life, as subsurface environments couple
relatively constant temperatures and protection from surface conditions with an
abundance of dissolved inorganic solutes and exposed mineral surfaces that can serve
as sources of energy and nutrients (Maher & Stevenson, 1988; Stevens 6c McKinley,
1995; Ghiorse, 1997; Bachofen et aL, 1998; Rogers et aL, 1999; Bennett et aL, 2000;
Ben-Ari, 2002). Consequently, in subsurface environments, chemosynthetic micro-
organisms are vitally important to global chemical and ecosystem processes because
they gain cellular energy from the chemical oxidation of inorganic compounds and
fixing inorganic C, and serve as catalysts for reactions that would otherwise not occur
or would proceed slowly over geological time.

Sulfur-based microbial ecosystems include communities that oxidize energy-rich, non-
surface-derived, reduced S substrates [S(-II)], coupled to molecular-oxygen reduction,
to fix C0 2 . In near-surface environments, such as caves and mines in shallow bedrock
(<200 m depth), these communities can be examined directly and chemolithoauto-
trophic S oxidizers form isolated but diverse and highly productive microbial-mat
communities that support higher trophic levels within the subsurface ecosystems (e.g.
Sarbu et aL, 1996; Angert et aL, 1998; Hose et aL, 2000; Sarbu et aL, 2000; Engel et aL,
2001). The ubiquity of reduced S compounds in anaerobic groundwater suggests
that microbial communities that utilize reduced S are dispersed more widely in the
terrestrial subsurface than is currently recognized. Equally important, S(-II) oxidation
to sulfate generates excess protons that accelerate rock weathering, porosity develop-
ment and mobilization of toxic metals, demonstrating a geological significance of
terrestrial S-based microbial ecosystems beyond the descriptive characterization
of novel populations. Understanding the dynamics of S-based subsurface communities
and the coupling of the S and C cycles is a step toward linking microbial ecology to
geological phenomena.

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Role of micro-organisms in karstification 347

Karst aquifers

Caves act as portals to subsurface karst environments, with the typical habitat charac-
teristics being complete darkness, nearly constant air and water temperatures and
usually oligotrophic nutrient conditions, due to a limited supply of allochthonous
organic material (Kinkle & Kane, 2000; Poulson &c Lavoie, 2000). The classic model for
karst development (speleogenesis) involves carbonic acid dissolution, usually at and
rarely below the water table. More recently, sulfuric acid speleogenesis was proposed
from work in Lower Kane Cave, WY, USA (Egemeier, 1981). Based on observations of
H 2 S-bearing thermal springs, extensive gypsum mineral deposits and gypsum-replaced
carbonate rock walls, Egemeier (1981) put forth the sulfuric acid speleogenesis model,
which included volatilization of H 2 S from the sulfidic groundwater to the cave atmos-
phere and H 2 S autoxidation to sulfuric acid on the moist cave walls:

H 2 S+20 2 ^H 2 S0 4 (equation 1)

where the acid then reacted with and replaced the carbonate with gypsum:

H 2 S0 4 + CaC0 3 + H 2 -* CaS0 4 . 2H 2 + C0 2 (equation 2)

Gypsum is dissolved easily into groundwater and the net result is mass removal and an
increase in void volume. Cave formation due to sulfuric acid has now been recognized
in several active sulfidic cave systems (e.g. Hose et al., 2000; Sarbu et al., 2000) and has
been a process linked to the development of at least 10 % of carbonate caves worldwide
(Palmer, 1991).

Sulfur-based cave ecosystems

Microbial communities colonizing sulfidic cave habitats have received recent attention
due to their chemolithoautotrophic metabolism (e.g. Sarbu et al., 1996) and their
geological impact due to acid production (Vlasceanu et al., 2000; Engel et al., 2001;
Northup & Lavoie, 2001). Early studies of sulfuric acid speleogenesis (Egemeier, 1981)
assumed that abiotic H 2 S autoxidation was the important process for cave formation,
and did not consider microbial S(-II) oxidation. In some of the active systems, e.g. in
the Movile Cave (Romania), the Frasassi Caves (Italy) and Parker Cave (USA),
filamentous aerobic to microaerophilic S-oxidizing bacteria (SOB) dominate sub-
aqueous microbial mats (Angert et al., 1998; Hose et al., 2000) and were found to fix
inorganic C, providing energy to sustain complex cave ecosystems (Sarbu et al., 1996,
2000). Movile Cave provides us with an extreme example of a highly evolved, terrestrial,
chemolithoautotrophically based ecosystem (Sarbu et al., 1996), where 33 endemic,
cave-adapted, invertebrate taxa have been identified (from 48 total taxa), including
24 terrestrial and nine aquatic animal species.

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348 P. C. Bennett and A. S. Engel

Culture-independent and -dependent approaches have been used to characterize
the microbial communities, and 16S rRNA gene-based phylogenetic analyses of the
microbial mats from Movile Cave, Parker Cave, Cueva de Villa Luz (Mexico), the
Frasassi Caves and Cesspool Cave (USA) show diverse SOB, belonging to distinct
lineages within different classes of the proteobacteria and including the genera
Thiothrix, Thiobacillus, Thiomonas, Tbiomicrospira, Thiovulum and Achromatium
(Sarbu et al., 1996; Angert et al., 1998; Hose et al., 2000; Vlasceanu et al., 2000; Engel
et al., 2001). The predominant microbial groups in many of the investigated sulfidic
caves were found to belong to novel lineages within the class 'Epsilonproteobacteria'
(Angert et al., 1998; Engel et al., 2001, 2003) and, in addition to caves, closely related
organisms have been described from many S-rich, oligotrophic natural habitats,
including sulfidic springs, groundwater associated with oilfields, marine waters and
sediments, deep-sea hydrothermal-vent sites and vent-associated metazoans (Engel
et al., 2004b). Despite the unique S-based microbial diversity of these habitats, there
are few detailed studies describing the occurrence of epsilonproteobacteria from
terrestrial environments. Moreover, little is known about either the ecophysiology of
most epsilonproteobacteria, as many assemblages have not been cultured; from the
work that has been done, most of these bacteria cycle S as SOB (e.g. Takai et al., 2003).
The SOB communities, including epsilonproteobacteria, occupy the redox boundary
where reduced S mixes with aerobic water, and these microbes can take advantage of a
potential energy gradient to produce sulfate and protons as the ultimate end products
of their metabolism.

Colonization of the sulfidic caves by non-SOB microbial groups has rarely been
addressed. Heterotrophic and anaerobic micro-organisms related to the phylum
i Bacteroidetes i have been reported from some of the caves from 16S rRNA gene-clone
libraries (Angert et al., 1998; Engel et al., 2001), and dissimilatory sulfate-reducing
bacteria (SRB) and fermenting bacteria have been characterized from molecular
investigations of few sulfidic aquifers, including those associated with oilfields (Voor-
douw et al., 1996; Ulrich et al., 1998). Based on geochemical data, SRB and SOB have
been characterized from Yucatan cenotes open to phototrophic activity (Stoessell et al.,
1993). However, the importance of SRB for recycling S compounds by generating
supplemental H ? S that SOB can use (e.g. Widdel & Bak, 1992) has only recently been
addressed (Engel et al., 2004b).

Coupled C and S metabolism

Whilst the microbiology and biochemistry of an S-based system are complex, only a
few aspects relevant to S(-II) oxidation in a non-photosynthesizing environment are
reviewed here. SOB use reduced S compounds as electron donors (Madigan et al.,
1997):

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Role of micro-organisms in karstification 349

H 2 S+20 2 -*S05" +2H + (-798 kj mol" 1 ) (equation 3)

HS-+ V20 2 +H + - S°+H 2 (-209 kj mol" 1 ) (equation 4)

S°+H 2 + 1Vz0 2 -> SOj-+2H + {-5S7 kj mol" 1 ) (equation 5)

S 2 0|-+H 2 0+20 2 ^2SO^ +2H + (-823 kj mol- 1 ) (equation 6)

These general reactions proceed both biotically and abiotically, with abiotic autoxi-
dation of H 2 S proceeding spontaneously in aerobic aqueous systems (Millero et al.,
1987). Almost all non-photosynthesizing, chemolithoautotrophic SOB are obligate
aerobes, requiring 7 as the electron acceptor for S(-II) oxidation (Ehrlich, 1996). The
mechanisms of biological S oxidation are still being elucidated. For example, whilst
thiobacilli oxidize sulfide to sulfate directly (i.e. via equation 3), many SOB initially
form S° as an intermediate species (equation 4) that is then stored intracellularly and
further oxidized (equation 5) during periods of limiting sulfide (Ehrlich, 1996). Others
oxidize reduced-sulfur intermediates, such as thiosulfate (equation 6). Thiobacillus
denitrificans (Justin & Kelly, 1978), Thermotbrix thiopara (Caldwell et al., 1976) and
some epsilonproteobacteria (Moyer et al., 1995) couple S oxidation to nitrate reduction
in dysoxic environments, whereas Thiobacillus ferrooxidans couples the anaerobic
oxidation of S° to iron reduction (Corbett & Ingledew, 1987). Chemo-organotrophs in
dysoxic marine environments conserve energy from S° conversion to sulfuric acid and
H 2 S by electron transfer in the presence of a sulfide scavenger (e.g. Fe^ + ) without 1
(sulfur disproportionation) (Thamdrup et al., 1993):

4S°+4H 2 O^SOf+3H 2 S+2H + (equation 7)

Anaerobic microbes are common in marine and freshwater habitats, from anoxic sedi-
ments and bottom waters. Of the anaerobic microbial guilds, specifically SRB, many
have not been studied in detail from groundwater and karst springs. SRB, obligate
anaerobes that use molecular hydrogen to reduce sulfate to H ? S, are divided into two
broad physiological subgroups: those that oxidize acetate and those that do not.
Certain species of each subgroup are capable of chemolithoautotrophic growth with
CO? as the sole C source. If sulfate concentrations are high, SRB oxidize fermentation
by-products completely to C0 2 . In low-sulfate environments, however, SRB compete
with methanogens for hydrogen and organic compounds (Widdel Sc Bak, 1992;
Scholten^^/,,2002).

MICROBIAL GEOCHEMISTRY METHODS

An interdisciplinary approach was used to examine the diversity of cave microbial
communities and their roles in rock modification and karstification. By combining
detailed characterization of both the geochemical characteristics of the habitat (Engel
et al., 2004a) and the composition and structure of the microbial community (Engel et

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350 P. C. Bennett and A. S. Engel

al., 2003, 2004b), the link between microbial metabolism and geochemical interactions
can be examined. Recent research on the S-based ecosystem in Lower Kane Cave reveals
complex microbial communities that tightly cycle S and C compounds across redox
boundaries, whilst generating acidity and accelerating rock weathering (Engel et al.,
2004a).

Geochemical characterization

The objective of the geochemical characterization was to define the microbial habitat
and to quantify the changes in solute concentration or speciation that are indicators
of biogeochemical interactions. Water samples were collected for complete analyses of
dissolved constituents, isotopic ratios and geochemical parameters by using standard
methods and methods developed specifically for sulfidic cave systems (Engel et al.,
2004a). Field parameters included pH, temperature, specific conductance and field
alkalinity on a filtered sample. Dissolved oxygen (DO) is a critical habitat constraint
and was surveyed by using multiple methods, including electrochemical, microelec-
trode fluorescence quenching, colorimetric analyses and gas chromatography (GC).
Dissolved Fe 2+ is a sensitive indicator of the presence of dissimilatory iron-reducing
bacteria (Lovley & Phillips, 1988) and this was measured immediately in the field by the
ferrozine method to prevent sample degradation. Total and dissolved metal concen-
trations are used to characterize the geochemical consequences of the microbial activity
and were determined from a filtered acid-preserved sample by inductively coupled
plasma-mass spectrometry (ICP-MS). Dissolved anions, nutrients and organic acids
were measured by ion chromatography (IC) and total inorganic and dissolved organic
C (DOC) were measured by a carbon analyser. Stream water-dissolved solute speciation
and activity, equilibrium gas partial pressure and saturation state with respect to
mineral phases were calculated by using the geochemical speciation model phreeqc
(Parkhurst & Appelo, 1999). Sediments and core were characterized by X-ray diffrac-
tion (XRD) analysis of dried samples and total elemental composition by dissolution
and ICP-MS analysis. Sediment and core samples were also examined by conventional
and environmental scanning electron microscopy (CSEM and ESEM, respectively) for
the presence of micro-organisms on mineral surfaces, mineral composition by energy-
dispersive X-ray spectrometry (EDX) analysis and for evidence of mineral weathering.

Microbial-community analysis

Both culture-dependent and -independent methods were used to characterize the
metabolically active microbes in Lower Kane Cave. Typically, only 1 % of environ-
mental bacteria can be cultured (Leff et al., 1995) and standard culturing methods often
introduce a selective bias toward micro-organisms that are able to grow quickly and to
utilize the substrates provided in the medium (McDougald et al., 1998). However,
culturing allows for quantification of metabolically active organisms in the samples

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Role of micro-organisms in karstification 351

and can lead to species identification through phylogenetic analyses (Palleroni, 1997).
Laboratory strain isolation is essential to quantify the potential geological significance
of microbes from the different habitats and can be used to help target unknown
environmental samples. In contrast, molecular methods allow for characterization of a
microbial community that may have populations that are difficult, if not impossible, to
cultivate (Head etal., 1998).

Biomass determination. Biomass, both viable and non-viable, is a basic attribute for
characterizing mineral-microbe interactions and for understanding the implications of
the more detailed metabolic and phylogenetic analyses. Biomass of filtered samples was
determined by 4,6-diamidino-2-phenylindole (DAPI) cell counts (Gough & Stahl,
2003). This method is excellent for individual cells, but it is less reliable for filamentous
organisms where the definition of a single cell is problematic. For filamentous mat
samples, total organic C content of an acidified sample (to remove carbonate-mineral
material) was analysed by using a carbon analyser and the total cell-count estimate was
done by using the conversion factor of 350 fg C per cell (Bratbak & Dundas, 1984). This
method offers a maximum biomass, as it assumes that all of the organic C is associated
with biomass. Parallel determination of bacterial biomass can also be done by using
phosphatidyl ester-linked fatty acid (PLFA) analysis (Vestal & White, 1989). Phospho-
lipids are membrane lipids that are turned over rapidly during metabolism (White et al.,
1979). Consequently, phospholipids indicate viable microbial biomass at the moment
that the microbial mats were extracted and are not subject to the positive artefact of the
DAPI cell-count or C biomass-estimate methods.

Molecular methods. Microbial-community composition and structure were assessed
by using established methods and a full-cycle rRNA approach (Engel et al., 2003,
2004b). Total micro bial-mat environmental DNA was extracted and near-full-length
16S rRNA gene sequences (rDNA) were obtained by PCR amplification using either
archaeal and bacterial universal primers or lineage-specific primers (e.g. Ausubel et al.,
1990; Lane, 1991; Engel et al., 2003). A TA cloning kit (Invitrogen) was used to facilitate
PCR-product transformation and cloning procedures. Sequence inserts from clone
plasmids were PCR-amplified with plasmid-specific primers, purified by using Sepha-
dex columns and sequenced by using an automated ABI sequencer (e.g. ABI Prism
377X Perkin Elmer sequencer at Brigham Young University, Provo, UT, USA). 16S
rDNA clone libraries were constructed from samples throughout Lower Kane
Cave and, to facilitate community-composition analyses, clone sequence inserts were
screened by using restriction-fragment length-polymorphism (RFLP) analysis, where
selected clones from each representative RFLP pattern were sequenced (Engel et al.,
2004b). Results were analysed by using several phylogenetic-analysis methods after
closely related rRNA gene sequences obtained from the Ribosomal Database Project

SGM symposium 65

352 P. C. Bennett and A. S. Engel

were aligned by using clustal_x (Thompson et al., 1997) and then adjusted manually
based on conserved primary and secondary gene structure. Phylogenetic methods have
included neighbour joining by phylip (Felsenstein, 1993), maximum likelihood,
minimum evolution and maximum parsimony by paup* (Swofford, 2002) and Bayesian
inference in MrBayes (Ronquist & Huelsenbeck, 2003).

To identify and reliably quantify micro-organisms directly from the microbial mats, 16S
rRNA-targeted oligonucleotide probes were used for fluorescence in situ hybridization
(FISH) (Aim et al., 1996). In addition to general archaeal and bacterial probes, gene
probes were designed to target two novel groups of epsilonproteobacteria found in
Lower Kane Cave (Engel et al., 2003).

Metabolic survey. Microbial groups were enriched in pre-reduced anaerobic media
specific for different metabolic groups, including chemolithoautotrophic, lactate/
formate- or acetate-utilizing SRB and fermenting bacteria. Enumeration of each meta-
bolic group was done by using the most probable number (MPN) estimates of tenfold
serial dilutions of cultivable micro-organisms from each enrichment medium. All
enrichments were incubated in the dark at room temperature in a Coy anaerobic
chamber with N ? : H 2 mixed gases to maintain anaerobic conditions. Following enrich-
ment, growth was monitored by measuring OD 590 , evolved gases produced by the
mixed cultures (e.g. CH 4 for methanogens) were measured by GC and SRB were
screened for by measuring H 9 S production by GC.

In situ microcosms

The microbial influence on speleogenesis was evaluated by using the in situ microcosm
approach that has been used extensively to examine silicate colonization and weather-
ing (Hiebert & Bennett, 1992; Roberts et al., 2004). Sterile and non-sterile field
chambers (in situ microcosms) containing 0-5-1-0 cm 3 chips of Iceland Spar calcite
(Wards Scientific) and native limestone were deployed in the cave to test whether micro-
organisms or the bulk cave stream-water chemistry controlled carbonate dissolution.
Microcosms were constructed from 2-5 x 5 cm PVC pipes with screw caps on both ends.
Sterile microcosms had 0-1 |im PVDF hydrophilic filters on the end, whilst non-sterile
microcosms had 0-5 mm polyethylene mesh on either end to allow for fluid flow and
also for microbial colonization of the chips. Paired sterile and non-sterile microcosms
were placed throughout the cave stream and within the microbial mats, and remained in
the cave for 2 weeks to 9 months. A technique similar to the buried-slide technique was
also used, such that thin, polished wafers (surface areas of 1-2 cmj of limestone and
pieces of Iceland Spar were attached to glass slides and then surrounded by a 0-5 mm
mesh sack. At each microcosm site, a mesh sack was also deployed. Chips and wafers
were examined by ESEM, confocal laser-scanning microscopy (CLSM) for FISH and

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Role of micro-organisms in karstification 353

laser ablation (LA)-ICP-MS. A parallel interaction between the habitat and the micro-
bial communities is the control exerted by the geological matrix, and colonization and
weathering patterns from the carbonate surfaces were compared to other mineral
substrates, such as chert or insoluble residues specific to the habitat being studied.
Results from other environments (Hiebert & Bennett, 1992; Rogers et aL, 1998; Bennett
et aL, 2000; Roberts et aL, 2004) suggest that colonization of a particular mineral
substratum is influenced directly by the chemical composition of that mineral.

Unpreserved chips were examined by using a Philips XL30 ESEM; chamber conditions
varied from 10 to 95 % relative humidity, using a Peltier cooling stage, and variable water
pressure [0-9-64 torr (120-853 Pa), with corresponding accelerating voltages from 4 to
20 kV]. For FISH, chips and wafers were fixed in two ways within 12 h of collection: (i)
with 4 % paraformaldehyde for 3 h; or (ii) with 50 % ice-cold ethanol (Engel et aL, 2003).
Fixed samples were air-dried and dehydrated by sequential washes with 50, 80 and
100 % ethanol for 3 min prior to hybridization. Gene probes specific for Lower Kane
Cave epsilonproteobacterial lineages, as well as probes targeting other microbial
groups, were applied to the rock surfaces and surfaces were examined by using CLSM.

The example of Lower Kane Cave

S-oxidizing bacteria are an underappreciated geological weathering force, particularly
in subsurface karst systems. In simple microbial systems, such as planktonic organisms
in groundwater, the production of acidity can accelerate rock weathering, but in
the complex, stratified microbial-mat communities that can attach to rock surfaces, the
microbial influences on weathering can be more difficult to elucidate. Influences may
involve local production of H 2 S or other S gases and transfer to the vapour phase, with
the net effect of transferring protons (or dissolution capacity) from aqueous to sub-
aerial environments.

Lower Kane Cave is located in the Bighorn Basin near Lovell, WY, USA, adjacent to
oilfields and localized sulfidic thermal and non-thermal springs (Egemeier, 1981). The
cave is formed in the Little Sheep Mountain Anticline along the Bighorn River in the
Madison Limestone and is ~350 m long. Thick, white, filamentous microbial mats,
3-10 cm thick and interconnected with white, web-like films (Fig. 1), are associated
with four sulfidic springs that discharge into the cave, and mats stretch for up to 20 m in
the cave stream below the springs (Engel et aL, 2003, 2004a, b). The biomass of the
filamentous mats was 10 12 cells ml -1 , based on the total organic C-estimate method.

The cave's springs are all of the calcium/bicarbonate/sulfate water type. Although the
cave is forming from sulfuric acid speleogenesis, the spring and stream pH is buffered to
circumneutral by ongoing carbonate dissolution. The dissolved sulfide concentration of

SGM symposium 65

354 P. C. Bennett and A. S. Engel

Fig. 1. Photograph of subaqueous microbial mats in Lower Kane Cave. Image width is 50 cm, water
depth is 10 cm.

incoming spring water was ~38 ^mol 1 (speciated as ~60 % HS~:40 % H 2 S, based on
the pH of the springs, pK 7-04), with non-detectable DO. The concentration of DO and
sulfide changed downstream, such that, at the end of the microbial mats, DO exceeded
40 ^moll -1 and sulfide was non-detectable. DOC (including methane) in all of the
waters was extremely low, at < 80 jimol l -1 .

Microbial-mat diversity, characterized by 16S rDNA clone library construction and
sequence analyses, revealed the presence of several different bacterial phyla. The
surface of the mats, predominantly of white filament bundles, was oxygenated based on
microelectrode DO profiles, whereas the grey mat interior was anoxic ~3 mm beneath
the surface of the mat-water interface. The majority of the sequences retrieved from
the white filaments belonged to the phylum Proteobacteria, specifically the classes
'Epsilonproteobacteria' (68%), 'Gammaproteobacteria' (12-2%), 'Betaproteo-
bacteria 9 (11-7 %) and 'Deltaproteobacteria' (0-8 %), as well as other bacterial phyla,
including the class Acidobacteria (5-6%) and the BacteroideteslChlorobi groups
(1-7 %). In comparison, ~50 % of the retrieved 16S rDNA clone sequences from the
interior mats were related to groups of SRB affiliated to the class 'Deltaproteobacteria'
and to unculturable members of the phylum Chloroflexi. Rarer clones belonged to the
phyla 'Gammaproteobacteria\ Planctomycetes, BacteroideteslChlorobi, 'Beta-
proteobacteria\ 'Epsilonproteobacteria', Verrucomicrobia, ' Alphaproteobacteria\
Spirocbaetes, Actinobacteria, Acidobacteria and different OP candidate groups. FISH
of the microbial mats revealed that ~70 % of the total bacterial biovolume was
dominated by filamentous epsilonproteobacteria and specifically dominated by one
novel lineage, referred to as LKC group II (Engel et ai, 2003) (Fig. 2). The microbial

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Role of micro-organisms in karstification 355

Fig. 2. FISH image of microbial-mat sample, (a) Overlapping probes; green, EUB338l-lllmix probes for
all eubacteria; light blue, GAM42a for gammaproteobacteria (e.g. Thiothrix species). Bar, 20 urn. (b)
EUB338l-lllmix-targeted filaments, being mostly epsilonproteobacteria belonging to the Lower Kane
Cave group II evolutionary lineage (Engel etal., 2003). (c) GAM42a probe only. All images were taken
by using CLSM. Photograph scale is the same in all images.

SGM symposium 65

356 P. C. Bennett and A. S. Engel

mats from Lower Kane Cave represent the first non-marine natural system demon-
strably driven by the activity of filamentous epsilonproteobacteria.

The high concentrations of S(-II) in the cave springs provide a rich energy source for
SOB, and the concentration of dissolved sulfide decreased rapidly through the micro-
bial mats. Abiotic sulfide autoxidation is extremely slow in disaerobic water at pH ~7-4
[autoxidation half-life was calculated to be > 800 h (Engel et al., 2004a)] and sulfide
volatilization from the water to the cave atmosphere accounts for <8 % of the sulfide
loss in the stream, based on gas-flux experiments (Engel et al., 2004a). With no other
mechanism for S(-II) loss, there would be, for example, much higher sulfide concen-
trations at the end of the microbial mats. The only other loss mechanism, microbial
consumption under microaerophilic conditions, could cause the observed rapid loss
in S(-II). Because epsilonproteobacterial filaments were found to dominate the micro-
bial mats, these organisms are presumed to be SOB. Future culturing of these microbial
groups will verify these observations.

The epsilonproteobacteria influence mat redox chemistry and ecosystem function
directly by colonizing the initially nutrient-poor habitat, providing an energy source
through chemolithoautotrophic C fixation, forming a dense mat and consuming 2 .
Chemolithoautotrophy in the cave serves as the base for the cave food web and bulk
white mats had a mean <5 13 C value of -36 %o, demonstrating chemolithoautotrophic
fractionation against 13 C from an inorganic C source in the cave-stream water of
-8-9 %o (Engel et al., 2004b). Most SOB and the epsilonproteobacterial filaments form
complex microbial mats and consumption of DO by the SOB creates an anaerobic
habitat within the mat interior (Engel et al., 2004b). The MPN method was used to
estimate the biomass of anaerobic metabolic guilds; up to 10 6 cells ml -1 were found,
with SRB and fermenters being the dominant culturable groups; iron reducers, S°
reducers and methanogens were in low abundance overall (<10 3 cells ml -1 ). Redox
stratification of the mats spatially separates metabolic guilds, such as SOB from SRB,
and nutrients are consequently cycled between the multiple ecosystem components,
whilst also being advected downstream. Based on the full-cycle rRNA approach,
microbial diversity is greater within the mat interior and downstream portions of the
mat, as mat density and organic C availability increase due to chemolithoautotrophic
input and heterotrophic C degradation. The mat possibly terminates when H 2 S is
consumed and 2 is too high for SOB metabolic function.

The dominant mechanism for S(-II) loss is subaqueous microbial oxidation and most
SOB oxidize reduced S compounds completely to sulfate, with a substantial energy
yield (e.g. equation 3). As a result of the energetic oxidation reactions, acidity is gener-
ated in the form of sulfuric acid, which can attack the geological matrix supporting the

SGM symposium 65

Role of micro-organisms in karstification 357

Fig. 3. ESEM image of a limestone fragment collected from the thick microbial-mat region. The
surface is deeply corroded, consistent with chemical weathering. Smooth crystals are authigenic
gypsum crystals. Filamentous microbial biofilm is barely visible covering the specimen under the

imaging conditions used. Bar, 5 urn.

microbial community. Although some SOB can be acidophiles, most are neutrophilic
and colonization of carbonate surfaces or habitats may buffer excess acidity and
maintain pH homeostasis. Consequently, observed deeply corroded native carbonate
rocks in the cave stream (Fig. 3), with dissolution effects only on surfaces exposed to the
stream water and filamentous microbial mats, can be attributed to the activity of SOB.
Examination of the surfaces by CSEM and ESEM revealed a complex reacting environ-
ment, with dissolving carbonate, secondary gypsum mineral deposits and a thick cover
of predominantly filamentous microbes in an exopolysaccharide matrix.

The results from the experimental in situ microcosm chips show that the filamentous
sulfide-oxidizing bacteria colonize the calcite surface (Fig. 4a). Over time, the sulfide
oxidation results in chemical corrosion of the surface, first producing shallow, etched
trenches along the filament (Fig. 4b) and culminating in a broadly etched surface
resembling the native limestone fragments (Fig. 5).

FISH probes were applied to the experimental microcosm carbonate surfaces to
determine the identities of the filaments colonizing the surfaces (Engel et aL, 2004a).
Gene probes targeting the LKC group II (' Epsilonproteobacteria') , 'Gammaproteo-
bacteria' (e.g. Tbiotbrix species) and all eubacteria were used. Positive hybridization
signals were observed for filamentous organisms and most of the filaments on the
carbonate surfaces hybridized simultaneously with the general eubacterial probe and

SGM symposium 65

358 P. C. Bennett and A. S. Engel

Fig. 4. ESEM image of chips of calcite collected from in situ microcosms exposed to microbial
colonization for 3 months, (a) Image of S-oxidizing Thiothrix sp. filament with abundant stored
elemental sulfur, (b) Thiothrix sp. filament on calcite surface, with visible etching of the mineral
along a filament trench forming in the mineral. Bars, 5 urn.

the LKC group II probe. Exceptionally bright hybridization signals for each of the
probes indicated high rRNA content and suggested that the microbes were active when
the microcosms were retrieved.

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Role of micro-organisms in karstification 359

Fig. 5. CSEM photomicrograph of calcite chip from in situ microcosm, showing dissolution of the
mineral surface (lower left) in association with filamentous micro-organisms that formed a thick
biofilm on the mineral surface. Bar, 20 urn.

The rapid loss of sulfide from the cave stream, carbonate dissolution associated with
microbial filaments and the dominance of epsilonproteobacteria on the experimental
carbonate surfaces support the hypothesis that these organisms, as SOB, have a direct
role in sulfuric acid speleogenesis in Lower Kane Cave. The epsilonproteobacteria
generate sulfuric acid as a by-product of their metabolism and locally depress the pH at
limestone surfaces, which subsequently focuses carbonate dissolution. The previous
cave-formation model was based on H 7 S volatilization from the cave stream to the
atmosphere, but negligible volatilization and abiotic autoxidation of S(-II) were found
in the cave. Instead, S(-II) was consumed by subaqueous SOB and cave enlargement
occurs via microbially enhanced dissolution of the cave floor.

CONCLUSIONS

The subsurface is a varied habitat for diverse microbial communities, as reactive rock
surfaces and mineral-rich groundwater provide a variety of energy sources for micro-
organisms, especially chemolithoautotrophs. Stream geochemistry and the spatial
relationships of aerobic and anaerobic guilds allow for tight cycling of C and S,
constrained by redox boundaries within the mat environment. In Lower Kane Cave, the
bulk of the subaqueous microbial mats were dominated by novel evolutionary lineages
of the class 'Epsilonproteobacteria', representing the first non-marine system driven by
the activity of this group. Ecologically, the epsilonproteobacteria are chemolitho-
autotrophic SOB and their genetic and metabolic diversity creates habitats within the
microbial-mat interior for other microbial groups, such as SRB and other anaerobic
metabolic guilds.

SGM symposium 65

360 P. C. Bennett and A. S. Engel

Nearly all of the S(-II) coming into the Lower Kane Cave is consumed by SOB within
the subaqueous microbial mats, and these bacteria drive subaqueous sulfuric acid
speleogenesis by attachment to carbonate surfaces and generation of sulfuric
acid, which focuses local carbonate undersaturation and dissolution. Prior to this work,
sulfuric acid speleogenesis was considered to be an abiotic, subaerial process or to be
limited to shallow groundwater depths because of oxygen requirements for abiotic
autoxidation of S(-II). However, chemolithoautotrophy linked to S(-II) oxidation
under microaerophilic conditions extends the phreatic depths to which porosity and
conduit enlargement can occur. The recognition of the geomicrobiological contri-
butions to subaqueous carbonate dissolution fundamentally changes the model for
sulfuric acid speleogenesis and supplements how subsurface porosity may develop in
karst aquifers.

ACKNOWLEDGEMENTS

We thank the US Bureau of Land Management for permission to access this field site. We are
grateful to M. L. Porter, S. A. Engel, T. J. Dogwiler, K. Mabin, M. Edwards, R. Payn, J. Deans and
H. H. Hobbs, III, for field assistance. Funding for this work was provided by the Life in Extreme
Environments (LExEn) program of the US National Science Foundation (EAR-0085576), the
National Speleological Society, the Geological Society of America and the Geology Foundation
of the University of Texas.

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SGM symposium 65

INDEX

References to tables/figures are shown in italics

Acetate oxidation 191
acetogenesis 111,189,241
Acbarax 309
Acbromatium 35—37

A. oxaliferum 36

A. uolutans 53

calcite deposition 51,52,53

genetic and ecological diversity 60—61, 62, 63

RY5, RYKS, RY8 spp. 61

sulfur cycle 46,50,51,348
acid mine drainage (AMD) 12, 14, 23, 25, 26
Acidobacteria 354

acidolysis, by fungi 207, 208-211, 214, 215, 218
Acineto b acte r sp. B0064 117,118
Acremomium -like hyphomycete 215—216
actinides, microbial reduction 286—291
Actinobacteria 354
Aeromonas 277 ', 279

fungal interactions 211, 215

microbial reduction 282

microbial resistance 115,273
aggregation

diffusion-limited (DLA) 142-143

reaction-limited (RLA) 142-143
Al, fungal interactions 212, 216
' Alpbaproteobacteria' 354
Alternaria alternata 337
Alteromonas putrefaciens see Sbewanella

one id en sis
americium (Am), microbial reduction 289
ammonia

anaerobic oxidation 152,164

marine cycling 250—251
ammonium, in primary production 181
anaerobic methanotrophs (ANMEs) 308, 310, 311,

312,313,314,316
Antarctic dry valleys 71—73

location 72

microbial activity in 73, 75—76

organisms in 73, 74

sources of resources 76—81
AOM see methane (anaerobic oxidation)
apatite 211
archaeol 307
archaeal 'strain 121' 277
Artbrobacter 113,115,117,119
As

fungal interactions 215

groundwater contamination 283—284

microbial influences 22

microbial reduction 164,283—286

microbial resistance 273, 284
Ascocratera manglicola 333
ascomycetes 206, 326, 327, 337
Aspergillus 218,220,330

A. fumigatus 325

A. nidulans 215

A.niger 208,218

A. sydowii 325
P37 sp. 215

Astrospbaeriella striatispora 333
ATPases, P IB -type 115-122
Aureobasidium pullulans 334
autunite 212
azo dyes 293-294
Azotobacter vinelandii 120

Bacillus

B. arseniciselenatis 285
B. selenitireducens 285
B. sp. U26 123
B.sp.V6 123-124

B. subtilis 87-88, 89-90
bacterial cell wall 85-104

cell surface complexity 94—99

cryo-transmission electron microscopy 86—93

hydrophobicity studies 99—104

metal— ion interaction and mineralization
91-93

potentiometric properties 93—99
'Bacteroidetes' 348, 354
basalt 205, 208, 234
Basidiomycotina 327
Beauveria caledonica 209, 210, 212-213
Beggiatoa 37-38

B.alba 45

carbon cycle 55, 56

nitrogen cycle 53, 54

sulfur cycle 44,45,46,47,309

within-genus diversity 57—59
benthic communities

organic matter content 1 78

organic matter degradation 1 78, 179, 180,
182-183, 188-191

oxygen uptake 182

respiration 186
bentonite 141,217,215,219

SGM symposium 65

366 Index

beta-imaging 5
'Betaproteobacteria* 354
biocatalysts 273,281,283-294
biofilms

existence, structure and function 22—23, 24
formation 24

in mycotransformation of minerals 205, 206
intraterrestrial environment 239
metabolic links to metals 23
metal geochemistry interactions 11—28
aqueous, fundamentals 12—18
emerging technologies 19,27—28
environmental aspects 23, 24, 25, 26
important solid phases 13—15
microbial functional metabolism 19—22
role of redox status 18
sorption reactions 15, 1 6, 17— 18
biogeochemical cycling, in coastal zones 173—193
functional groups 188—191
organic matter degradation and recycling

182-183
physico-chemical differences in latitudinal

regions 173—175
primary production 174, 175—182, 183
role of coastal wetlands 191—193
temperature— substrate concentration
interaction 183-188
biogeochemical cycling, marine ecosystems
fungal contribution 336—338
role of fungi 330—336
C,NandP 331-335
sulfur 331,332,335-336
biomagnetite 160—161
biomarkers, lipid 133,306—308
biomass, microbial

in intraterrestrial environment 239
marine, resource availability 261—265
subsurface microbial mats 110, 351, 353, 356
biomineralization

accidental biominerals 156, 157—158
bacterial cell wall interactions see bacterial

cell wall
biofilms 23—26
biomagnetite 160—161
eukaryotic 159-160
prokaryotic 157-159, 160
purposeful biominerals 158, 159—160
see also biosilification, mycotransformation
bio remediation
fungal role 220
metal reduction 273, 278, 279, 282, 287, 289,

291
metal resistance 112—114
radioactive waste 280, 286-291

biosensors, Hg 282
biosignatures 165—166, 216
biosilification 131—145

chemistry of silica 133—138
opal-A 137,138

precipitation rate, calculation 136
saturation state 135—136, 137
cyanobacterial biomineralization pathways

and colloid aggregation 141—144
cyanobacterial surface properties and
function 138-141
biotite 208-209,216
bioweathering, defined 203

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