Life at Vents and Seeps 9783110493672, 9783110494754

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Jens Kallmeyer (Ed.) Life at Vents and Seeps Life in Extreme Environments

Life in Extreme Environments

| Edited by Dirk Wagner

Volume 5

Life at Vents and Seeps |

Editor Dr. Jens Kallmeyer GFZ – German Centre for Geosciences Telegrafenberg, 14473 Potsdam Germany E-mail: [email protected]

ISBN 978-3-11-049475-4 e-ISBN (PDF) 978-3-11-049367-2 e-ISBN (EPUB) 978-3-11-049211-8 ISSN 2197-9227 Library of Congress Cataloging-in-Publication Data A CIP catalog record for this book has been applied for at the Library of Congress. Bibliographic information published by the Deutsche Nationalbibliothek The Deutsche Nationalbibliothek lists this publication in the Deutsche Nationalbibliografie; detailed bibliographic data are available on the Internet at http://dnb.dnb.de. © 2017 Walter de Gruyter GmbH, Berlin/Boston Cover image: Minyang Niu, Shanghai, China Typesetting: le-tex publishing services GmbH, Leipzig Printing and binding: Hubert & Co. GmbH & Co. KG, Göttingen ♾ Printed on acid-free paper Printed in Germany www.degruyter.com

Preface “What are extreme environmental conditions?” Most of the answers that one gets from undergraduate students confirm a rather anthropocentric view: An environment that humans perceive as unpleasant is classified as extreme. Many of Earth’s ecosystems are characterized by “extreme” environmental conditions, because they deviate from those conditions that humans would consider “normal” with regard to temperature, water availability, pressure, salinity, nutrient supply and so on. Vents and Seeps, especially deep-sea hydrothermal vents, are synonyms for life under extreme conditions because they exhibit everything that we, as surfacedwelling humans, consider extreme: high temperature, high pressure, in some cases high concentrations of toxic compounds. However, there are many more environments that are characterized by moving fluids, as different as their geologic setting, temperature regime, flow rates and fluid compositions might be. What they all have in common are very steep chemical and often thermal gradients as well as high concentrations of chemical compounds. Despite being considered extreme, these habitats are colonized by a large number of organisms that thrive under the given conditions. In recent times, more and more scientists from various disciplines became interested in vent and seep environments. Through multidisciplinary research completely new concepts were developed of how life can possibly survive and even thrive in extreme ecosystems, mainly due to the development of new analytical techniques. The different chapters in this volume of the series “Life in Extreme Environments” provide an overview of the wide range of different vent and seep environments and their inhabitants. The first chapters describe the various habitats, followed by studies about specific locations. The final chapters present exciting new technological developments as well as studies about specific microbial groups that are commonly found in vent and seep environments. I hope that this book will provide you with an up-to-date overview of the topic and spark your interest in these truly extreme ecosystems. Jens Kallmeyer

https://doi.org/10.1515/9783110493672-201

Volumes published in the series Volume 1 Jens Kallmeyer, Dirk Wagner (Eds.) Microbial Life of the Deep Biosphere ISBN 978-3-11-030009-3

Volume 2 Corien Bakermans (Ed.) Microbial Evolution under Extreme Conditions ISBN 978-3-11-033506-4

Volume 3 Annette Summers Engel (Ed.) Microbial Life of Cave Systems ISBN 978-3-11-033499-9

Volume 4 Blaire Steven (Ed.) The Biology of Arid Soils ISBN 978-3-11-041998-6

Contents Preface | V Contributing authors | XIII Stefan Krause, Helge Niemann, and Tina Treude 1 Methane seeps in a changing climate | 1 1.1 Introduction | 1 1.2 What is a cold seep and what drives it? | 1 1.2.1 Methane seeps | 2 1.2.2 Oil seeps | 2 1.2.3 Pockmarks | 3 1.2.4 Mud volcanoes | 4 1.2.5 Brine seeps | 4 1.3 Seep chemistry and biology | 5 1.4 Cold seeps and gas hydrate | 7 1.5 Gas hydrates, methane bubbles | 9 1.6 Anaerobic methane oxidation, sulfide oxidation, and large seep fauna | 10 1.7 Carbonate depositions at cold seeps | 12 1.8 Effects of climate change on methane seeps and vice versa | 13 1.9 Methane and past climate | 13 1.10 Cold seeps and recent global climate change | 14 1.11 Heat transfer from the atmosphere to marine sediment | 14 1.12 Gas hydrate destabilization in the near future | 16 1.13 Areas most affected by gas hydrate destabilization | 18 1.13.1 The Arctic | 18 1.13.2 Blake Plateau | 19 1.14 The fate of methane | 19 1.14.1 Methane in the sediment | 20 1.14.2 Methane in the water column | 20 1.15 Methane seeps – relevant for climate change? | 21 1.16 Atmospheric methane concentration over time | 22 1.17 Closing remarks | 23

VIII | Contents

Samantha B. Joye and Sara Kleindienst 2 Hydrocarbon seep ecosystems | 33 2.1 Overview | 33 2.2 Introduction – discovery of hydrocarbon seeps | 33 2.3 Geology of hydrocarbon seeps | 35 2.3.1 Types and formation mechanism | 35 2.3.2 Gas hydrates | 37 2.3.3 Authigenic carbonates | 39 2.4 Biology and biogeochemistry of hydrocarbon seeps | 40 2.4.1 Biogeochemistry | 40 2.4.2 Biology | 41 2.4.3 Microbiology | 42 2.5 Closing remarks | 46 Walter Menapace, Achim Kopf, Matthias Zabel, and Dirk de Beer 3 Mud volcanoes as dynamic sedimentary phenomena that host marine ecosystems | 53 3.1 Abstract | 53 3.2 Introduction | 53 3.2.1 Types of MV, expelled products, morphologies and dimensions | 53 3.2.2 Identification on the seafloor and classification | 55 3.3 Geological significance | 57 3.3.1 MV number estimation and distribution | 57 3.3.2 Formation models and long-term evolution | 58 3.4 Mass transfer and fluid cycling fuel ecosystems | 61 3.4.1 Composition and sources of gas, water, mud, and clasts | 61 3.4.2 Seafloor and subseafloor ecosystems on MVs | 63 3.4.3 MV episodicity | 66 3.4.4 MV short term evolution | 67 3.4.5 Methane in MVs: the Håkon Mosby case study | 69 3.5 Conclusion | 73 Solveig I. Bühring and Stefan M. Sievert 4 The shallow submarine hot vent system off Milos (Greece) – a natural laboratory for the study of hydrothermal geomicrobiology | 85 4.1 Abstract | 85 4.2 Introduction | 85 4.2.1 Background | 85 4.2.2 Shallow marine systems: overview | 86 4.3 Milos – an extreme environment as a model system | 88 4.3.1 Geophysical and geochemical conditions | 89 4.3.2 Seabed features | 91

Contents | IX

4.4 4.5

General aims in hydrothermal geomicrobiology and how they could be addressed by using Milos as a natural laboratory | 92 Conclusions and future directions | 100

Matthew O. Schrenk 5 Life in serpentinite hosted alkaline springs | 107 5.1 Abstract | 107 5.2 Introduction to serpentinization influenced ecosystems | 107 5.2.1 Serpentinization is common | 108 5.2.2 Chemical reactions associated with serpentinization | 109 5.2.3 Inorganic geochemistry of serpentinites | 111 5.2.4 Hydrogeology of serpentinite springs | 113 5.2.5 Interfaces associated with serpentinization | 114 5.2.6 Mediators of carbon exchange | 114 5.3 Where do serpentinization processes occur? | 116 5.3.1 Submarine systems | 116 5.3.2 Continental serpentinite springs | 118 5.3.3 Wells in serpentinization influenced aquifers | 119 5.3.4 Serpentine cores | 121 5.3.5 Serpentine soils | 122 5.4 Constraints upon the microbial ecology of serpentinite habitats | 122 5.4.1 Challenges of hyperalkaliphily | 124 5.4.2 “Plenty to eat, nothing to breathe . . . ” | 124 5.4.3 Carbon availability | 126 5.5 Are there endemic species in serpentinite hosted systems? | 127 5.5.1 Serpentinomonas, here, there, and everywhere? | 127 5.5.2 What about the Clostridiales? | 128 5.6 Evidence of activity and function | 128 5.7 Relevance to early Earth, deep biosphere and mineral carbonation | 129 5.7.1 Carbon sequestration in serpentinites | 130 5.8 Summary and future perspectives | 131 Mingyang Niu, Qianyong Liang, Dong Feng, and Fengping Wang 6 Ecosystems of cold seeps in the South China Sea | 139 6.1 Introduction | 139 6.2 Framework geology | 140 6.2.1 Southwest of Taiwan | 141 6.2.2 Dongsha area | 143 6.2.3 Shenhu area | 144 6.2.4 Qiongdongnan Basin | 144 6.2.5 Xisha area | 145

X | Contents

6.3 6.4 6.4.1 6.4.2 6.5

Macroecology in cold seeps of the northern South China Sea | 145 Microbial community | 148 Formosa Ridge | 148 Jiulong Reef and Haima Seep | 149 Outlook | 154

Jian Ding and Yu Zhang 7 Life at the hydrothermal vent field of the Southwest Indian Ridge | 161 7.1 Overview of the hydrothermal vents | 161 7.2 Hydrothermal vents on the Southwest Indian Ridge | 163 7.3 The biologic communities distributed on the SWIR | 164 7.3.1 Faunal communities | 164 7.3.2 Microbial communities | 165 7.3.3 A case study on the microbial communities at the Longqi field | 167 Scott D. Wankel, Annie Bourbonnais, and Chawalit Charoenpong 8 Microbial nitrogen cycling processes at submarine hydrothermal vents | 179 8.1 Introduction | 179 8.1.1 Hydrothermal fluid venting and mixing | 180 8.1.2 Nitrogen species found in hydrothermal vent fluids | 180 8.2 Energetic considerations | 183 8.3 Microbially catalyzed nitrogen cycling processes | 186 8.4 Stable isotopes as indicators of microbial processes | 188 8.5 Genetic evidence for nitrogen cycling processes at vents | 192 8.6 Nitrogen fixation | 193 8.7 Nitrogen assimilation | 195 8.8 Denitrification and anammox | 197 8.9 Dissimilatory nitrate reduction to ammonium (DNRA) | 202 8.10 Nitrification | 204 8.11 Concluding remarks and future directions | 209 Jeffrey Marlow and Roland Hatzenpichler 9 Assessing metabolic activity at methane seeps: a testing ground for slow growing environmental systems | 223 9.1 Introduction | 223 9.2 Observational approaches to quantifying seep hosted activity | 225 9.2.1 Modeling rates of activity from geochemical profiles | 225 9.2.2 Colonization Rates | 226 9.3 Catabolism based methods | 227 9.3.1 Tracking methane catabolism | 227 9.3.2 Tracking sulfur catabolism | 232

Contents | XI

9.3.3 9.4 9.4.1 9.4.2 9.4.3 9.5 9.5.1 9.6 9.6.1 9.6.2 9.6.3 9.6.4 9.6.5 9.6.6 9.7

Tracking catabolism of ‘nontraditional’ electron acceptors | 233 Anabolism based methods | 233 Cell quantification | 233 Stable isotope probing | 234 BONCAT | 239 Key outstanding issues and challenges | 240 Linking the lab with the real world | 240 Metabolic activity and tools of the future | 243 Preserving microscale spatial arrangements | 244 Single-cell growth rate | 244 Microcalorimetry | 245 Raman spectroscopy | 246 Replication rates from metagenomic data | 246 Single-cell omics | 247 Conclusions | 247

Barbara J. MacGregor, Beverly Flood, Jake Bailey, and Matthew Kanke 10 Multiplication is vexation: a genomic perspective on cell division and DNA replication in the large sulfur bacteria | 261 10.1 Abstract | 261 10.2 Introduction | 262 10.2.1 Overview of sequenced large sulfur bacteria | 265 10.3 Septation | 266 10.3.1 The division and cell wall (dcw) gene cluster is fragmented in several of the large sulfur Beggiatoaceae compared to close relatives | 266 10.3.2 The available Thiomargarita genomes are missing genes for septum formation | 267 10.3.3 What might substitute for ZipA and FtsA as an FtsZ membrane anchor in Thiomargarita septum formation? | 276 10.3.4 How might central vacuoles divide? | 282 10.4 Disruption of some gene clusters seems characteristic of marine LSB | 283 10.5 DNA replication | 287 10.5.1 LSB are lacking some or all of the typical gene features of bacterial replication origins | 287 10.5.2 The two complete LSB genomes lack GC skew | 291 10.5.3 Parallels to Cyanobacteria | 296 10.5.4 Alternatives to DnaA | 299 10.6 Summary and perspectives | 300

XII | Contents

Weishu Zhao and Xiao Xiang 11 Life in multiextreme environments: cross-stress response in Thermococcales | 307 11.1 Abstract | 307 11.2 Introduction and background | 307 11.3 Basic characteristics of Thermococcales | 308 11.3.1 Physiological traits | 308 11.3.2 Genome traits | 313 11.4 Adaptation related pathways of Thermococcales | 314 11.4.1 Energy conversion | 315 11.4.2 Amino acid metabolism | 316 11.4.3 Compatible solute | 317 11.4.4 Composition of membrane lipid | 319 11.4.5 Antioxidant pathway | 320 11.5 Common adaptation strategies for different stresses | 321 11.5.1 Common adaptation strategy | 322 11.5.2 Summary of adaptation strategy | 324 Index | 331

Contributing authors Jake Bailey Department of Earth Sciences University of Minnesota Minneapolis, USA e-mail: [email protected]

Annie Bourbonnais Woods Hole Oceanographic Institution Department of Marine Chemistry and Geochemistry Woods Hole, Massachusetts USA e-mail: [email protected]

Solveig I. Bühring Marum – Zentrum für Marine Umweltwissenschaften Universität Bremen Bremen, Germany e-mail: [email protected]

Chawalit “Net” Charoenpong Woods Hole Oceanographic Institution Department of Marine Chemistry and Geochemistry Woods Hole, Massachusetts USA e-mail: [email protected]

Dong Feng CAS Key Laboratory of Ocean and Marginal Sea Geology South China Sea Institute of Oceanology Chinese Academy of Sciences Guangzhou, China e-mail: [email protected]

Beverly Flood Department of Earth Sciences University of Minnesota Minneapolis, USA e-mail: beverlyfl[email protected]

Ronald Hatzenpichler Montana State University Department of Chemistry and Biochemistry Bozeman, MT, USA e-mail: [email protected]

Samantha Joye Department of Marine Sciences The University of Georgia Marine Sciences Bldg. Athens, Spain e-mail: [email protected]

Dirk de Beer MPI für Marine Mikrobiologie Bremen, Germany e-mail: [email protected]

Sara Kleindienst Universität Tübingen Angewandte Geowissenschaften Hölderlinstraße 12 72074 Tübingen Germany e-mail: [email protected]

Jian Ding School of Life Sciences and Biotechnology State Key Laboratory of Microbial Metabolisms Shanghai JiaoTong University Shanghai, China e-mail: [email protected]

Achim Kopf Marum – Zentrum für Marine Umweltwissenschaften Universität Bremen Bremen, Germany e-mail: [email protected]

https://doi.org/10.1515/9783110493672-202

XIV | Contributing authors Stefan Krause Marine Geosystems GEOMAR Helmholtz Centre for Ocean Research Kiel Germany e-mail: [email protected]

Mingyang Niu School of Life Sciences and Biotechnology State Key Laboratory of Microbial Metabolisms Shanghai Jiao Tong University Shanghai, China e-mail: [email protected]

Matthew Kanke Cornell University College of Veterinary Medicine Box 17 Ithaca, NY 14853, USA e-mail: [email protected]

Matt Schrenk Michigan State University Department of Earth and Environmental Sciences Department of Microbiology and Molecular Genetics Natural Science Building East Lansing, USA e-mail: [email protected]

Qianyong Liang Guangzhou Marine Geological Survey China Geological Survey Guangzhou, China e-mail: [email protected] Barbara MacGregor University of North Carolina Department of Marine Sciences Chapel Hill, USA e-mail: [email protected] Jeffrey Marlow Harvard University Department of Organismic and Evolutionary Biology Cambridge, MA, USA e-mail: [email protected] Walter Menapace Marum – Zentrum für Marine Umweltwissenschaften Universität Bremen Bremen, Germany e-mail: [email protected]

Helge Niemann Department. of Environmental Sciences Aquatic and Stable Isotope Biogeochemistry University of Basel Basel, Switzerland e-mail: [email protected]

Stefan M. Sievert Biology Department Woods Hole Oceanographic Institution Woods Hole, USA e-mail: [email protected] Tina Treude Geology Building Los Angeles, California, USA e-mail: [email protected] Fengping Wang School of Life Sciences and Biotechnology State Key Laboratory of Microbial Metabolisms Shanghai Jiao Tong University Shanghai, China e-mail: [email protected] Scott D. Wankel Woods Hole Oceanographic Institution Department of Marine Chemistry and Geochemistry Woods Hole, Massachusetts USA e-mail: [email protected] Xiang Xiao School of Life Sciences and Biotechnology State Key Laboratory of Microbial Metabolisms Shanghai JiaoTong University Shanghai, China e-mail: [email protected]

Contributing authors

Matthias Zabel Naturwissenschaften 2 Universität Bremen Bremen, Germany e-mail: [email protected] Weishu Zhao School of Life Sciences and Biotechnology State Key Laboratory of Microbial Metabolisms Shanghai JiaoTong University Shanghai, China e-mail: [email protected]

| XV

Yu Zhang School of Life Sciences and Biotechnology State Key Laboratory of Microbial Metabolisms Shanghai JiaoTong University Shanghai, China e-mail: [email protected]

Stefan Krause, Helge Niemann, and Tina Treude

1 Methane seeps in a changing climate 1.1 Introduction Areas of fluid seepage occur globally in the marine realm, predominantly on continental shelves, margins and subduction zones. Here, ascending fluids become charged with hydrocarbons, especially methane, on their way up to the sediment surface. Since their discovery, methane seeps have been the subject of intensive multidisciplinary research approaches, aiming to constrain the fate of CH4 including its genesis, dynamics of sedimentary deposition, microbial depletion, and the transfer from the oceanic realm into the atmosphere, where it acts as a potent greenhouse gas [1]. According to recent calculations, increasing CH4 concentrations in the atmosphere from combined natural and anthropogenic sources might have been responsible for about 20% of the global warming since preindustrial times [2]. In order to constrain the methane budget for climate predictions all relevant sources have to be evaluated, including seep locations. Currently, the annual methane emission from marine seeps to the atmosphere is anticipated to be comparatively small. However, estimates show considerable uncertainties spanning three orders of magnitude (0.4– 65 Tg CH4 yr−1 ) [3, 4].

1.2 What is a cold seep and what drives it? The earliest reports of cold seeps in the marine realm date back almost 2,000 years, when inhabitants of the Syrian Island Aradus obtained fresh water from a submarine spring. However, cold seeps are far less well known to the public as the famous ‘black smokers’, which were only discovered in 1977. Cold seeps (sometimes referred to as cold vents) represent areas where gases and fluids migrate through the sediment and discharge into the overlying water body. The term ‘cold’ implies that the emanating fluid is cooler than the adjacent bottom water. However, this is usually not the case. In contrast, the fluid temperature is usually slightly higher than that of the adjacent bottom water [5], but considerably cooler compared to hot vents. The morphology of cold seeps shows considerable variation as a result of the different fluid sources and formation mechanisms [6]. Although the vast majority of seeps are associated with emanating fluids, locations featuring gas ebullition without advecting fluid are also often categorized as seeps.

https://doi.org/10.1515/9783110493672-001

2 | 1 Methane seeps in a changing climate

Cold seeps are predominantly found on continental slopes, which extend more than 320,000 km in length, covering an area of approximately 41 million km2 or 11% of the global ocean area [7]. During their ascent, the fluids may become charged with high salinity, hydrogen sulfide, or hydrocarbons, predominantly methane. The latter is the focus of intensive research in various scientific disciplines. Currently, all global cold seeps may contribute between 1 and 5% to annual global methane emissions into the atmosphere [8–11]. Modern cold seep research started in 1983 in the Gulf of Mexico [12]. Since then, cold seeps have been the subject of intensive research resulting in the still ongoing discovery of numerous seeps at various water depths broadly ranging from the intertidal zone of continental margins down to oceanic trenches in the hadal zone (> 6,300 m) [13, 14]. The vast majority of cold seeps occur on top of geological fissures at active and also passive continental margins, facilitating the ascent of chemically enriched fluids. Depending on the geological setting and the fluid composition, we can distinguish between the following types of seeps: (i) methane seeps, (ii) oil seeps, (iii) pockmarks, (iv) mud volcanoes, and (v) brine seeps.

1.2.1 Methane seeps Methane seeps are encountered typically at the seafloor of continental slopes [15–17]. The leakage of natural gas, predominantly methane, is one of the key features associated with the vast majority of cold seeps. Methane reservoirs have formed along biologically productive continental margins, mainly caused by microbial conversion of refractory organic matter. Tectonic activity at passive and active continental margins, creating a network of fissures, facilitates the ascent of fluids and natural gas from deeper parts of the crust. The advection of methane charged fluids facilitates a diverse endemic faunal community, ranging from microbes to megafauna [18–20], which is essentially dependent on the activity of chemosynthetic microbes. Areas of methane seepage are frequently associated with calcium carbonate deposits, which result from the microbial anaerobic consumption of methane. These deposits are present throughout the Phanerozoic rock record [21], representing unique archives to constrain past fluid flow, methane budgets, and climate variability [22–24].

1.2.2 Oil seeps Crude oil or petroleum seeps are geographically common and can be encountered in the marine realm [25] as well as on land, even including biblical references to asphalt seeps in the vicinity of the Dead Sea [4]. Probably, the world’s most spectacular marine oil seeps are encountered at Coal Oil Point, Santa Barbara Channel, California, with

1.2 What is a cold seep and what drives it?

|

3

hydrocarbon gas and oil emission rates from the seabed of 1.7 ± 0.3 × 105 m3 d−1 and 1.6±0.2×103 L d−1 respectively [26]. Crude oil seepage is the result of reservoir failure due to overpressuring. The formed network of cracks and fissures act as conduits for the uprising oil. The light oil components, which are positively buoyant, may rise to the water surface, while the heavy oil components accumulate at the seafloor for prolonged periods of time [27]. Petroleum seeps are generally an indicator of underlying oilfields expelling an extremely broad range of organic compounds [28]. Recent studies demonstrate that sulfate reduction may play a key role in the degradation pathways of higher hydrocarbons constituting crude oil [29–32]. Similar to methane seeps, locations of emanating petroleum are also commonly associated with a specific fauna, including assemblages of chemoautotrophic clams, exhibiting long-term persistence and resilience [33]. In numerous instances a combined discharge of oil and gas can be observed. An analysis of 141 globally distributed petroleum systems revealed that oil and gas, primarily methane gas, are simultaneously advected at 89% of them [34]. Current estimates of the annual crude oil seepage rate range between 0.2 and 2 mM t, representing about 47% of the oil entering the oceans [25].

1.2.3 Pockmarks Pockmarks occur worldwide and represent crater-like structures or pits at the seafloor. They were originally discovered off the coast of Nova Scotia in the 1960s [35]. The shape of pockmarks is usually subcircular, but elongated and ellipsoidal craters have also been observed [36]. More complicated, asymmetrical structures are the result of two or more pockmarks in close proximity to each other [37]. Although pockmarks have been the subjects of investigations over several decades, their exact formation mechanism still needs to be deciphered. Depending on the geographical region, pockmarks may be caused by initially explosive emissions of water, brine, oil, or gaseous hydrocarbons leading to the translocation of sediment [38, 39]. Alternatively, evidence also exists indicating formation over geological timespans due to the continuous slow leakage of gas and fluid. Recent investigations indicate that vortical flow may be causative for pockmark maintenance and potentially also for pockmark excavation, providing evidence for a postformation evolution in the absence of fluid venting [40]. The dimensions of pockmarks display large variation, spanning from 10 m to up to 1.5 km in diameter [41, 42]. The study of pockmark fields and even individual craters allows for conclusions to be made about changes in the hydraulic activity of the seabed, making them ideal indicators of deep fluid pressure increase prior to earthquakes [42]. Inactive pockmarks do not displace sediment anymore. In contrast, they can act as sediment traps by reducing bottom water currents [43], facilitating an increased settlement of larvae compared to the surrounding sediment [44]. Therefore, inactive pockmarks might represent recruitment areas for bottom-dwelling larvae and species [45].

4 | 1 Methane seeps in a changing climate

1.2.4 Mud volcanoes A mud volcano is not a real volcano according to the geological definition, as it does not expel any lava. Instead, the term refers to a seafloor or terrestrial surface structure resulting from the exudation of mud or sediment slurry, together with gaseous hydrocarbons, primarily methane [46]. The size of a mud volcano may range from less than one meter up to ten kilometers in diameter [47]. Mud volcanoes are associated with a variety of geological features including accretionary prisms at subduction zones or slope failures at continental margins. In areas with high sedimentation rate, liquid is produced as a result of dewatering processes in deeper sections of the sediment. The resulting liquid blends with subterranean sediments forming liquefied mud, which is forced upwards, often carrying gaseous and liquid hydrocarbons. At the seafloor the upward movement of the mud forms characteristic crater-like structures [46]. One of the most intensely studied marine mud volcanoes is Håkon Mosby, named after the famous Norwegian oceanographer. The structure was discovered in 1990 on the Barents Sea continental slope [48], representing an unusual example of venting in an Arctic location and under noncompressional tectonic circumstances [49]. Pioneering work at the Håkon Mosby mud volcano included the first detailed measurements of chemical gradients and methane-related microbial activities at this type of geological feature as the basis for numerical modeling of chemical fluxes [50–52].

1.2.5 Brine seeps The seepage of brine can be observed globally where sediments overlie large-scale salt deposits. The tectonic movements of the salt result in a network of fissures facilitating the rapid loading of fluids with dissolved salt components, often accompanied by oil and/or gaseous hydrocarbons. The emanating brine has a higher density than the ambient seawater, causing the pooling of brine at the seafloor [53]. These hypersaline environments harbor a diverse group of microbial organisms, well adapted to the prevailing chemical challenges. On the basis of fossilized stromatolites it is assumed that halophilic (salt-loving) organisms may have evolved during the Archean age, approximately 3.5 billion years ago [54]. Therefore, the study of microbial communities and metabolic pathways opens a time window onto early Earth environments. Marine brine seeps may also serve as analogues for extraterrestrial environments [55, 56]. The activity, or more exactly the rate and dynamics of fluid advection and consequential fluxes of chemical components, represents one of the most important aspects of seep research. Regardless of the type, seepage activity may undergo considerable temporal variations, exemplified by ephemeral, episodic and/or variable rates over time [57, 58]. Seep locations may also become permanently inactive. This state can be induced by geologic (tectonic) activity, blocking the fluid conduits or the precipita-

1.3 Seep chemistry and biology |

5

tion of cap carbonates as a consequence of prolonged biological (microbial) activity, restricting the diffusion between bottom water and sediment.

1.3 Seep chemistry and biology The chemical composition of advecting fluids at cold seeps differs from pore water of nonseepage areas and from the overlying bottom water. During their passage through the strata of sedimentary rock towards the seafloor, fluids often become highly enriched in hydrocarbons such as natural gas, of which methane represents the dominant fraction (󳶳 Fig. 1.1). Principally, two sources of methane have to be distinguished. Thermogenic methane reservoirs are usually encountered at subbottom depths exceeding 1,000 m [59]. The genesis of thermogenic methane requires high temperatures and the presence of organic carbon compounds. Natural observations combined with experimental approaches have shown that thermogenic methane production involving the breakdown of kerogens (high molecular weight organic compounds) takes place between 60 and about 180 °C [60]. At temperatures of 160 to 220 °C the breakdown of kerogens and asphalt also results in the formation of thermogenic methane [61].

Fig. 1.1: Scheme summarizing the chemical-biological coupling and ecologically relevant processes at cold seep sites with hydrocarbon-charged fluid advection. (Image source: SFB574, GEOMAR Kiel)

6 | 1 Methane seeps in a changing climate

By contrast, microbial methane is mostly formed at much lower temperatures (< 80 °C) [62, 63] and at shallower sediment depths. However, active methanogens have also been identified at temperatures exceeding 100 °C and as part of the deep biosphere microbial community [64]. The type of microorganism carrying out biogenic methanogenesis belongs to the taxonomic domain of archaea, representing singlecelled prokaryotic microorganisms, which are distinct from bacteria [65]. Microbial CH4 can be produced from competitive and noncompetitive substrates. Competitive substrates include carbon sources (e.g. CO2 ) which can be utilized by other microbial groups (e.g. sulfate-reducing bacteria). In contrast, a noncompetitive substrate is of exclusive metabolic value for only one microbial group. For methanogens these include methanol, methylated amines as mono-, di-, and trimethylamines [66], and certain organic sulfur compounds, e.g. dimethylsulfide [67]. Methanogenesis utilizing CO2 is termed ‘the carbonate reduction pathway’ CO2 + 8H+ + 8e− → CH4 + 2H2 O The net reaction of the acetoclastic pathway is CH3 COO− + H2 O → CH4 + HCO−3 An example for methanogenesis using a noncompetitive, methylated substrate (CH3 −A) is represented by CH3 −A + H2 O → CH4 + CO2 + A−H [68] The pathways for methanogenesis listed above involve the activity of archaea in anoxic conditions. However, occasionally oversaturation of methane (with respect to the atmospheric equilibrium) is also observed within the oceanic mixed surface layer, which cannot be explained by sedimentary methane production. This phenomenon is referred to as the ‘oceanic methane paradox’ [69]. Contrary to the current understanding, some methanogens are tolerant of increased levels of oxygen [70], allowing for the production of methane in oxygendepleted microniches of organic aggregates in an oxic environment [71]. Under such conditions dimethylsulfoniopropionate (DMSP) may serve as a substrate for methanogenesis [72]. In addition, aerobic bacteria have been identified that have the potential for methanogenesis under oxic conditions as a by-product during methylphosphonate (MPn) decomposition, which is used as a phosphate source [73, 74]. Recent investigations also revealed that the marine unicellular algae Emiliania huxleyi, a globally occurring coccolithophorid, is capable of producing methane under oxic conditions when using methionine [75]. Consequently, the combination of these processes may partly explain the observed paradox. The thermogenic and biogenic formation pathways (heat vs. enzymatic activity) result in varying stable isotope signatures, due to kinetic isotope fractionation, with regard to carbon and hydrogen (󳶳 Fig. 1.2) [68]. This allows for the discrimination between the two sources. In addition, thermogenic methane is usually accompanied

1.4 Cold seeps and gas hydrate |

7

Bacterial Carbonate Reduction Bacterial Bacterial Methyl-type Fermentation

mix & transition

Atmospheric

early mature

Geo

Artifical, Bit Metamorphic

Thermogenic

ther

mal,

associated

Hyd

roth

erm

humic

al, C

rysta

lline

Abiogenic ? Mantle ?

Fig. 1.2: C-, H-isotope signatures of microbial and thermogenic CH4 sources (modified from [68])

by higher hydrocarbons, which may be present as traces to concentrations of up to 40% [76]. Consequently, the ratio of methane to other hydrocarbons can also be employed to identify its origin of genesis. In comparison, thermogenic methane is generally characterized by less negative 13 δ C values, ranging roughly between − 20‰ and − 50‰ (VPDB). The corresponding hydrogen isotope ratios range approximately from δD − 100‰ to − 275‰ (SMOW). Biogenic CH4 spans a considerable range of δ13 C and δD ratios. The carbon isotopic signature is generally more negative than the thermogenic one, starting from around − 50‰ down to − 110‰ (VPDB). The δD values range from − 150‰ to − 400‰ (SMOW), showing substantial overlap between thermogenic and biogenic CH4 . The large range of encountered δ13 C ratios is due to the variable C isotope composition of the used carbon sources. In 󳶳 Fig. 1.2 two main fractions in biogenic δ13 C can be observed: the methyl-type fermentation and the carbonate reduction. The first pathway typically results in δ13 C ratios between − 50‰ and − 70‰ (VPDB). The carbonate reduction pathway generates greater depletion in 13 C and also shows less negative δD ratios.

1.4 Cold seeps and gas hydrate The leakage of CH4 at cold seeps is frequently associated with the occurrence of shallow gas hydrates, also called gas clathrates. Gas hydrates are crystalline cages composed of water, encasing a gas molecule. In this chapter, we will focus on methane

8 | 1 Methane seeps in a changing climate

Fig. 1.3: Schematic diagram of the marine gas hydrate stability zone and phase boundary for gas hydrates. The deep phase boundary is governed by the geothermal gradient. (Modified from [83])

hydrates and henceforth refer to them as gas hydrates. The formation of gas hydrates is governed by a function of temperature and pressure, exemplified in 󳶳 Fig. 1.3. Depending on these two factors, methane bearing gas hydrates may form in the marine environment at water depths between 300 to 900 m [77]. The vertical zone in which the physical conditions allow for the formation of gas hydrates is termed the gas hydrate stability zone (GHSZ). The phase boundary conditions for the stability of gas hydrates demonstrate that water depth (pressure) and temperature are inversely correlated. This is highly relevant for the vertical distribution of natural gas hydrate deposits. The different latitudinal regions of the Earth exhibit considerable variation regarding the extension of the GHSZ. Due to solar warming temperate and tropical oceans are characterized by a relatively warm surface layer, which either persists during the entire year (tropics) or develops during the summer (temperate). Therefore, the water body is stratified, resulting in reduced vertical mixing. In tropical regions, the layer of warm water might reach down to water depths of around 500 m, below which the temperature rapidly declines (thermocline). In arctic regions, a thermocline is often not observed and intensified mixing results in similar cold water temperatures from the surface to the bottom. Consequently, in an arctic setting, slight warming of the ocean surface can shift the phase boundary considerably to shallower water depths (< 300 m), in contrast to areas located at lower latitudes, where the presence of a thermocline greatly reduces vertical heat transport and thus expansion of the GHSZ to shallow water depths. Under natural conditions, gas hydrates may form in shallow sediments and at the sediment-water interface. Within the GHSZ methane may escape from the sediment as

1.5 Gas hydrates, methane bubbles | 9

gas hydrate-coated bubbles or gas hydrate flakes [78–80], which disintegrate rapidly during their ascent due to dissolution of methane into the undersaturated water column and when exiting the GHSZ. Considering the preconditions necessary for gas hydrate formation (low T, high P and CH4 ), all continental slopes may host an extensive fringe of gas hydrate deposits [81]. Estimates of the global CH4 inventory stored in hydrates differ substantially, ranging from 4.08 to 107 × 106 Gt C. However, recent modeling results suggest a total mass of 1,146 Gt C [82].

1.5 Gas hydrates, methane bubbles A useful tool in the quest for cold seep identification is the mapping of methane bubbles, which can be visualized as so-called flares with hydroacoustic techniques [84, 85] (󳶳 Fig. 1.4). Methane flares can be observed at cold seep locations inside, at the upper boundary of, and also outside the GHSZ [86]. In the first scenario, the sediment within the GHSZ may become saturated with methane and gas hydrate, caus-

(a)

(b)

(c) Fig. 1.4: Example of observed methane flares during an acoustic survey offshore of western Svalbard [89]

10 | 1 Methane seeps in a changing climate

ing the formation of methane bubbles, which eventually leave the sediment as flares. Methane leakage is also found at the upper boundary of the GHSZ. As a consequence of decreasing hydrostatic pressure and/or increasing sediment temperature, the GHSZ may retract, causing the destabilization of present gas hydrates and the formation of gaseous methane at the upper phase boundary. Depending on seasonal variations of bottom water temperature and underlying sediment, the upper boundary of the GHSZ may undergo appreciable temporal vertical shifts, causing periodic destabilization of gas hydrates [87]. In productive shallow coastal regions way outside the GHSZ, methane emission can also be frequently observed. Here, methane forms as a result of intensified degradation of buried organic matter [88]. In addition, CH4 originating from deeper reservoirs may migrate upwards, as in areas of mud volcanism or leakage through tectonic faults and fissures. If pore water saturation is reached, methane ebullition into the overlying water body takes place. As the occurrence of gaseous methane emanating from the seabed may be indicative of global warming, numerous large-scale research activities are dedicated to identifying the reasons for methane flare formation at continental slopes, especially at the vicinity of the upper boundary of the GHSZ. Pressing questions include to what extent might gas hydrate deposits destabilize and what impact will this have on the global climate?

1.6 Anaerobic methane oxidation, sulfide oxidation, and large seep fauna Although spatially large, the global ocean contributes relatively little to the atmospheric methane budget. This is the result of specialized aerobic and anaerobic microbes in the water column and in the sediment, respectively, consuming the majority of available methane within the marine realm. Similar to hot vents, cold seep locations also host habitats with a highly specialized microbial and faunal community [90]. The backbone of this thriving community is a group of chemosynthetic microorganisms, which typically utilize reduced molecules in the geofluids such as methane and sulfide. Research from the 1970s and 1980s already indicated that within marine sediments, CH4 is oxidized with sulfate as the terminal electron acceptor [91–93]. More than two decades later the process of anaerobic oxidation of methane (AOM) was revealed [18], showing that a consortium of archaea and sulfate-reducing bacteria carry out effective methane consumption in a distinct zone within marine sediments. To date, the exact mechanisms of this process are still a matter of debate. The archaeal and bacterial partners appear to form a syntrophic relationship, carrying out AOM in a combined effort. However, the detailed interactions between the partners, such as which intermediates are exchanged, remain poorly constrained [18, 94, 95]. Methano-

1.6 Anaerobic methane oxidation, sulfide oxidation, and large seep fauna |

11

Fig. 1.5: Typical cold seep fauna with endosymbiotic sulfide-oxidizing bacteria. The presence of these key organisms indicates increasing methane advection, resulting in respective increasing sulfide production through AOM, from left to right of the scheme (modified from [20])

trophy yields comparatively small amounts of energy, resulting in very slow growth and cell doubling rates on the order of several months [96], which severely hinders laboratory culturing efforts. A recent study suggests that AOM may also be carried out by the archaea alone and that sulfate-reducing bacteria are rather commensals, disproportionating a produced sulfur compound [94]. Irrespective of the actual pathway, AOM is carried out according to the following net reaction: − − CH4 + SO2− 4 → HCO3 + HS + H2 O − The reduction of the sulfate ion (SO2− 4 ) causes the genesis of sulfide (HS ). Together with the non-oxidized methane this sulfide is transported towards the sediment surface by diffusion and ascending fluids. At the sediment-water interface, where oxygen or nitrate is present, it is reoxidized by sulfide-oxidizing bacteria [97]. In addition to free-living microbes, a typical chemosymbiotic metazoan community characterizes cold seeps [20], which forms a symbiotic relationship with endosymbiotic sulfide-oxidizing bacteria (󳶳 Fig. 1.5). This community includes bivalves (Solemyidae, Lucinidae, Vesicomyidae, Thyasiridae, and Mytilidae) and also tube worms (Siboglinidae) [98]. Typically, one body part of the metazoan organism reaches down into the sediment to grant access to sulfide. The other part acquires oxygen. Alternatively, some bivalves take up sulfide liberated from the sediments directly from the surrounding seawater together with oxygen through their gills. Both sulfide and oxygen are transported to the body parts hosting the symbiotic bacteria, granting the controlled oxidation of sulfide. Some metazoans host symbiotic bacteria, which carry out methane oxidation under oxic conditions [99]. These symbiotic bacteria use methane as an electron donor and as a carbon source [100]. As this trophic relationship relies on the co-occurrence of oxygen and sufficient quantities of dissolved methane, metazoans with methanotrophic symbionts are restricted to areas of cold seeps and hot vents.

12 | 1 Methane seeps in a changing climate

In addition to chemosynthetic organisms, numerous faunal species inhabit cold seeps, including scavenging or predatory crabs (Lithodidae, Galatheidae, and Chirostylidae), which further benefit from chemosynthetic biomass [101]. The strength of methane flux and resulting sulfide production can be roughly distinguished macroscopically, as the seep fauna shows variability in composition according to the sulfide flux (󳶳 Fig. 1.5). Habitats characterized by high methane flux, and therefore also sulfide flux, are indicated by dense mats of sulfide-oxidizing bacteria at the sediment surface [20, 102]. A medium flux of sulfide is often encountered in habitats dominated by epibenthic mussels (e.g. Bathymodiolus sp., Calyptogena sp.). Vicinities of rather low sulfide flux are frequently inhabited by the bivalve genus Acharax, which lives inside the sediment, and tubeworms.

1.7 Carbonate depositions at cold seeps Another key feature encountered at numerous cold seeps is the presence of massive carbonate deposits [15, 103, 104]. Due to the chemical reaction of AOM, the pore water of cold seep locations is often enriched in bicarbonate (HCO−3 ), which is in equilibrium with carbonate ions (CO2− 3 ). Consequently, prolonged AOM ultimately increases the concentration of carbonate. When the buildup reaches supersaturation, carbonate may precipitate with calcium (Ca2+ ) or magnesium (Mg2+ ), forming Ca/Mg-carbonate minerals (󳶳 Fig. 1.6). These mineral deposits are termed authigenic carbonates, as they are formed autochthonously, i.e. in the seafloor where they are observed. Cold seep carbonates represent a unique geological archive, allowing the history of fluid flow and gas hydrate stability on continental margins to be constrained by the study of trace element ratios and isotope signatures [87, 105, 106]. In addition, cold seep associated carbonate deposits represent a long-term sink for carbon originating from advected methane, which would otherwise ultimately end up as carbon dioxide (CO2 ). Surficial authigenic carbonates add to habitat diversity by creating irregularly textured hard substrates in an otherwise sediment dominated environment. The increased habitat diversity facilitates an increase in metazoan species abundance, span-

Fig. 1.6: Cold seep calcium carbonate deposits at the Hikurangi Margin, New Zealand, at 1,098 m water depth with tube worms (modified from [108])

1.9 Methane and past climate

| 13

ning from organisms directly dependent on the availability of seep-specific chemical components to predators of higher trophic levels [107].

1.8 Effects of climate change on methane seeps and vice versa The stability of gas hydrate deposits over long time scales is of global relevance due to the effects of atmospheric methane on climate change. The stability of gas hydrates is essentially governed by the prevailing hydrostatic pressure and temperature (󳶳 Fig. 1.3). According to the United States Global Change Research Program National Climate Assessment in 2014 [109], an average sea level rise of between 300 and 1,200 mm is to be expected by the end of the century. Consequently, decreasing hydrostatic pressure is generally not an issue of concern in the near future. However, increasing global average temperatures, which we are currently experiencing, threaten to destabilize gas hydrates and subsequently liberate CH4 . Once in the atmosphere, CH4 acts as a potent greenhouse gas, being > 25 times more efficient at trapping heat than CO2 [110, 111]. The residence time of methane in the atmosphere is about a decade, after which it is oxidized to the longer-lived CO2 [112]. Consequently, global warming and gas hydrate destabilization can create, at least in principle, a positive feedback loop, amplifying the greenhouse effect and global warming.

1.9 Methane and past climate Research suggests that large-scale marine gas hydrate destabilization might have contributed to dramatic changes in the global climate in Earth’s history, such as during the Permian–Triassic extinction event 250 million years ago, when approximately 95% of marine invertebrate species were wiped out [113]; or during the Paleocene–Eocene Thermal Maximum (PETM) 56 million years ago [114]. During both events, the prolonged release of methane from destabilized gas hydrates over several thousands of years (for the PETM approximately 4,000 years) might have resulted in global temperature increases of ∼5 °C. In 2003 Kennett et al. put forward the so-called clathrate gun hypothesis [115], stating that a temperature increase of seawater and/or a drop in sea level could cause an abrupt release of large quantities of methane from destabilized gas hydrates in marine sediments and subseafloor permafrost. The released methane would lead to a rapid increase of atmospheric and oceanic temperatures, causing the liberation of more gas hydrates and markedly amplifying the greenhouse effect within a human life span, like a bullet from a gun that, once fired, cannot be stopped. The clathrate gun hypothesis was proposed specifically as a potential explanation for warming events during and after the last glacial maximum 13,000 to 10,000 years ago. However, subsequent investigations strongly reject this hypothesis, be-

14 | 1 Methane seeps in a changing climate cause δ13 CH4 analyses from ice core records show conclusive evidence for elevated methane production by tropical wetlands and/or terrestrial plants under oxic conditions [116–118].

1.10 Cold seeps and recent global climate change Ongoing global warming is a generally accepted climate phenomenon caused via a direct consequence of anthropogenic influences [119]. According to the Intergovernmental Panel on Climate Change (IPCC) [120], a temperature increase of 1–4 °C is expected within the next 100 years if greenhouse gas emissions remain at current levels. In addition to the atmosphere, the global ocean is also subject to warming. From 1971 to 2010, the average temperature of the surface ocean (0–75 m water depth) increased by 0.11 °C per decade, while the deeper layers below 700 m water depth showed an increase of 0.015 °C per decade over the same period [121]. Consequently, the potential climatic impact of gas hydrate dissociation gives rise to the following pressing questions: – How much of the atmospheric heat energy is actually reaching the gas hydrate bearing sediments of the continental slopes? – Is the heat energy sufficient to destabilize present gas hydrates? – Over what time scale could the gas hydrates destabilize? – How much of the liberated CH4 will actually reach the atmosphere?

1.11 Heat transfer from the atmosphere to marine sediment Considering global warming and consequential heat transfer between Earth’s different realms, we need to familiarize ourselves with the physical term ‘heat capacity’ or ‘thermal capacity’. Heat capacity is the ratio of heat added to an object and its resulting temperature change. The higher the heat capacity of a substance, the more energy per unit mass is required to heat it up. Water has a comparatively high heat capacity, illustrated by a rather high boiling point, as a consequence of the numerous hydrogen bonds between individual water molecules. Before boiling can occur in a water body these intermolecular hydrogen bonds have to be broken, which requires energy, allowing the gas molecules to enter the gaseous phase. Due to global warming, gas hydrate destabilization requires heat transfer from the atmosphere via the global ocean to the seafloor. As a consequence of its physical properties, water has a considerably higher heat capacity (by a factor of 4–5) than the atmosphere or the terrestrial environment. This difference in heat capacity plays a crucial role in regulating the global climate [122], especially as the marine realm is the Earth’s largest water, heat, and carbon reservoir [123]. Because of its physicochemical properties, seawater takes up heat approximately 4,000 times faster than the same

1.11 Heat transfer from the atmosphere to marine sediment | 15

Fig. 1.7: Energy accumulation in ZJ (1 ZJ = 1,021 J) by different parts of the Earth’s climate system relative to 1971 for the period between 1971 and 2010. Dashed lines denote 90% confidence intervals for all variables combined (figure from [125])

volume of air, resulting in the transport and storage of large quantities of heat. According to recent investigations, [124] the oceans have taken up more than 93% of the heat created by anthropogenic warming since 1971. Most of this heat energy is stored in the upper (< 1,500 m) ocean (󳶳 Fig. 1.7). In contrast to the turbulent heat transport in oceanic water masses, the much slower process of thermal conduction largely governs heat propagation within sediments. Calculating the impact of sustained temperature change, heat is transported to a maximum sediment depth of 18 m and 178 m after 10 and 1,000 years respectively, regardless of the degree of temperature change [126]. As a consequence, even prolonged (millennia-scale) heat transport may affect only gas hydrates situated in close proximity to the seafloor. In the case where the temperature increase in the sediment remains below the threshold of gas hydrate destabilization, no CH4 is released. In contrast, if the temperature increase exceeds the threshold of the stability field, a vertical retreat of the GHSZ is the consequence (󳶳 Fig. 1.3). Gas hydrate deposits now situated

16 | 1 Methane seeps in a changing climate

outside (on top of) the GHSZ will start to destabilize and liberate CH4 . Considering the gas hydrate distribution at continental margins, less than 5% of the deposits are prone to destabilization during the next 1,000 years [126].

1.12 Gas hydrate destabilization in the near future Numerical climate modeling, including projections of temperature increase, form the backbone for estimates of gas hydrate dissociation rates and methane liberation. Current modeling approaches suggest an atmospheric temperature increase of approximately 0.2 °C per decade (IPCC) [141]. Given these conditions, between 0.6 and 0.03% of today’s gas hydrate inventory is likely to destabilize within the next century [82]. Not all gas hydrates face the same risk of destabilization; their susceptibility is largely governed by the geographical region and the position on the continental slope [127] (󳶳 Fig. 1.8). Generally speaking, the higher up the slope a gas hydrate deposit is situated, the higher the probability of destabilization and consequential

(b)

(a)

(c)

Fig. 1.8: Predicted change in the distribution of gas hydrates (kg C/m2 ) for (a) the Arctic and the Norwegian Continental Margin, (b) the Cascadia Margin, Gulf of Mexico, Blake Plateau, and Gulf Stream, and (c) the Sea of Okhotsk (figure from [82])

1.12 Gas hydrate destabilization in the near future | 17

Fig. 1.9: Schematic overview of gas hydrate locations across the continental margin [128]

methane release. Gas hydrates can be encountered in five different regimes along the slope, including the terrestrial and marine realm (󳶳 Fig. 1.9). The combined gas hydrate deposits in onshore and arctic subsea permafrost account for less than 1.5% of the global inventory [126]. Numerical modeling of the susceptibility to prolonged temperature increase reveals that time scales of > 1,000 years are required to destabilize the topmost part of permafrost associated gas hydrates while the vast majority remain unaffected. Following the continental margin, gas hydrate formation starts once the stability field has been reached, which corresponds to combined water and sediment depths of 300 to 500 m [126], depending on prevailing water temperature. In these parts of the continental slope, the susceptibility of gas hydrates to warming is high [129] as the entire GHSZ (approx. 40 m thickness) is situated in close proximity to the surface. According to recent estimates [126], < 3.5% of global gas hydrate deposits may occur in this setting. Here, prolonged warming may result in the destabilization of the entire gas hydrate reservoir within the next 100 years [126]. Gas hydrates located at water depths exceeding 500 m and 300 m in temperate and high latitudes, respectively, represent about 99% of the global inventory [130]. The probability of destabilization in this setting due to bottom water warming is low within the next 100 to 1,000 years as the GHSZ protrudes considerably into the water column above the sediment [129]. Under these conditions, large volumes of water need to be warmed in order to push the GHSZ towards the sediment surface, inducing gas hydrate destabilization.

18 | 1 Methane seeps in a changing climate

1.13 Areas most affected by gas hydrate destabilization 1.13.1 The Arctic A large proportion of the global gas hydrate inventory is stored in the continental margins of the Arctic Ocean (󳶳 Fig. 1.10). In this setting, hydrates occur primarily close to the sediment surface at shallow water depths, stabilized by prevailing cold temperatures throughout the year [131]. Consequently, it can be argued that warming of shallow water masses in the Arctic in particular and consequential heat transfer to the sediment may result in the destabilization of substantial amounts of gas hydrates and catastrophic methane liberation. In 2009, more than 250 methane flares at water depths at and above the upper boundary of the GHSZ were reported from the continental margin off the coast of West Spitsbergen [85] giving rise to the hypothesis that gas hydrate destabilization at this location is taking place as a result of recent warming of the West Spitsbergen current. Follow up investigations identified seasonal temperature fluctuations as the main driver for episodic gas hydrate destabilization in this area [87]. Extensive methane venting to the atmosphere from sediments of the East Siberian Arctic Shelf were reported in 2010 [132]. To date, it is unclear whether this process has been going on steadily for an extended period of time or if it is the consequence of recent temperature change, with increasing methane liberation from sediments to be expected in the future [133]. Although there is broad consensus among the scientific community regarding the high susceptibility of gas hydrates in shallow Arctic Ocean sediments, the fate of these methane deposits under climate warming conditions cannot be constrained easily. Recent investigations clearly demonstrate that the warming of Arctic waters

Fig. 1.10: Global map of gas hydrate deposits in the global regions [138]

1.14 The fate of methane |

19

and sediments will simultaneously induce several physical and biological processes, which may have contrary effects on the sea-air methane flux [89]. To constrain these processes in the future, a multidisciplinary research effort focusing on the microbial methane oxidation potential [134, 135], methane transport [136], and understanding of ice cover [137] is needed. As a matter of fact, surface air temperatures in the Arctic show warming rates approximately twice as high as the global average during the last four decades [139], and according to recent climate modeling approaches [140], warming is expected accelerate over the next 100 years in the case of continued greenhouse gas emissions [141]. However, the susceptibility of gas hydrates in the Arctic Ocean to increasing water temperature varies strongly, depending on geographic region and position on the continental slopes. Primarily shallow areas, which are subjected to Atlantic inflow, are under high risk of destabilization. Within the next 100 years, approximately 25% of shallow and middepth hydrate bearing sediments may be affected by warming [142], causing the destabilization of 1.4–3.5% of the arctic gas hydrates [82, 126].

1.13.2 Blake Plateau This site is located in the western Atlantic off the coast of South Carolina and Georgia and belongs to the Blake Ridge formation (󳶳 Fig. 1.8). At the plateau, the water depth declines to about 500 m. This location has long been recognized as a major gas hydrate province within the US Exclusive Economic Zone [143]. The amount of deposit stored at this site is comparatively small, representing < 1 ‰ (approx. 1 Gt C) of the global hydrate inventory. However, according to recent modeling approaches, continuing ocean warming will lead to a decrease of up to 12% of the Blake Plateau gas hydrate reservoir in the next century [82], representing the highest rate of gas hydrate destabilization known so far.

1.14 The fate of methane In the case of gas hydrate destabilization, the amount of liberated CH4 reaching the atmosphere depends on a combination of physical and biological factors. In the sediment, the release of CH4 may result in the saturation of pore water with dissolved gas. Under diffusion dominated conditions, a considerable fraction of the CH4 may be retained in the sediment over long time periods [144]. If the concentration of liberated CH4 exceeds the maximum solubility of the pore water, gas bubbles may form. Depending on the seafloor conditions, these bubbles may be either trapped, forming long-term gas pockets, or move upward towards the sediment surface.

20 | 1 Methane seeps in a changing climate

1.14.1 Methane in the sediment During its ascent in the sediment, a large fraction of dissolved CH4 may be consumed by microbial methanotrophs during the AOM process. The efficiency of this ‘microbial methane filter’ in marine sediments is inversely correlated with the CH4 flux rate. Under low flux conditions the efficiency of this filter may reach > 95% [11, 145, 146]. By contrast, at locations with high flux, a considerable amount of CH4 can escape into the water column, mainly as methane bubbles [52, 147, 148], but will the benthic microbial filter also maintain its efficiency in the case of sudden gas hydrate destabilization and increased advective methane flux? According to bioenergetics reactiontransport simulations, the sedimentary AOM community will show a lag phase on the order of 60 years before the methane efflux will reach original levels, due to the slow growth rate of the involved microorganisms [149]. However, recent studies showed that a thriving AOM community may also develop within less than 20 years in the case of prolonged high advective methane flux [150]. Although these timespans appear to be rather long from a human perspective, this time window may be far too short from a geological point of view to allow for the escape of relevant amounts of methane from the sediment into the water column.

1.14.2 Methane in the water column Depending on the flux, variable amounts of methane may enter the water column above cold seep sites. Once dissolved, methane is readily available to be consumed by bacteria carrying out aerobic methane oxidation (MOx) according to the following reaction: CH4 + 2O2 → CO2 + 2H2 O MOx bacteria represent a group that uses methane as its sole carbon source. Members of the group are ubiquitously present in various environments, including the global ocean [151]. In contrast to the sedimentary AOM consortium, MOx bacteria produce carbon dioxide (CO2 ) as an end product of methane oxidation. In water, CO2 reacts to carbonic acid, which immediately dissociates into bicarbonate and carbonate ions, thereby releasing protons, which cause a decrease in pH. The increased uptake of CO2 from the atmosphere is currently the focus of intensive research activities, as this factor is responsible for the observed ocean acidification, which has severe implications for marine calcium carbonate producing plants and animals [152]. An increased liberation of methane into the water column and subsequent activity of MOx bacteria may have the potential to aggravate this effect. However, CO2 production due to organic matter respiration of oxygen in the water column is by far greater than from MOx activity. Consequently, increased MOx activity is unlikely to contribute significantly to decreasing seawater pH in the near future.

1.15 Methane seeps – relevant for climate change? | 21

Only a small fraction of water column CH4 is actually reaching the ocean surface. During the rise towards the ocean surface, the CH4 in the gas bubbles exchanges with dissolved oxygen and nitrogen from the surrounding seawater. If a bubble is emitted at water depths exceeding approximately 100 m, these two other gases replace nearly all of the CH4 when the bubble reaches the water surface [153]. Considering that cold seeps usually occur at water depths exceeding 100 m, it can be assumed that the vast majority of emitted methane dissolves into the water column. A large fraction of the dissolved CH4 is rapidly consumed by aerobic (oxygen-dependent) methanotrophic bacteria [154, 155], producing CO2 and water. Due to the high efficiency of the combined microbial methane filters in the sediment and the water column, almost 90% of the produced methane is consumed in the ocean before it can reach the atmosphere [11]. However, the hydrographic regime controls the capacity of the water column methane filter. While rapid water mass exchange leads to translocation of water column methanotrophs away from the methane point source and thus lowers this capacity substantially [134], stratification provides more stable conditions for the methane-consuming microbes [135, 156].

1.15 Methane seeps – relevant for climate change? The concentration of atmospheric methane is the result of global emission (sources) and processes of consumption (sinks) (󳶳 Fig. 1.11). The total annual flux of CH4 , combined from all sources, amounts to approximately 555 Tg yr−1 [2]. The major sink for methane is the atmosphere itself. In the stratosphere and the troposphere, spanning

Fig. 1.11: Contribution of relevant terrestrial and atmospheric sources and sinks for the average global methane budget from 2000–2009 in Tg CH4 yr−1 . Modified from Global Carbon Project 2013; figure based on Kirschke et al. (2013) [2]

22 | 1 Methane seeps in a changing climate

the first 50 km above the Earth’s surface, CH4 is partially oxidized photochemically to CO2 and water vapor, which are also important greenhouse gases [157]. In addition, CH4 reacts with natural chlorine gas forming chloromethane and hydrochloric acid during free radical halogenation. In the case of CH4 liberation from cold seeps, which are often fueled by gas hydrates, we still face severe knowledge gaps, which preclude a fully justified estimate of whether cold seeps and gas hydrates play an important role in the future methane budget of the atmosphere. The magnitudes of heat transfer from the atmosphere to the ocean’s deeper water layers, which ultimately determines potential gas hydrate destabilization, have been debated recently [158]. Furthermore, our knowledge about the efficiency of the microbial water column methane filter, the last sink for methane before entering the atmosphere, is not well constrained [11, 134, 159, 160]. However, at present levels, the combined sediment and water column sinks reduce the amount of oceanic CH4 efficiently. The future efficiency of the microbial methane filter will depend on the release rates of hydrate bound methane, which are probably rather slow considering current global warming scenarios [128].

1.16 Atmospheric methane concentration over time The concentrations of CH4 and CO2 have undergone considerable variations during the last 800,000 years [128], showing cyclic glacial-interglacial variations (󳶳 Fig. 1.12). During the last seven decades, atmospheric concentrations of both greenhouse gases continuously increased, resulting in methane concentrations approximately 150% higher than in preindustrial times. This drastic increase in atmospheric methane fuels the discussion over a potential positive feedback to climate warming.

Fig. 1.12: Methane and CO2 concentrations in the atmosphere over the last 800,000 years (from the European Project for Ice Coring in Antarctica (EPICA) [128]). LMG – Last Glacial Maximum

1.17 Closing remarks

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23

Today’s atmospheric methane originates from a variety of sources, which can be divided into natural and anthropogenic ones. The latter include the production of rice, ruminant farming, biomass and fossil fuel burning, as well as landfills and waste depositions. Relevant natural methane emission sources comprise geological sources such as volcanoes, naturally occurring biomass burning, termites, and microbial production in freshwater systems and wetlands. Under the current climate conditions, the destabilization of gas hydrates contributes about 0.9% to the global methane sources, while combined natural methane emissions from wetlands, anthropogenic fossil fuel burning and cattle farming account for approximately 60% of methane emissions.

1.17 Closing remarks Worldwide destabilization of deep ocean gas hydrate reservoirs leading to a massive liberation of methane in the near future seems unlikely under the current rates of global warming. Furthermore, substantial transport from the seafloor to the atmosphere by methane bubbles is also unlikely as the vast majority of natural seeps are located at water depths > 100 m, resulting in the complete dissolution of methane in the water column. In the case of liberated methane, the microbial filter in the water column will largely reduce the amount of methane reaching the atmosphere. Considering the projected temperature increase for the near future, the contribution of methane from present and future cold seepage can be assumed to be rather low. Nevertheless, this estimate has to be regarded with caution because the rate of warming and heat transfer to deeper water layers as well as the efficiencies of the microbial methane filter in sediments and the water column are not well constrained. Systems that are primarily susceptible to gas hydrate destabilization include permafrost and shallow regions of the Arctic Ocean. On a global scale, a maximum of about 0.6% of the gas hydrate inventory is prone to dissociate, facilitating the liberation of CH4 . Under the current climatic situation, only 51% of the atmospheric methane originates from natural sources, while the remaining 49% are manmade, most importantly related to agriculture and waste, fossil fuels and biomass burning [2]. Consequently, efforts to diminish the anthropogenic production of methane could effectively mitigate the accelerating influence of methane in global warming.

24 | 1 Methane seeps in a changing climate

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Samantha B. Joye and Sara Kleindienst

2 Hydrocarbon seep ecosystems 2.1 Overview Hydrocarbon seep ecosystem are spectacular deep sea oases where hydrocarbon flux from deep reservoirs through sediments and then into the water column fuels diverse biological assemblages that mediate a variety of biogeochemical processes. A number of different types of cold seeps exist, including oil and/or gas seeps, brine seeps, brine pools, and mud volcanoes. Several geological processes contribute to the formation of hydrocarbon seeps but these seeps, located across the globe, support diverse, yet similar, biogeochemical processes and microbial and animal fauna. In addition to serving as a sink for petroleum and low molecular weight gases, like methane, seeps support tremendous biodiversity, both endemic fauna and fauna which utilize these habitats to support their own populations. The microbiology of hydrocarbon seep ecosystems is an area that has received extensive attention because it is these microbial processes that form the geobiological engine that drives and sustains these fascinating ecosystems.

2.2 Introduction – discovery of hydrocarbon seeps One of the most important oceanographic achievements of the 20th century was the discovery of hydrothermal vents on the seafloor in the late 1970s [1–5]. Release of super-heated fluids from the seabed fuels geochemical reactions that generate sulfide minerals towers reaching meters above the seafloor and the released geochemical energy supports stunningly diverse and novel biological ecosystems. Many ventassociated animals were new to science and lived in symbiotic association with microorganisms [6] that take advantage of the abundant geochemical energy released from the seabed [7] in the form of hydrogen gas, hydrogen sulfide, reduced metals, and methane. About a decade later, similar biological oases along the seabed were discovered in the Gulf of Mexico, first along the West Florida Escarpment [8] and then along the continental slope of the Northern Gulf [9] and then off the West coast of Oregon (U.S.A.; [10]). Similar to the processes occurring at hydrothermal vents, these hydrocarbon seeps, often referred to as “cold seeps”, are formed when fluxes of fluids from the deep subsurface that flow through conduits within the seabed before being expelled into the overlying water column. This material flux represents a recycling process that regenerates and returns materials from ancient, isolated geologic reservoirs

https://doi.org/10.1515/9783110493672-002

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to actively cycling biogeochemical domains and occurs on time scales of tens of thousands to tens of millions of years. In contrast to hydrothermal vents, cold seeps fluids are not super-heated. The discharging fluids are only slightly warmer (by a few °C) than bottom water (note: mud volcanoes are an exception and can discharge fluids that are 10’s of °C warmer than the overlying bottom water, [11]). The circumneutral pH (∼7) of the fluids discharged at hydrocarbon seeps carry oil, gas (mainly methane) and/or brine instead of the acidic mixture of reduced gases (e.g. hydrogen sulfide, carbon dioxide) and metals (e.g. Fe(II), Mn(II), etc.) discharged at hydrothermal vents. The animal communities at hydrocarbon seeps are also different from those observed at hydrothermal vents. Though some species are related (e.g. Riftia sp. tube worms at vents vs. Lamellibrachia sp. at seeps) phylogenetically and physiologically, e.g. the ability to oxidize reduced sulfur species, the mechanisms of activity were different. Following the discovery of hydrocarbon seeps along the West coast of the US and in the Gulf of Mexico, similar ecosystems were discovered across the globe [12], on both active [13–15] and passive [16–18] margins and in semi-enclosed ocean

Fig. 2.1: Global distribution of gas hydrates and by inference, cold seeps. Predicted volume distribution of methane in hydrate expanded to STP in a 1° × 1° area (color scale), overlaid with points showing locations of recovered gas hydrate (light blue) and inferred gas hydrate (yellow) based on references compiled through the middle of December 2013 by the USGS Gas Hydrate Project. Volume distribution (color scale) reprinted and modified with permission from Klauda JB, Sandler SI. Global Distribution of Methane Hydrate in Ocean Sediment. Energy Fuels 2005, 19(2):459–470 (DOI: 10.1021/ef049798o). Copyright 2005 American Chemical Society. Locations of recovered gas hydrate and inferred gas hydrate from USGS Gas Hydrates Project, woodshole.er.usgs.gov/projectpages/hydrates/database.html, accessed April 7, 2017

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basins [19]. Hydrocarbon seeps pepper continental margins across the globe forming a nearshore network of deepwater chemosynthetic ecosystems similar to the distribution of hydrothermal vents offshore, which tracks plate boundaries across ocean basins (󳶳 Fig. 2.1). Like hydrothermal vents, each hydrocarbon seep ecosystem is somewhat unique due to variations in the physical (e.g. temperature, pH and salinity) and chemical (e.g. concentrations of hydrogen sulfide, ammonia, organic matter, methane, petroleum, silica, phosphorus, and metals) characteristics. These physical and chemical factors exert constraints on the nature and quality of habitat for seep-endemic animals and microorganisms and determine the metabolic energy available to fuel chemosynthesis by symbiotic as well as free-living microorganisms. Hydrocarbon seep ecosystems have captivated the imagination of scientists and the general public since their discovery. Much remains to be learned about these fascinating ecosystems. In this chapter, we overview the geology, biogeochemistry, and microbiology of these unique deepsea ecosystems.

2.3 Geology of hydrocarbon seeps 2.3.1 Types and formation mechanism Hydrocarbon seeps form via several mechanisms but their formation has a common requirement: availability of sufficient organic matter that is buried to a depth where thermogenic (or biogenic in the case of gases) production of petroleum and dissolved alkanes (e.g. methane, ethane, propane, etc.) ensues. The formation of oil and gas at depth generates reservoirs of hydrocarbons that are often isolated within distinct geological strata [20]. Fracture faults, generated by salt tectonics, created by gas overpressure, or driven by structural tectonics, create conduits through which these deeply sourced hydrocarbons can migrate upwards towards the seabed [16, 21, 22] (󳶳 Fig. 2.2). Hydrocarbon seeps can be characterized by the geochemical nature of seepage [23]: some seeps discharge mainly methane, often with small amounts of higher alkanes; some seeps discharge methane with petroleum (oil); some seeps discharge oil, methane and brine; and others discharge brine and gas. Roberts [23] grades hydrocarbon seeps along a spectrum of mud prone (young seeps, e.g. mud volcanoes) to mineral prone (old seeps, e.g. dominated by carbonates and hydrates); this spectrum reflects not only geological characteristics but the relative rate of fluid discharge. Mud volcanoes are high flow features that discharge copious amounts of fluidized mud, brine and hydrocarbons into the environment [24]. Older, mineral prone, systems typically support only modest rates of seepage. The nature of seepage influences the biological ecosystem associated with the seep and this can vary over time at a particular seep: as fluid flow rate and geochemistry evolve, so do the associated biological communities.

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Fig. 2.2: Types of hydrocarbon seeps. (a) Brine pool. Credit: Ocean Exploration Trust. (b) Mud volcano, Flower Garden Banks National Marine Sanctuary, Gulf of Mexico. Image courtesy of Emma Hickerson (NOAA) & Ocean Exploration Trust. (c) Oil Mountain, oil seeping from an oil-saturated hydrate mound in GC60, Gulf of Mexico. (d) Methane seep, upper slope, offshore Virginia. Image courtesy of NOAA Okeanos Explorer Program, 2013 Northeast U.S. Canyons Expedition. (e) Methane hydrate on the New England margin (1,400 meters). Methane bubbles visible in foreground. Image courtesy of NOAA Okeanos Explorer Program, 2013 Northeast U.S. Canyons Expedition. (f) Carbonate outcrop populated by tube worms and anemones next to a brine pool, Gulf of Mexico. Image courtesy of NOAA Okeanos Explorer Program, Gulf of Mexico 2014 Expedition

Along accretionary prisms at active margin, hydrocarbon seepage is dominated by methane of biogenic origin (e.g. Hydrate Ridge, [25] or SW Taiwan, [26]. Clay dewatering contributes to an enhanced fluid flux and this flux enriches surface sediments with methane and carries dissolved methane into the water column, generating plumes that can reach hundreds of meters above the seabed [27] (󳶳 Fig. 2.2). These active margin environments can support accumulation of substantial amounts of methane (gas)

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hydrate, which forms a bottom simulating reflector (BSR) that is detectable through acoustic surveys [27] and these hydrate deposits are sensitive to water column temperature regimes and thus global climate [28] [discussed further below]. In salt rich domains, such as the Gulf of Mexico or Congo Basin, brine fluids form at depth through clay dewatering processes and these brines transport oil and/or gas, along with other chemical species, as they migrate towards and are expelled from the seabed [24, 29, 30]. Methane in these environments can have either thermogenic or biogenenic isotopic signatures [22]. In basins such as the Gulf of Mexico, where an expansive salt body is present, the sediment overload creates spectacular basin and ridge topography along the seabed [30, 31]. Mud (brine) volcanoes, brine basins [32], brine lakes and brine streams/rivers are present across the system [24, 30] (󳶳 Fig. 2.2). Brine discharge can be explosive [11, 24] and mud volcanoes can discharge copious amounts of gas and oil (󳶳 Fig. 2.2b). In other scenarios, brine discharge is slower and more diffuse, leading to accumulation of pools of brine along the seabed or with in enclosed basins (󳶳 Fig. 2.2a) [29]. The variability in geochemistry of the expelled brine fluids varies markedly across the Gulf system [24, 28] yet fluid discharge rates remain poorly constrained, as does the microbiology and biogeochemistry of these brine habitats. Along passive margins, significant deposition of organic matter leads to generation and accumulation of methane and methane hydrate as well as oil (󳶳 Fig. 2.2c), in some places (e.g. Arctic) over geologic time (e.g. Atlantic margin, [18], and Arctic Ocean, [33]). Much more is known about hydrocarbon seep ecosystems along active margins and in petroleum rich basins like the Gulf of Mexico than is known about the hydrocarbon ecosystems along passive margin environments. However, seep fauna similar to those in the Gulf of Mexico, specifically rich communities of chemosynthetic Bathymoliolus sp. mussels, are known to exist in western Atlantic margin environments (e.g. Blake Plateau and elsewhere, [8, 34]) (󳶳 Fig. 2.2d). Gas hydrate deposits are pronounced in this area, just as in the Gulf of Mexico, but, in general, western Atlantic margin seeps are underexplored. Authigenic carbonates are a by-product of hydrocarbon seepage – a result of production of inorganic carbon and an increase in alkalinity as hydrocarbons are oxidized; this creates ideal conditions for precipitation of carbonate nodules, rocks, and pavements along the seabed. In the Arctic Ocean, numerous methane flares, called pingos, have been documented [35], but whether that flux of hydrocarbons across the seafloor supports chemosynthetic communities is not known.

2.3.2 Gas hydrates Gas hydrates are a common feature of hydrocarbon seeps and constitute an important component of the global carbon cycle [36]. The amount of methane present in gas hydrates likely exceeds conventional fossil fuel reserves of oil and gas [37]. Though the

38 | 2 Hydrocarbon seep ecosystems largest fraction of the gas hydrate reservoir lies beneath ∼200 m of sediment at the base of continental margins [38], the most dynamic hydrate deposits are located within a few tens of meters of the sediment-water interface in high flow, advective regions such as the Gulf of Mexico, where fluid flow driven by salt tectonics connects deep hydrocarbon reservoirs to the seafloor through advective transport processes. Gas hydrates can form in well-characterized temperature (T) and pressure (P) regimes (the Hydrate Stability Zone, or HSZ) in association with high fluxes of methane and other alkanes, carbon dioxide, hydrogen sulfide, nitrogen and oil [38–41]. In general, gas hydrates form spontaneously in continental margin sediments where a steady source of gas occurs at T < 10 °C and P of 1–5 MPa [37, 39]. Methane hydrate systems are inherently dynamic way-stations for light hydrocarbons and oil migrating up through the upper few hundred meters of continental margin sediments (󳶳 Fig. 2.2c and d, 󳶳 Fig. 2.3). Closely coupled fluid (oil and brine) and gas fluxes, driven by salt tectonics [42–45] and associated upward advective processes [46–48], feed gas into the HSZ at rates that vary as a function of the developmental stage of gas hydrate deposits. Solid masses of gas hydrate occur at the seafloor

Fig. 2.3: Seafloor expressions of gas vents and gas hydrate in the Gulf of Mexico. (a) an active oil/gas vent at GC600; (b) surface breaching gas hydrate at GC234; (c) oil wicking from the gas hydrate mound at GC234; (d) a massive surface breaching hydrate at GC600 with oil/gas chimneys (front below the scale bar), note the darker coloration which possibly reflects a higher oil content in the hydrate larger view of the mound shown in 󳶳 Fig. 2.2c; (e) an eroding gas hydrate inside a crater at GC232; (f) gas hydrate from GC234 recovered using a temperature-pressure certified recovery chamber; (g) massive carbonate pavement covered with sea fans and cold water corals

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interface [49, 50] where they form mounds and dissociate under the influence of physical, geological and biogeochemical processes at rates that are presently not well constrained [51, 52]. These mounds serve as a home for microorganisms, such as microbial mats, and megafauna, such as chemosymbiotic mussels. A frequently quoted estimate of the global methane hydrate resource is 20,000 trillion cubic meters, or about 700,000 trillion cubic feet [53]. The majority of hydrates are found in slope environments below 400 m water depth, and the Gulf of Mexico is thought to contain the largest reservoir of hydrate in the U.S. (607 trillion m3 of gas in place, [54]). However, the gas hydrates in the Gulf are found at shallow depths in comparison to other continental margin locales. Methane hydrate deposits are a significant potential future energy resource as well as globally significant and dynamic carbon reservoir.

2.3.3 Authigenic carbonates Hydrocarbon seep ecosystems have many fascinating and visually compelling geological features and authigenic carbonates certainly fall in that category (󳶳 Fig. 2.2f and 󳶳 Fig. 2.3g). Authigenic carbonates form spontaneously when the saturation state of calcium carbonate exceeds that required for spontaneous mineral precipitation. At the seafloor, carbonates precipitate as the result of dissolved inorganic carbon accumulation, which results from the oxidation of hydrocarbons, both oil and methane (or higher alkanes), by microbial processes [21, 55]. Authigenic carbonates are the primary source of hard substrate at hydrocarbon seeps and the amount of authigenic carbonates present at seeps can be extensive: at Builders Seep, a methane seep off New Zealand, extensive carbonate pavements cover over 70,000 km2 [56]. Such carbonates play an important role in the evolution of seepage regimes as carbonate pavements can occlude or divert locations of hydrocarbon discharge, driving hydrocarbon seep evolution, and they also generate topography and substrate for sessile animals, such as mussels, tubeworms and corals. The porous nature of carbonates makes them an ideal substrate for microorganisms as well [57, 58] though the role of carbonates as a habitat for microorganisms is poorly constrained. Interestingly, the mineralogy of authigenic carbonates in the modern ocean and at most hydrocarbon seeps is dominated by aragonite, due to thermodynamic and geochemical considerations [59]. However, at brine seeps where the ratio of calcium to magnesium deviates from seawater values, low magnesium calcite (i.e. < 5 mol% Mg) can dominate and in some cases account for a majority of carbonate present [59]. At brines in the Gulf of Mexico, magnesium concentrations are significantly depleted relative to seawater values [22] and the Mg/Ca molar ratio is as low as 0.6 (compared to ∼4.9 in seawater). Thus the mineralogy of these carbonate deposits can reveal much about the type of hydrocarbon seep from which they originated, long after active seepage has stopped.

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2.4 Biology and biogeochemistry of hydrocarbon seeps 2.4.1 Biogeochemistry Hydrocarbon seeps are areas of dynamic biogeochemical cycling of hydrocarbons and a multitude of other elements [60]. The discharge of hydrocarbons is spatially and temporally variable, and at most seeps, the rates of discharge are very poorly constrained, though recently new approaches provide a straightforward way to estimate and quantify discharge rates [61]. Given that discharge rates determine the availability of energy-rich substrates to fuel microbial processes and drive the geobiological engine of hydrocarbon seep ecosystems, this new approach to quantifying rates of hydrocarbon discharge rates is a very important development. Most efforts aimed at understanding elemental cycling at cold seeps has focused on the anaerobic oxidation of methane coupled to sulfate reduction [62], but oil oxidation is a critical process at hydrocarbon seeps as well. The paucity of knowledge regarding oil biodegradation rates at seeps stems from methodological difficulties: oil is a complex mixture of thousands of compounds making it extremely difficult to quantify turnover rates of this pool. Furthermore, even determining rates of oxidation of model compounds (i.e. hexadecane oxidation rates as an indicator for alkane turnover or naphthalene oxidation rates as a proxy for PAH oxidation rates) is problematic given the high background concentrations of oil in many hydrocarbon seep sediments [63]. Determining rates of anaerobic oxidation of methane (AOM) is more straight forward, since methane is the dominant alkane present in seep sediments. However it is critical to determine rates of AOM under quasi-in-situ conditions [64–66]. Determination of AOM rates at 1 atmosphere (shipboard or laboratory conditions) results in a dramatic underestimation of rates and an improved understanding of the rates and dynamics of AOM could reshape our understanding of how methane is processed at deep ocean hydrocarbon seeps. Oxidation of hydrocarbons at seeps generates bicarbonate, which drives authigenic carbonate formation. However, oil contains only minute amounts of nutrients and one of the biggest biological conundrums of hydrocarbon seep ecology involves the source of nutrients – namely nitrogen and phosphorus – that fuels the production of biological biomass [21]. Biological nitrogen fixation is one way to introduce new nitrogen into the system [67–70]. At brine seeps, the brine fluids themselves are enriched with ammonium and phosphate [22, 28] and these brines may serve as a unrecognized source of nutrients to fuel production of microbial and animal biomass in the deep sea. Achieving a better understanding of nutrient dynamics at hydrocarbon seeps is a fruitful area for future research.

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2.4.2 Biology The endemic fauna associated with hydrocarbon seeps are as diverse as they are beautiful (󳶳 Fig. 2.4). While some hydrocarbon seeps (i.e. actively venting mud volcanoes) are too volatile to support animal life, oil and gas seeps provide metabolic fuel and substrates for a diverse array of endemic animal life [60]. Seep fauna can be segregated into three general groups: Vestimentiferan tube worms, mytilid mussels and clams (vesicomyid, lucinid, or thyasirid). Clams are often found in habitats where mussels and tube worms cannot survive (i.e. associated with mud volcanoes). As seeps age and change, so does the fauna associated with them. Symbiotic mussels, clams and tube worms dominate when concentrations of hydrogen sulfide and methane are sufficient to support growth and provide maintenance energy. As seep-

Fig. 2.4: Hydrocarbon seep associated fauna. (a) Deepwater coral community with brittle stars & anemones; (b) Purple octocoral; (c) Deepwater corals and sea stars; (d) Deepwater corals and crabs; (e) Venus flytrap anemone; (f) Mussels, tubeworms and chemosynthetic bacteria; (g) Tubeworms; (h) Spider crabs; (i) Cold seep clam (Calyptogena pacifica); (j) Fish near oil seep; (k) shrimp and snails on gas hydrate; (l) Ice worms on methane hydrate; (m) Fish near gas hydrate; (n) Gulf hake near gas hydrate. Image credits: © ECOGIG (j,k); Dr. Ian MacDonald (ECOGIG) (e,g,h,l,m,n); Ocean Exploration Trust and ECOGIG (a-d,f); image of Calyptogena pacifica (i) Lovell and Libby Langstroth © California Academy of Sciences

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ages rates slow, or move to another location due to carbonate precipitation, cold water corals settle on authigenic carbonates and these animals provide habitat for a variety of associated fauna, such as brittle stars (󳶳 Fig. 2.4). Foraging fish, crabs, eels, gastropods, and deepwater sharks come in and out of seep areas, where they take advantage of an abundance of food that is ultimately derived from energy fed to the seabed from deeply sourced oil and gas. The surfaces of animal tubes and shells are often colonized by microorganisms or associated fauna (snails, crabs) who graze these microbial populations [60]. Gas hydrates are a particularly interesting example of a seep geological feature that serves as a home to diverse biological fauna. Methane “ice worms” (Hesiocaeca methanicola) were discovered in 1998 on gas hydrate mounds in the Gulf of Mexico [71]. These interesting polycheate worms create burrows in the surface of gas hydrates (󳶳 Fig. 2.4l) and are believed to graze on microorganisms inhabiting the hydrate surface.

2.4.3 Microbiology At marine gas and oil seeps, diverse microbial communities thrive in response to gas and oil exposure. At these sites, certain microorganisms have the ability to use the main constituents of gas and oil, i.e. hydrocarbons, as their carbon and energy/electron sources. Of the complex hydrocarbon mixture, microorganisms utilize methane, alkanes, cycloalkanes, and polycyclic aromatic hydrocarbons (PAHs). Methane is present in high concentrations at gas-dominated sites but also at oil seeps since methane can be part of the complex hydrocarbon mixture in crude oil. Shortchain alkanes are mainly found at gas-dominated seeps particularly at sites with elevated concentrations of propane and butane such mud volcanoes from the Gulf of Cadiz [72], mud volcanoes from the Central Nile deep-sea fan [73], and Gulf of Mexico seeps with structure II gas hydrates [74]. A broad range of short-chain and longer-chain alkanes as well as PAHs is typically found at oil seeps [74]. In anoxic sediments, sulfate is the most important electron acceptor since it is available in high concentrations (28 mM in seawater) and sulfate reduction is the dominant metabolic process accounting for up to 50% organic matter mineralization in many environments [75]. In the water column and in the upper few mm sediment depth, oxygen is the main electron acceptor and aerobic hydrocarbon oxidation processes dominate. 2.4.3.1 Aerobic methane oxidation Aerobic methane oxidation has historically been associated with methanotrophs belonging to Gammaproteobacteria (Type I methanotrophs), Alphaproteobacteria (Type II methanotrophs) and Verrucomicrobia (Type III methanotrophs). However, it is now

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clear that much greater diversity of aerobic methanotrophs exists and that the true diversity of aerobic methanotrophs has been vastly underestimated [76–78]. At marine seeps, methanotrophs occur in oxic environments such as the water column and the upper sediment layers but also in symbiosis with chemosynthetic communities, e.g. mussels [79]. Methanotrophs use methane as their sole carbon and energy/electron source, employing the key enzyme methane monooxygenase (pMMO). Since pMMOs lacks substrate specificity, several compounds (including short-chain alkanes) can be metabolized by these enzymes. Much remains to be learned about the aerobic methanotrophs, both in seep sediments, associated with seep animals and in the water column around seeps. 2.4.3.2 Aerobic non-methane hydrocarbon oxidation Aerobic non-methane hydrocarbon degraders affiliate mostly with Gammaproteobacteria (e.g. Oeanospirillum spp., Colwellia spp., Alcanivorax spp., Cycloclasticus spp., Oleiphilus spp., Oleispira spp., and Thalassolitus spp. but also with Alphaproteobacteria (e.g. Roseobacter and Rhodospirillales) [80]. Key enzymes of aerobic alkane oxidation are particulate copper monooxygenase enzymes (pMMOs or CuMMOs) [78]. These CuMMO enzymes are promiscuous and can also oxidize other alkanes (up to C8 ), branched and cyclic alkanes, and some aromatics. The oxidation of longer chain alkanes can be mediated by the AlkB enzymes, membrane-bound di-iron (non-heme) monooxygenases, or cytochrome P450 enzymes that also generate primary alcohols. For aromatic hydrocarbons, mono- and dioxygenases hydroxolate aromatic compounds and generate a range of intermediates (catechol, protocatechuate, etc.) [81]. 2.4.3.3 Anaerobic methane oxidation A very important methane sink at marine gas and oil seeps is the anaerobic oxidation of methane (AOM) [62]. AOM is mediated by anaerobic methane oxidizing Archaea (ANME), that are closely related to known methanogenic Archaea and found in a free-living state or in close association with bacterial partners, i.e. microbial consortia in densely packed aggregates or microbial mats. Three major clades of ANME were described: ANME-1 are abundant at diffusive methane-sulfate interfaces [82– 84], methane seeps in the Black Sea [85] and in hot seep habitats such as Guaymas Basin [86–88], ANME-2 are most abundant at cold seeps [89, 90], and ANME-3 occur at mud volcanoes [91] and the Eastern Mediterranean seeps [92]. Bacterial partners affiliate with Deltaproteobacteria of the groups SEEP-SRB1a [93], SEEP-SRB2 [94], HotSeep-1 (involved in thermophilic AOM) [87], and Desulfobulbus [91]. AOM was first discovered to be linked to sulfate reduction, which was confirmed by the tight association of ANME and sulfate-reducing bacteria [62]. The interrelationship between consortia-associated populations remained puzzling and it was hypothesized that methane was oxidized to carbon dioxide while reducing sulfate to hydrogen sulfide, splitting the energy (−34 kJ per mol substrate turnover at stan-

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dard conditions) between the populations. However, the intermediate(s) of these AOM consortia (metabolites or electrons) remained unknown. Later on, alternative mechanisms of AOM were discovered: methane oxidation was shown to be coupled to metal-oxide reduction [95]. In addition, the bacterium Methoxymirabilis oxyfera oxidizes methane under nitrite-reducing conditions by producing oxygen internally [96]. Recently, ANME-2 were shown to perform AOM coupled to sulfide disproportionation, eliminating the need for a microbial partner [97]. Under anaerobic conditions, methane is oxidized in cooperative microbial consortia or by single organisms and Archaea play a key role. In contrast, anaerobic non-methane hydrocarbon oxidation is mainly performed by Bacteria; however, a couple of exceptional discoveries were made for Archaea as well. 2.4.3.4 Anaerobic non-methane hydrocarbon oxidation An exciting discovery of late was the recently proposed genus ‘Candidatus Syntrophoarchaeum’. ‘Candidatus Syntrophoarchaeum’ are anaerobic thermophilic Archaea that live in partnership with the HotSeep-1 group and oxidize butane using enzymes similar to methyl-coenzyme M reductase, a key enzyme believed to be specific for methane consumption and production [97]. In addition, long-chain alkane (i.e. hexadecane) degradation under methanogenic conditions was demonstrated for anaerobic microorganisms including methanogenic Archaea and Bacteria of the genus Syntrophus [98]. Sulfate-reducing short-chain alkane degraders were enriched and isolated from marine seeps e.g. the strain BuS5, which was isolated from Guaymas Basin [99] and the enrichment culture Butane-GMe12 from Gulf of Mexico [99, 100]. Also aromatic hydrocarbon-degrading sulfate-reducing bacteria include isolates from marine environments: strain EbS7, which oxidizes ethylbenzene, was isolated from Guaymas Basin [101]. Cultivation-independent studies demonstrated that Deltaproteobacteria and more specifically Desulfosarcina/Desulfococcus dominated gas and oil seeps [74, 94, 102]. Indeed, diverse subgroups of the Desulfosarcina/Desulfococcus clade were identified as key alkane degraders at marine seeps sites (Amon Mud Volcano and Guaymas Basin) [103]. These alkane-degrading groups were, according to their substrate usage of butane (short-chain alkane; SCA) or dodecane (long-chain alkane; LCA), termed “SCA1”, “SCA2”, “LCA1”, and “LCA2”. Interestingly, group SCA1 comprises strains Bus5, Butane-GMe12, and related spp., while all other groups comprise uncultured microorganisms, suggesting that most of the in situ active hydrocarbon degraders are currently uncultivated. Possible hydrocarbon degradation pathways of sulfate-reducing bacteria are mainly analogous to that of the denitrifying bacteria [104]. The most common activation mechanism under anoxic conditions for non-methane alkanes and alkyl-substituted aromatic hydrocarbons is fumarate activation [105]. Glycyl radical enzymes

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that were proposed for the anaerobic hydrocarbon-activation are the benzylsuccinate synthase (Bss; [106, 107], the (1-methyl)alkylsuccinate synthase (Mas or Ass; [108, 109] and (2-naphthylmethyl)succinate synthase (Nms; [110]). 2.4.3.5 Local vs. Global patterns/distributions Local and global patterns of microbial communities at seeps are influenced by environmental factors such as the bioavailability of gas- and oil- derived hydrocarbons, the availability of electron acceptors (e.g. aerobic vs. anaerobic hydrocarbon degradation), nutrients (e.g. phosphorus, nitrogen, essential trace metals), pH, temperature, pressure, and water depth. Therefore, the microbial communities among or within seep sites are highly influenced by the local environment. Microbial communities at marine seeps may resemble each other at the phylum or class level, but they separate with increasing phylogenetic resolution, i.e. genus or species/operational taxonomic unit (OTU) level. With the new genomic era including next-generation sequencing technologies as well as sophisticated computational tools that allow sub-OTU resolution, global comparisons of microbial communities among seeps and in-depth analysis of the diversity of microbial communities at specific seep sites are possible. Natural hydrocarbon seep communities harbor distinct bacterial and archaeal taxa linked to key biogeochemical functions, such as hydrocarbon degradation. These bacterial and archaeal taxa are globally widespread but appear with elevated abundance at seep sites. For instance, a global study of methane-dominated seeps showed high relative sequence abundance of ANME, aerobic Methylococcales, sulfate-reducing Desulfobacterales, and sulfide-oxidizing Thiotrichales [111]. 2.4.3.6 Microbial mats One of the most striking features of hydrocarbon seep ecosystems is the presence of chemoautotrophic microbial mats [112]. The dominant microorganisms in microbial mats oxidize hydrogen sulfide to elemental sulfur and/or sulfate and fix carbon dioxide into biomass. They thus recycle reduced sulfur to sulfate and trap carbon dioxide, potentially derived from oil and/or gas oxidation into biomass, making them key biogeochemical players at hydrocarbon seeps. Typically, these microbial mats are comprised of sulfide-oxidizing Gammaproteobacteria, often of the genus Beggiatoa [113], but other sulfide oxidizers, such as Thiomargarita spp. also form mats at hydrocarbon seeps [114]. These microbial mats can be colorful, especially Beggiatoa spp., due to the dominant respiratory pigments employed by the bacteria. As such, microbial mats form a “bulls-eye” target that reveals areas of diffuse fluid release along the seabed. At high flow mud volcano systems, mats of Epsilonproteobacteria have been noted [115].

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2.5 Closing remarks Hydrocarbon seep ecosystems are some of the most remarkable deepsea environments known. The diverse geological settings and drivers, the spectacular seafloor features created by hydrocarbon seepage and the complex biological communities supported – all fueled by deeply sourced metabolic substrates – illustrates the dynamic connections between deep carbon reservoirs and processes at the seabed. Of the known and suspected hydrocarbon seeps on earth, only a fraction have been explored and studied in detail. Hydrocarbon seeps are thus ripe regions for exploration and discovery in the future. Acknowledgment: Funding for the preparation of this chapter was kindly provided by the National Science Foundation Emerging Frontiers Program (grant EF-0801741) and the Gulf of Mexico Research Initiative’s “Ecosystem Impacts of Oil and Gas Inputs to the Gulf-2” (ECOGIG-2) Program. We thank Dr. Eric Marty for assistance with preparation of the figures. This is ECOGIG contribution number 484.

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Walter Menapace, Achim Kopf, Matthias Zabel, and Dirk de Beer

3 Mud volcanoes as dynamic sedimentary phenomena that host marine ecosystems 3.1 Abstract Mud volcanoes (MVs) release reduced substances, especially methane, which is a powerful greenhouse gas with the potential to feed deep sea ecosystems. In this chapter we provide basic geological information on submarine MVs for the curious nonspecialist, who would like to better understand how MV ecosystems and biodiversities are shaped. Crucial for life is a suitable chemical and physical environment, and the temporal dynamics thereof. Here we describe the appearance of MVs, where and how we can find them, the chemistry of their sediment-water interface and their physical dynamics. We will define how most submarine MVs are formed, and provide a classification based on the formation mechanism. Their global distribution is also strongly linked to the same formation mechanism, be it driven by tectonics or sedimentation. The formation mechanisms often define the depth of origin of the MV products and hence determine the chemical compositions of the released fluids and solids. Products of MVs include elevated dissolved inorganic carbon (DIC), methane and ammonium, fine mud, breccia and occasionally asphalt. The fluid flow velocity also shapes the ecosystem: higher flow velocities supply more reduced substances but velocities that are too high to suppress the supply of oxidants. A distribution of fluid flow, with the highest near the venting sites, may control the heterogeneity of the ecosystem. Finally, we will discuss the physical stability of MVs as ecosystems, since the intervals between eruption determine the maximal age of the ecosystem and its inhabitants. Conversely, the inhabitants can provide information on the stability and age of the MV.

3.2 Introduction 3.2.1 Types of MV, expelled products, morphologies and dimensions MVs are often described as raised sedimentary expulsion features that present inner and outer structures very similar to magmatic volcanoes (conical main edi fice, feeder channel, secondary craters, calderas, stratification; 󳶳 Fig. 3.1a) [1, 2]. Their venting activity is related to mud, clast, gas, and water emissions. Due to the high volume of these emissions and the profound roots of the edifices, they are considered one of the most effective ways of solid and fluid release from deeper sediments to the surface [4]. https://doi.org/10.1515/9783110493672-003

54 | 3 Mud volcanoes as dynamic sedimentary phenomena that host marine ecosystems

(a)

(b)

Fig. 3.1: (a) Schematic cross section of a typical MV structure summarizing the main features likely to occur in such a setting [3]; (b)1,2 Examples of MV sedimentary clasts in a gray mud matrix with the typical moussy structure from degassing (Venere MV, Calabrian Arc)

The ejecta of MVs are generally termed mud breccia, according to Cita [5, 6], who described it as a “structureless pebbly mud with dominantly angular semi-indurated clasts” (󳶳 Fig. 3.1b). The clast fraction of the mud breccia has a wide compositional range, while the matrix is mainly made of silty clay. The high variability of the clasts is linked to the characteristic of the rocks that are passed by the ascending mud; they can range in diameter from millimeters to meters, depending on the size of the conduit, the viscosity of the mud matrix, the physical properties of the clasts themselves, and the velocity of ascent [7]. The viscosity of the matrix and the ascent velocity are both direct functions of the amount of fluids available in a given MV (see also next paragraph). MVs manifest mostly as cone-shaped, positive seafloor morphologies, according to the viscosity of the erupted mud [8]. They form shield-like edifices (low viscosity, high fluids presence, flat morphology), or more pronounced conical forms (high viscosity, low fluids presence) [3, 9]. Occasionally there have been reported accumulations of soupy/liquid mud in craters of collapsed MV edifices, which were called mud pools [10]. The edifices are built by the generation of different mudflows which, in a volcano-like style, pile up on the flanks of the structures after every subsequent eruption, forming mud domes [11]. The size ranges from tens to hundreds of meters in height and hundreds of meters to a few kilometers in diameter [9], although smaller

3.2 Introduction

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features are often found in areas subject to mud volcanism [12]. The conduit diameter can reach tens of meters according to seismic images [13–15], and several kilometers deep, based on the geochemical signature of the expelled fluids sampled on top of MVs [16, 17]. Most of the time, however, the singular structures have undergone different phases of evolution, resulting in composite MV systems.

3.2.2 Identification on the seafloor and classification MVs are identified through a different suite of methods that evolved over time, from the first evidence of submarine mud volcanism obtained with side scan sonar surveys in the 1980s [18] to the state-of-the-art detailed microbathymetric mapping using autonomous underwater vehicles (AUVs) [19]. Most of the current techniques are complementary and often applied together (󳶳 Fig. 3.2a,b). The location of MVs on the seafloor is obtained through ship based hydroacoustic surveys, which include bathymetry mapping, measuring of seafloor backscatter intensity, as well as water column imaging (locating gas flares by fishing sonars as in 󳶳 Fig. 3.2b) [20–22]. Seismic images and AUV microbathymetry, although less common, are useful for the remote identification of MVs (󳶳 Fig. 3.2a) [13, 19].

Fig. 3.2: (a) Seismic profile through a MV located offshore of Trinidad [23]; (b) Example of a gas flare in fan (b1) and stack (b2) view on top of MV#1 (Kumano Basin, Japan) as imaged through a multibeam echosounder

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The in situ techniques applied to MVs usually aim to test their current activity, age, and depth of the source and constitute sediment and pore water sampling, heat flow measurements, and visual inspection with the help of remotely operated vehicles (ROVs) [24–26]. An important advancement in in situ techniques for MVs is long-term measurements, which generally focus on the determination of pressure, temperature, water geochemistry and volume of gas emissions [27]. These measurements, normally conducted through monitoring stations, can provide more meaningful insight into MV evolutionary patterns due to the extended time component. Different publications have addressed the issue of classifying MVs, although most focus on the terrestrial realm [1, 28]. One of the first field description of MVs was carried out by Shih in 1967 [29], who categorized the Taiwan MV types according to their geomorphology. A robust classification system is based on the paroxysms and frequency of the eruptions and type of products [8]. In the submarine realm, various classifications have been compiled with the intent of describing these structures, following the subsequent criteria: – morphology of the edifice (shape, size) [30] – morphogenesis (tectonic/geophysical correlations) [31] – type of emissions (liquid, solid and/or gaseous) [32] – evolutionary stage [32] – seismic reflector shapes [30] Three main groups of MVs exist: 1. MVs in collision zones and compressional tectonic settings (approx. 60–70%) 2. MVs driven by compaction (approx. 20%) 3. Other MVs (driven by hydrothermal fluids, hydrocarbon leaking, groundwater infiltration, permafrost thawing, deep mantle processes) For more detailed information about the geodynamical settings see Sect. 3.3.

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3.3 Geological significance 3.3.1 MV number estimation and distribution MVs are morphological features known from the beginning of the 17th century [3], first identified on land and subsequently (1970s) recognized on the sea floor in areas of high tectonic activity. The distribution and number of MVs worldwide is not known exactly but it is estimated to be between 1,000 and 100,000, for confirmed and inferred respectively [7, 33, 34] (󳶳 Fig. 3.3). This phenomenon is widespread and, at the same time, still poorly defined due to the large extent of unexplored oceanic seafloor and the lack of a global updated census of MVs, as has recently been done for the Mediterranean Sea [35]. The occurrence of MVs is more prominent in subduction zones, where there is convergence of two tectonic plates [9]. Plate collision creates overpressured sediments and fluids at depth, which tend to escape towards the seafloor, producing MVs. Mud volcanism has been observed in the nonsubducting plate over back arcs, forearc basins, and within the frontal domain of accretionary prisms, where sediments are

Fig. 3.3: Occurrence of MVs in densely populated areas on Earth [9]

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accumulated; the most affected areas are the Mediterranean Sea, the Gulf of Cadiz, the Black Sea, Japan, offshore Barbados, and the Cascadia Basin [9, 36–39]. Other key areas of mud volcanism are submarine fan and deltaic complexes of very large rivers, e.g. the Nile [40, 41] or the Niger [32, 42], where the origin of mud volcanism resides in compaction of deltaic sediments under the increasing load of the overlying strata or in the production of methane from the decomposition of organic material transported through the rivers. The presence of MVs has also been observed on continental slopes in the Beaufort Sea [19], in the Barents Sea [26] and in the Gulf of Mexico [43]; these features are driven respectively by gas hydrate decomposition, compaction of glacial deposits or from complex interactions with salt diapirs and hydrocarbon reservoirs at depth (possible oil leaking through MV conduits). To a smaller degree, faulting and mud diapirism can also contribute to the formation of MV features, e.g. in the Alboran Sea [44]. The Mariana Convergent Margin, with its huge serpentinite mud domes (reaching 50 km diameter and 2 km height, [45], is the typical example of a region affected by fracture controlled mud volcanism, where the subduction of seamounts on the Pacific plate is causing the deformation of the overlying Philippine plate, creating preferential pathways for the ascent of the serpentine muds [46, 47]. The driving force for the formation of the Mariana MV is the release of water from the descending slab due to compaction, which in turn fuels the hydration processes in the mantle, creating the serpentine muds. Those muds are then pushed upwards through the fracture zones, mainly driven by buoyancy contrasts with the surrounding sediments. The expulsion mechanism hypothesized for serpentine MVs does not include a major contribution of hydrocarbon gases, which in sedimentary MVs are considered to have an important role in developing overpressure conditions at relatively shallow depths below the seafloor (bsf).

3.3.2 Formation models and long-term evolution MVs are formed due to pressure build-up at depth from numerous physical and chemical sources such as mineral dehydration, sediment compaction, fluid migration, gas hydrate decomposition, earthquakes or biogenic/thermogenic methane production [3, 9]. Given the complex nature and inaccessibility of many MVs, only ocean drilling and core dating can provide reliable data on the temporal evolution of MV features [48]. The sources’ origins can be generally grouped into shallow (0–5 km) and deep (5–15 km) types. The shallow sourced products (e.g. mud breccias, water, gases) are related to sediment compaction, gas generation and tectonic features such as faults and folds [14, 49], while the deep sourced ones (mainly water and mud) often originate

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Fig. 3.4: Ideal sketch of MV genesis according to (a) solid models and (b) fluid models [23]

directly from the subducting slab and are mainly associated with mineral dehydration, high thermal gradients and buoyancy contrasts [16, 50]. A feeder channel or a high permeability zone below the edifice is thought to be the primary way for the deeper materials to reach the surface, as confirmed by seismic images [13, 51]. Structural deformations of the hosting sediments, i.e. folds and rocks, may favor such feeder pathways [15, 52] or on the contrary, the presence of more competent strata during the mud ascent could enhance the formation of a mud chamber at shallow depths [50, 51]. Different models for the origin of MVs have been considered in the literature, which are mainly driven either by buoyancy contrasts, the so called ‘solid models’, or by hydrofracturing of sediments and mud fluidization, also known as ‘fluid models’ (󳶳 Fig. 3.4) [3, 4, 9]. Buoyancy contrasts are formed due to differences in density between the rising mud (diapir) and the surrounding sediments, while hydrofracturing of sediments happens either at depth (plate interfaces) or on top of mud intrusions owing to the high fluid pressure released from sediment compaction, mineral dehydration or hydrocarbon cracking, which breaks apart the overlying strata [23]. Mud fluidization implies the rapid and abundant supply of fluids in order to liquefy a source layer and foster its rise through the sediment column, thus forming pillar-like structures [23]. These processes are strictly related to one another and often interdependent, according to the geology of a specific area, the characteristics of the MVs sources and the type of hosting sediments.

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Fig. 3.5: Four fundamental stages of MV evolution [9, 53]

MVs evolve, similarly to magmatic volcanoes, through different stages (󳶳 Fig. 3.5) mainly dependent on deep sediment/fluid input and sedimentation rate [9, 53]. These stages could be essentially divided into four successive phases (󳶳 Fig. 3.5): Eruption Hydraulic failure of strata within the overpressured feeder channel, formation of the MV edifice with subsequent mud flows; Depletion Migration of fluids to the surface through cracks, gryphons and adjacent high permeable surroundings; Quiescence and build-up Subsidence of the MV edifice, eventual formation of caldera (depending on the presence of a mud chamber) and accumulation of pore pressure; Reactivation Breakout of an eventual ‘plug’ and resuming of the activity. In addition, these sedimentary features have been found to eject large quantities of mud breccia, water and gas, following episodic patterns [54–56]. The principal methods applied on MVs, upon which the episodicity idea is based, are fluid and heat flow measurements [24, 57], bathymetric monitoring [58], seismic images of “Christmas tree” structures in the subbottom [30, 55], geochemical analyses [59], modeling of the eruptions [2] and occurrence of consecutive mud flows in sediment cores [12, 48, 60]. Every one of these techniques highlights changing aspects of MVs over time, supporting the occurrence of multiple eruption episodes during the lifetime of a MV.

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3.4 Mass transfer and fluid cycling fuel ecosystems 3.4.1 Composition and sources of gas, water, mud, and clasts The compositions of the ejecta from MVs are of interest for geologists, as they contain information on underlying geological processes, and for biologists, as they provide data on the ecosystems of MVs. In order to unravel the source depth of the various MV products in a given region, various chemical and mineralogical techniques have been used. For the solid phase, one has to distinguish between the origin of the muddy matrix and of the clasts that are dragged upward with the mud. Naturally, the former is from the same depth or a greater one than the latter, because the clasts usually originate from the wall of the diapiric intrusion [61]. There are numerous ways to estimate the depth of clast origins, e.g. by comparing analog lithologies in neighboring boreholes, onshore and on outcrops. Water is liberated from clay minerals in a process called smectite-illite transformation, which takes place under elevated pressure and temperature conditions (p-T) (> 100–150 °C and > 2,000 m depth) [62]. The illite crystallinity measurements that are possible in the ejecta, owing to the smectite-illlite transformation, provide information on their depth and pressure of formation. The ratio between smectite and diagenetically formed illite is used to calculate the maximum temperature of burial [63]. The light reflectance of vitrinite (organic matter residues from plant material) is used to determine the p-T conditions of the ejecta’s source. Vitrinite analysis can be used when organic matter occurs in sufficient abundance and the geothermal gradient of the area is well constrained. All the aforementioned methods have been performed on many MVs; however, a particularly good example is the MVs on the Eastern Mediterranean Ridge, where the whole suite of methods was extensively applied (e.g. [64]). For the aqueous fluids, a large number of studies attest that both mineral desorption and dehydration of hydrous clays (smectite group) are the main causes for the formation of deep-seated MV fluids [65]. Both processes are closely coupled to the burial and heating of sediments in subduction zones. During dehydration, clays lose their interlayer water or change mineral phase (smectite-ilite transformation, see glossary and 󳶳 Fig. 3.6). The release of hot and usually weakly ionized water (with the exception of fluid venting related to brine seepage [5]) results in other, secondary reactions. These range from the release of solutes (e.g. due to cation exchange, which could result in an increase in Na+ and decrease in NH+4 concentrations) to kinetic or equilibrium driven processes of isotope fractionation. Although many of these chemical reactions are theoretically well understood, the interpretation of distinct chemical signatures is quite complicated and often a challenge due to the effects of alteration processes and mixing during the advective migration of the fluids. Last but not least, the progressive degradation of organic matter has to be considered (e.g. associated with NH+4 and carbonate 2+ alkalinity increase, SO2− 4 decrease and the release of reduced constituents like Fe ).

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Fig. 3.6: Temperature distribution vs. different sources of fluid expulsion in a subduction zone (after [66])

Of course, compositions of the initial rock/sediment, the nature of burial/subduction processes, and p-T conditions of the source have a significant influence on the characteristics of primary fluids. In the next paragraph some of the main geochemical markers are described. Further information on fluids in subduction zones can be found in a recent detailed overview [65]. For the gaseous phase, methane accounts for > 95% of all MV gases found in the various geological settings [9]. As a result of either burial to several kilometers depth or tectonic processes such as underthrusting, sediments rich in organic matter are exposed to elevated p-T conditions so that methane and, to a much smaller extent, higher hydrocarbons, will form (e.g. [49, 65]). By far the most abundant hydrocarbons are the ones of thermogenic origin. These gases are identified by characteristic δ13 C ratios, showing a relatively higher content of 13 C than those from biogenic processes [67]. The reason is that enzymatic processes have a slight preference for isotopically lighter substrates. In settings with hydrothermal fluid circulation (e.g. at the flanks of Mt. Etna, Mediterranean Sea), CO2 may be the dominant gas phase in MVs, due to magma degassing that provides the buoyant phase, rather than the maturation of organic matter as elsewhere [68]. It has also been reported that contributions from mantle gases such as He are found in MVs, highlighting that the source of MV ejecta could be as deep as the mantle [61]. General indicators of a freshening (decrease in ionic strength) in MV fluids, mainly by melting methane hydrates and entrapment of ions in clay are: (i) depleted K+ concentrations when poorly hydrated K+ is incorporated between crystal layers during smectite-illite transformation [69], (ii) low chlorinity and (iii) enrichment in δ18 O (these latter two are due to the dissociation of methane hydrates, which are enriched

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in H2 18 O and strongly depleted in ions). In combination, the 35/37 Cl isotope ratio variations in Cl concentration can also be used to distinguish between different source reactions (clay dehydration, gas hydrate dissociation and the formation of high temperature hydrous serpentine minerals). For similar purposes, the comparison between δ18 O and δD, as well as boron and lithium concentrations together with δ11 B, can also be used [70–73]. Some of these parameters also serve as geothermometers (e.g. δ11 B ratios; [61]), so that the origin of the fluids can be estimated as long as the geothermal gradient in a determined MV area is known. Episodic expulsion of fluids can also be recorded by isotope composition (87 Sr/87 Sr) in mineral clasts ([74]; Section 3.4.3). Whereas the isotope ratios are of interest for geologists seeking the origins of ejecta, essential for life is the presence of reduced species, trace elements/metals and nutrients, which are also transported to the uppermost sediment layers and the seafloor. Their origin and release are mainly associated with the transformation of organic compounds, either via microbially mediated degradation, and the accompanied reduction of electron acceptors (e.g. release of NH+4 , Fe2+ , H2 S, biogenic CH4 ), or via thermally activated breakdown under high p-T (thermogenic CH4 ). However, besides methane, efflux rates of these reduced constituents are usually very low or below detection (e.g. [75, 76]). In most cases, the reasons for this observation may be a combination of two factors: the availability of oxidizing agents (primarily free oxygen), and the transport mechanisms of the fluids – slow vs. fast. As a consequence, not only can the composition of released fluids be crucial for the nature and diversity of benthic habitats on MVs, but the velocity of the advective flow can also shape the ecosystem. The latter especially controls the balance between the release of reduced substances and the availability of oxidants. Additionally, fluid compositions vary from MV to MV in subduction zones [45, 75, 77], but also within one single structure over relatively short distances. A study on the Håkon Mosby MV [78] shows that microbial communities in the center of the active area depend on the oxidation of methane, while in the surroundings, with lower flow rates, biological oxidation of sulfide is much more important than chemical anaerobic oxidation. We will come back to the critical factor of porewater upflow velocity in the next two sections.

3.4.2 Seafloor and subseafloor ecosystems on MVs Apart from CO2 seeps, most MVs eject reduced substances into an oxidized water column; thus a thermodynamic disequilibrium is generated at the sediment-water interface. This can drive primary production by chemosynthetic bacteria, which is the reason MVs can be considered oases of life. We will explain briefly how the chemical energy in MVs can lead to primary production and fuel an ecosystem. Primary production by chemosynthesis is the energy consuming incorporation of CO2 into biomass, where the energy is supplied by the oxidation of inorganic compounds. MVs harbor complete food webs analogous to those in the photic zone that

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rely on solar energy, where radiation is the energy source for primary production. The main oxidants in seawater are sulfate (28 mM) and oxygen (∼0.2 mM) [79]. The most common substrates for chemolithoautotrophic microorganisms are CH4 , H2 , CO, sulfide, sulfur, Fe2+ , and ammonium. Whereas all these compounds are formed by thermal processes, not all will reach the surface. H2 is readily converted by methanogens to methane (4H2 +CO2 → CH4 +2H2 O) [80]. CO is anaerobically converted (CO+H2 O → H2 + CO2 ) [80], and the produced hydrogen reforms the substrate for methanogens. Typical reductants in surficial mud fluids are methane, sulfide, Fe2+ and ammonium, which can all be oxidized biologically [80]. The most abundant reductant in MV fluids is normally methane, which can be oxidized with either oxygen or sulfate as an electron acceptor. The C atoms from methane can be directly inserted into biomass [81]; however, old and new evidence suggest that many strains produce their cellular biomass from CO2 fixation [82, 83]. Methane oxidizers are autotrophs, as they generate biomass from C1 compounds. Although aerobic methane oxidation has a high yield (one gram dry weight as biomass per gram of methane), the areal conversion rates are limited by low oxygen supply (approximately 0.2 mM in seawater), making this process a weak basis for a food web. Most methane is oxidized by anaerobic methane oxidation (AOM) using the abundant sulfate (28 mM in seawater) as electron acceptor. The net process is + CH4 + SO2− 4 + 2H → H2 S + CO2 + 2 H2 O

The process is driven by a symbiosis between archaea (anaerobic methane oxidizers: ANME) that convert the methane, and sulfate reducers that transfer the electrons to sulfate [84]. It is still debated how the redox equivalents are transferred between the archaea and the sulfate reducers. Evidence has been presented for direct transfer via electrical conductive wiring [85, 86], while diffusive transfer via chemical redoxmediators has also been shown to be possible [87]. A concept that avoids electron transfer between populations was proposed by F. Widdel [88]. In this model the archaea perform both the methane oxidation and the sulfate reduction, however only to the level of S(0), which then in the population of sulfate reducers is converted by sulfur disproportionation to sulfide and sulfate. Whatever the details of the process are, AOM has been shown to have a very low biomass yield and the involved microbial consortia have very slow growth rates with doubling times of months [89], which is not a good basis for a food web. However, AOM produces sulfide, and aerobic sulfide oxidation is coupled to rapid biomass growth with high yield. Therefore, sulfide oxidation is an important process for primary production and the basis of the food webs in marine seeps. Additionally, the combination of AOM and sulfide oxidation shields the biosphere from methane and sulfide input [84]. An overview of these processes is given in 󳶳 Fig. 3.7. Sulfide oxidizers often form conspicuous mats that are white from intra- and extracellular sulfur globules. Typical mat-forming organisms are Arcobacter [90] and the so-called Big Bacteria [91]. In fact, cold seeps are often found by searching for these white mats. Arcobacter is an aerobic organism and its mats can develop in hours [92].

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Fig. 3.7: Pathways of methane oxidation: anaerobic methane oxidation is visualized by the red and green consortium, respectively ANME archaea and sulfate-reducing bacteria. When AOM is well established, no methane passes the methane-sulfate transition zone and aerobic methane oxidation does not occur (from F. Stams, Wageningen University)

Examples of Big Bacteria are the long filamentous Beggiatoa and Thioploca up to 160 µm thick. Their slime jets provide a rapid gliding motility in the 2–5 cm thick suboxic zone (where both oxygen and sulfide are absent) of sediments. Most of their cell volume is a vacuole that can be filled with up to 0.5 M nitrate [91]. By respiring the nitrate, they can ‘scuba dive’ in sulfidic sediment for up to 40 days before refueling at the surface [93]. Access to deeper sources of sulfide provides an obvious competitive advantage over aerobic sulfur oxidizers. Interestingly, aerobic sulfide oxidizers often live in symbiotic associations with marine animals like tube worms, snails, bivalves, crustaceans and nematodes [94]. These symbiotic animals can reach enormous densities near or on seeps. Their fossil remains can be used to recognize extinct cold seeps. Energetics predicts that, based on yields and reductant supply, most growth in such environments is based on sulfide oxidation. Nevertheless, a bewildering diversity of symbionts can be found in a single host, occasionally combining sulfate reducers and hydrogen, methane and sulfide oxidizers in one host [94]. The importance of each of these processes for the host is not easy to assess. Remarkably, the growth rates of these organisms are highly variable. For example, Riftia tubeworms can colonize a seep area in a few weeks to months, whereas vestimentiferan tubeworms can be 250 years old [95]. Cold seeps are thought to follow a sequence of colonizers: first aerobic microbial mats develop, followed by Beggiatoa mats, then relatively fast growing bivalves that harvest sulfide from up to 10 cm bsf, followed by the very slow growing tube worms that ‘mine’ for sulfide up to 2 m deep [96]. The latter can reach ages of at least 250 years; thus if these are found, the seeps are old and stable. Such a sequence will restart after each eruption, when the sediment surface is strongly perturbed and the ecosystem is destroyed.

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3.4.3 MV episodicity The episodic activity of MVs starts with the formation of the features and is highly episodic. As stated in Sect 3.3.2, MVs may evolve rather suddenly and rapidly form a cone of clast-bearing mud and breccia. Since the majority of the features occur in convergent margin settings, geodynamic forces govern mud volcanism. Similar to earthquakes within the seismic cycle, mud volcanic eruptions, mudflows or even variations in seepage rates are driven by the overall tectonohydrological regime. On many occasions, stress changes at depth have been responsible for mud extrusion, most importantly because there is a causal link between MV activity and earthquakes. In the Caucasus, historic eruptions of MVs on the Caspian Sea coast occurred prior to large earthquakes (EQs). Chemists also attested that the chemistry of waters and gases changed a few days before a given EQ. Such studies are easier to perform onshore, where the detection of radon anomalies has become a powerful tracer for EQs [97], sometimes even prior to EQs [98]. Similarly, methane gas may migrate shortly before, during and shortly after EQs through sedimentary bodies, especially in pockmark areas on the shelf [99] or on MVs [100]. Also, ionic concentrations of some elements have been shown to vary in the case of fault slip events in subseafloor pore waters [101, 102]. The wealth of processes sketched above all rely on the enhancement of permeability, which is causally connected to seismic wave propagation [103] and subsurface fluid flow onshore (e.g. [104]) and offshore [105]. One such recent example is the massive release of methane from a submarine MV in the Kumano Basin, Japan after an EQ [106]. However, nontectonic triggers of mud volcanic activity have also been reported (Håkon Mosby, Nile Delta) and include Earth tides, thermal pulses or discontinuous dewatering. Very scarce direct observations exist of the development of submarine MVs, but MVs on land can develop within days, and then stay stable for days, weeks, years or much longer. The longevity is believed to be a function of mud breccia physical properties (i.e. clast rich ejecta withstand erosional forces more efficiently; [9]) or the nearstress field and mud reservoir/source layer volume (which may keep the conduit open and flux alive). This perception is indirectly supported by scientific drilling into multimillion-year-old mud domes [107, 108] in the Eastern Mediterranean Sea. Drilling as well as seismic reflection surveying attest to episodicity, as apparent for the Christmas tree structure (󳶳 Fig. 3.2a), where each mud flow event represents a branch of the tree while hemipelagic background sedimentation takes place during periods of MV quiescence. Ocean Drilling Program (ODP) Expedition 160 drilled into two such MVs in the Eastern Mediterranean sea where mud flows and interbedded hemipelagites were recovered and dated to be several million years old [107]. Grain size distribution of the oldest mud breccia recovered further allowed researchers to associate the initial deposits with clast supported sediment while the youngest flows showed fine grained, gassy strata [109]. This mechanism is typical for deep seated MVs in subduction settings since the mud has to first clear the fault zone of breccia and gouge it out before

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mud may ascend. Once completed, the ascent of gassy mud may be more rapid and limited in clasts, which are only collected as wall rock samples on the mud mass’s way to the seafloor. Note that core dating utilizing biostratigraphic methods (using indicative microfossils) or from stratigraphic correlation as well as assuming certain sediment accumulation rates work best in the range of myrs to several kyrs, helping researchers to understand only the most important episodes in MV evolution (as described in Chap. 4).

3.4.4 MV short term evolution If we regard more recent variations in mud volcanic activity, there are two fundamentally different approaches: (i) to study sites where mud extrusion occurs suddenly for the first time, or (ii) to monitor transient changes on existing, well-known MVs. As for category (i), sites of first time mud eruption are usually spectacular and are associated with a large volume of extruded mud (breccia) since the reservoir/layer of origin is tapped into for the first time. Very recent examples of massive mud extrusion occurred in the Caspian Sea, offshore Pakistan (Makran margin), and in Indonesia where the mud pierced the seafloor at shallow water depth and accumulated sufficient volume to form an island. These islands, however, were built of homogenous, fine grained mud almost free of clasts and usually lasted only a few weeks after extrusion ceased, as wave action eroded the island. However, one prominent example that has now been active since 2016 is Lusi MV in NE Java, Indonesia (e.g. [110]). The mud flow rate reached up to 180,000 m3 /d, covered a region > 7 km2 , and forced the evacuation of several villages. Several factors may have contributed to the sudden outburst: nearby drilling activity [111] or reactivation of a fault after a M6.3 EQ in May 2006 [112]; the latter fluidizing some deeper strata and triggering mud ascent. As for studying category (ii), which represents sites where eruptive activity is less violent and thus biological inhabitation is likely, changes on timescales longer (kyrs to months) than those of category (i), researchers rely on two techniques: monitoring devices deployed on MVs (see list below), and repeated seismic/bathymetric mapping of the subsurface/ seafloor to identify micromorphological changes, new mud flow deposits, etc. Both approaches are powerful and can be used in combination. The monitoring devices, which have proven to be effective in measuring several MV parameters, can mainly be grouped in these categories: Osmotic pumps, which sample MV fluids (on the seafloor or deployed in boreholes) for further geochemical analyses [27, 113, 114] – here changes in fluid geochemical signature may be indicative of the depth of origin and hence attest to variability over time. p-T sensors or sensors strings, normally inserted in the sediment or in boreholes for measuring anomalies or gradients in pressure or temperature [24, 115–117] – here the episodic increase in p and T is usually associated with hydraulic connec-

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tivity to deeper levels, as experienced in EQs. Pore pressure fluctuations provide possible links between earthquakes and MV episodicity [100, 118] and ongoing studies are refining these links [117]. High temperature gradients around MVs are on the order of several hundreds of °C/km, whereas typical marine sediment background values are less than 40 °C/km [24] – note that high gradients do not represent the geotherm at the site, but are rather a measure of how recent the flow is and if continuous fluid discharge from depth takes place. Flow meters and benthic chambers, lying on the seafloor, mainly dealing with quantification of flow rates [114, 119, 120] – similar to the previous point, higher flux is associated with stress changes at depth, with the MV basically acting as a valve in deformation episodes. Flow rate quantifications using flow meters or inverse modeling from geochemical or thermal measurements give typical values of fluid flow on MVs in the order of tens of cm/yr [24, 115, 119], whereas quiescent features may show less than 1 cm/yr (e.g. [17]). Sonars or ROV observations, essentially bubble quantification of seeping gases, but also visual variations of surface sediments [116, 121] – bubbling may be associated with tectonic forcing, but has also been reported to be linked to temporal

Fig. 3.8: Example of mud volcano #2 in the Kumano Basin, Japan: (a) differential bathymetry map, indicating an increase in the fill of the moat around the dome by several meters; (b) in situ HF (heat flow) data across the mud dome, showing anomalously graphs with depth that indicate gassy mud flow deposits even outside of the moat (N. Kaul, unpublished data)

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dissociation of gas hydrate in MVs in the Beaufort Sea. Quantitative estimates are scarce; however the methane flux out of MVs obtained from visual observations and amount of anaerobically oxidized CH4 ranges from 104 to 106 mol/yr for single features [116, 122]. Another powerful means to determine short term episodicity in MVs is the repeated mapping of the seafloor [24, 58, 123]: to use differential bathymetry to identify the most active zones of mud deposition. In two recent expeditions in the Kumano Basin, Japan, it was shown that a gas hydrate bearing MV had massive mud flows between summer 2012 and fall 2016, which filled most of the moat around the feature (󳶳 Fig. 3.8). Evidence for mud extrusions further came from in situ T measurements and coring next to the mud volcano, which also showed gassy mud breccia and elevated heat flow. If ground truthing is not available, changes in multibeam backscatter intensity of the surficial sediments may also attest to mud breccia deposition (owing to high clast content) or consolidation patterns (i.e. high or low backscatter values corresponding to a younger/older mud flow; [124, 125]). Both the in situ measurements and the geophysical mapping provide hints concerning episodic hydrological or geological activity and have repercussions for biological activity on MV features.

3.4.5 Methane in MVs: the Håkon Mosby case study The release rate of methane is a critical parameter as it is a greenhouse gas and a food source for the local ecosystem. Methane is released as gas bubbles and in dissolved form. The gas plumes are clearly visible on fishing sonar and disappear at approximately 300–400 m below the ocean surface [121]. This coincides with the methane hydrate stability depth. Below 400 m depth a methane hydrate skin protects the bubbles from dissolution. The fate of the dissolved methane is of interest, where it seems that under normal conditions only a minor fraction reaches the atmosphere and most is converted by aerobic methanotrophs [126, 127]. It is however argued that strong eruptions of MVs can lead to climate change [9]. To contribute to methane in the atmosphere, submarine MVs must have episodes of enhanced activity, so that the plume reaches the sea surface. Such episodicity in periods of eruptions has been shown by direct observations on the Håkon Mosby MV. Also, studies on this MV have shown how mud volcanism, i.e. transport of methane rich mud from large depth, can shape the ecosystem of the MV’s surface. Methane arrives at the surface of MVs mostly dissolved in pore water, and is then saturated. Occasionally fountains of bubbles are found. The release of methane from a MV is not homogenous, because the fluid flow velocities are not constant; it is often highest in the central area of a MV edifice, or where the main vents are situated. The warm fluids flowing upwards from large depths gradually lose heat to the cold surrounding sediments. The resulting temperature gradient causes a higher viscosity

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near the wall of the mud volcano than in the middle of the fluid stream. This leads again to the observed velocity gradient in the pore water upflow. AOM was thought to form an effective biofilter that protects the biosphere from significant methane input [84]. The upflow velocity of pore water determines how effective this filter can be. At very high flow velocities of 3–6 m/yr, as observed in the center of the Håkon Mosby MV, the sulfate penetration is limited to the upper few mm of the sediments and no AOM can occur [57]. Most methane that is dissolved in the upwelling pore water will escape to the water column, and where the warm fluids destabilize hydrates, bubble streams are observed [121]. Only in zones where the flow velocity is reduced to less than 1 m/yr are anaerobic methanotrophic archaea (ANMEs) abundant at a depth of 2–3 cm, and here the AOM rate is so high that all methane is converted [57]. It was however argued that even under calm conditions, in between eruptions, a large fraction may escape the biofilter [128]. The relation between flow velocity and structure of the ecosystem, fueled by the sulfide produced by AOM, requires some further explanation. AOM is only possible in the sediment zone where sulfate and methane overlap. The rate of AOM is determined by methane and sulfate supply into that zone. Where the upflow velocity of the pore water is very high, sulfate cannot penetrate into the sediments. Conversely, when the upflow is very low, the methane supply becomes limiting [57]. This concept is used to explain the ecological structure of the Håkon Mosby MV (󳶳 Fig. 3.9). The center of this very large MV (1 km in diameter) is colonized only by aerobic methane oxidizers and most methane escapes into the water column. The upflow velocity is too high for sulfate penetration and no AOM can develop. Further to the outside of the MV the upflow velocity decreases and sulfate penetration to 3–4 cm bsf is possible [57]. The periphery of the volcano is sealed by methane hydrates and the upflow is blocked. The large areas of hydrates are colonized by vestimentiferan tubeworms (called pogonophorans in 󳶳 Fig. 3.9) that reach up to 70 cm bsf, to the surface of the methane hydrates. The worms pump oxygen- and sulfate-rich seawater down their tubes, supplying the ANME near the methane hydrates with sulfate. The worms are autotrophic, with symbionts growing on sulfide produced by ANME from the hydrates [129]. The worms cannot live in the center of the mud volcano, as they are too weak to pump water downwards against the upflowing pore water and would suffocate. A second hypothesis for the absence of ANMEs and AOM in the central part of the Håkon Mosby was that this area is too often disturbed by eruptions so that the ANMEs cannot colonize the sediments [57]. Whereas the absence of AOM in the central area was explained by the high upflow velocity, the second hypothesis could not be disproven. A critical point of the studies investigating the possible dynamics of MVs is that observations during expeditions are snapshots. Eruptions may well be missed, or ongoing eruptions may not even be recognized by the scientific party that is present. The only method to document episodic events is found in long term observations. The first data of this kind were collected from a 12 m long temperature sensor deployed on

3.4 Mass transfer and fluid cycling fuel ecosystems |

14°42'36''E

14°43'12''E

14°43'48''E

14°44'24''E

14°43'48''E

14°44'24''E

72°0'11''N 72°0'0''N

72°0'0''N

72°0'11''N

72°0'22''N

14°43'12''E

72°0'22''N

14°42'36''E

71

>50% Beggiatoa patches

>50% pogonophorans

>50% Beggiatoa spots

20–50% pogonophorans & smooth mud >50%

>50% Beggiatoa & smooth mud 50%

20–50% Beggiatoa Patches & smooth mud >50%

no pogonophorans

20–50% Beggiatoa & smooth mud >50%

100% structured mud

0–20% Beggiatoa Spots & smooth mud >50%

50–100% structured mud

0–20% Beggiatoa & smooth mud >50%

100% smooth mud

pogonophorans >50% & Beggiatoa −0.7 °C, attributable to mud volcanic activity. The latter are only found 12 m or more from the frame. The lines indicate trails of hot spots, which are interpreted as sediment movement. Clearly sediment movements are episodic

beginning. Thus the absence of ANMEs in the center can be explained by two mechanisms, which act in concert: high porewater upflow velocity and regular habitat destruction by eruptions. In conclusion, the ecological structure reflects the stability of the seafloor of the MV. The center is replaced a few times per year while the hydrates on the periphery, with the 70 cm long tubeworms, probably have a stability of centuries [95]. An important global implication of these findings is that the most active areas of MVs, with the highest supply of methane, are not shielded by the biofilter of AOM and sulfide oxidation. Rather, most of the methane brought to the surface from marine MVs is ejected into the water column [27].

3.5 Conclusion Whereas the general mechanisms of mud volcanism are reasonably well understood, the importance of marine mud volcanoes for global cycling of elements is poorly constrained. MVs form interesting ecosystems based on anaerobic methane oxidation and sulfide oxidation. However, these ‘biological filters’ do not prevent large methane emissions in the water column. Their contribution to global or even regional methane budgets is difficult to determine. Firstly, the real number of active volcanoes is not known, and secondly, mud emission is typically not a constant process. Although recent improvements in seafloor acoustic surveys and increasing use of

74 | 3 Mud volcanoes as dynamic sedimentary phenomena that host marine ecosystems

optical methods and long term recording sensors have led to considerable advances in our knowledge and understanding of mud volcanoes, the extent and frequency of eruptions remain uncertain. During eruptions, emissions increase very drastically. It is likely that most measurements during cruises are performed in periods of inactivity, resulting in underestimated emissions. What triggers an eruption is not fully understood. An eruption may not even be recognized when we look at one, as we need to know the difference from background activity. The only way to deepen our knowledge on the activity of mud volcanoes is long term observations, by autonomous or cabled observatories. Which technique is better depends on economic options and methods used. Cabled observatories have practically unlimited power supply and data transfer, and in principle can be interactive: when activity is observed special instruments can be moved in or remotely controlled mobile platforms can be deployed to the most active areas. Disadvantages of cabled observatories are costs and lack of flexibility. The technology for data storage and transmission, power supply and instrumentation (sensors, lab-on-chip) has drastically improved in the last decades, strongly enhancing the usefulness of autonomous observatories. Ideally, a few different types of mud volcanoes should be selected for a concerted large project, e.g. mud volcanoes driven by tectonics, compaction and/or hydrocarbons. Moreover, microbiological studies resulting in knowledge on growth rates, colonization rates and spatial distributions of key species will help to better understand the spatial and temporal variability of mud and fluid flow as well as the distribution of methane seepage at the scale of entire MVs.

Glossary Accretionary prism sediment wedge scraped off of the subducting plate during the collision between lower and upper plate ANME Anaerobic methane oxidizing archaea. The member of the AOM consortium that oxidizes methane AOM anaerobic oxidation of methane. A microbial process where the electron acceptor is sulfate. The process is thought to be conducted by a consortium of an archaea that oxidizes methane and a sulfate reducing bacteria that passes the electrons to sulfate. Areal conversion rate process rate expressed per sediment surface area (mol m−2 s−1 ) Autotrophy production of biomass from CO2 or methane. It is primary production that forms the basis of ecosystems, where other (heterotrophic) organisms thrive by degrading this formed biomass. Energy for biomass production is, in the deep sea, supplied by oxidation of inorganics like sulfide (see chemolithoautotrophy). In the photic zone, primary production is mainly driven by photosynthesis: using light to generate biological energy. Backarc part of a subduction zone landward of the volcanic arc

3.5 Conclusion

| 75

Caldera rimmed depression affecting the whole MV edifice, formed due to compaction of sediments by overloading of freshly erupted material and depleting of an underlying mud chamber Chemolithoautotrophy Metabolic regime where the carbon for biomass is obtained by CO2 fixation and the energy is generated by oxidation of inorganic compounds such as ammonium, hydrogen, iron, or sulfide Christmas tree sedimentary structure identified in seismic profiles that resembles a pine tree formed by the interdigitation of mud flows and hemipelagic sediments Clast fragment of a sedimentary rock detached from its original sedimentary unit Fold geological structure created by tectonic forces that deforms the strata in bent surfaces with a varying degree of curvature Forearc basin area of a subduction zone seaward of the volcanic arc; often terrigenous sediments accumulate here Fault fracture through the different layers of a lithospheric plate mainly caused by interplate movements Geotherm temperature gradient in geological strata Geothermometer indicator of the temperature at which a geological or geochemical event occurred. Reliable geothermometers are stable isotopes, mineral phase transformations, fluid inclusions and mineral alterations Gryphon secondary, minor emitting structure forming on the flanks or at the crest of MVs Heavy/light hydrocarbons hydrocarbons with respectively higher or lower viscosity and density, e.g. crude oil vs. methane Hemipelagic muddy sediment deposited in the deep sea close to continental margins due to settling of fine particles, composed of a mix of biogenic and terrigenous material Hydrocarbon cracking endothermal chemical process that breaks heavy hydrocarbons into light hydrocarbons Hydrofracturing (also fracking) drilling in rock that is fractured by a pressurized liquid Hydrothermal fluid circulation movements of fluids in the sediments driven by the presence of a heat source nearby Illite A nonexpanding crystalline clay mineral belonging to the phyllosilicates group. Illite is a secondary mineral which, in marine sediments, is often formed after dehydration of smectite. Illite crystallinity measure of X-ray angles and intensity of reflection upon the illite mineral, which provides information about the metamorphic facies and pressure/ temperature of formation of the mineral itself Marine subsidence sinking of the seafloor surface due to compaction of sediments in response to increasing load (burial) Mineral dehydration inorganic chemical reaction where water is subtracted from the crystal structure of a mineral forming another one

76 | 3 Mud volcanoes as dynamic sedimentary phenomena that host marine ecosystems

Mud breccia main ejecta of MVs, formed by a clay matrix of varying viscosity within rock pieces; sediment clasts of different provenance and characteristics are included Mud chamber relatively shallow reservoir where an amount of overpressured sediment is stored ODP Ocean Drilling Program Outcrop subaerial exposure of bare rock Overpressured sediments sediments having porosities higher than predicted from compaction law under a defined lithostatic load Paroxysm set of events that constitute the maximum intensity phase of a MV’s eruptive period Underthrusting characteristic of the subducting plate sliding beneath the upper plate Salt diapir geological structure in the form of a plume intruding into the overlying strata by pushing them upwards or piercing them due to overpressured conditions Salse satellite crater filled by a low viscosity pool of bubbling mud and gases Seafloor backscatter signal received from an echosounder or a sonar representing the seafloor interaction with the emitted sound waves. Different signal intensities, if supported by ground samples, can provide a rough surficial sediments classification Sedimentation rate rate of sediment accumulation on the seafloor in a specific sedimentary environment Slime jet A molecular motor for locomotion of filamentous bacteria (Beggiatoa and Cyanobacteria). A row of very small pores, spiraling around the filaments along their axis, excrete slime at an angle. The viscosity of the slime causes gliding motion in the opposite direction. The slime jets are coordinated, as the gliding motion can suddenly be reversed by switching the angle of the motor Smectite An expandable, water rich clay. It is the main source of deep water in mud volcanoes. The most prominent representative may be the phyllosilicate montmorillonite, which precipitates from water as microscopic crystals e.g. during weathering/alteration of basaltic rocks. When dehydrated under high p-T it forms illite Subduction zone region of convergence between two lithospheric plates, with the lower plate being subducted below the upper plate Sulfur disproportionation A metabolic process to generate energy by splitting one S-containing compound into sulfide and sulfate. This can use elemental sulfur, sulfite, or thiosulfate to produce both hydrogen sulfide and sulfate Symbiosis a close interaction between two different biological species, to their mutual benefit. Near seeps, many symbiotic relations develop between bacteria and animals, where autotrophic bacteria are housed inside or on top of animals that harvest the produced biomass

References | 77

Vitrinite one of the main organic components of coal, derived from cell wall material or woody tissue of plants. Its chemical composition comprises polymers, cellulose and lignin Vitrinite reflectance estimation of the maximum temperature history of vitrinite (maturity) through the measurement of percentage of reflected light on the latter. The vitrinite glass-like gloss increases with increasing degree of coalification. Therefore, this measurement can be used to estimate maximum p-T conditions.

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Solveig I. Bühring and Stefan M. Sievert

4 The shallow submarine hot vent system off Milos (Greece) – a natural laboratory for the study of hydrothermal geomicrobiology 4.1 Abstract Hydrothermal vent systems are increasingly recognized as extreme environments of tremendous scientific importance. Their study is of great significance to geochemistry, geobiology and microbiology, because they act as windows into the Earth and its multiple linkages between the bio- and geosphere. Ever since their discovery 40 years ago, deep sea hydrothermal vents have captured the imagination of scientists and the public alike, resulting in many groundbreaking discoveries. In contrast, shallow marine hydrothermal vents have received less attention compared to their deep sea counterparts and are consequently vastly understudied. This is despite the fact that investigations of marine shallow water hydrothermal vents present enormous opportunities, because they (1) provide a window into a microbiome characterized by high metabolic versatility and unique adaptations to extreme conditions; (2) directly influence processes in the surface ocean and finally, (3) combine easy accessibility with a wide range of geological, chemical, and biological processes similar to those taking place at deep sea vents. The shallow water hydrothermal vent system off the volcanic Greek island Milos (Cyclades, Aegean Sea) is characterized by predictable steep geochemical and thermal gradients and fluids loaded with toxic chemicals like arsenite, making it ideally suited to study the staggering variety of microbes inhabiting the sediments and their adaptations to these extreme conditions. In the present review, we provide an overview of the current knowledge of the geology, biogeochemistry and geomicrobiology of the shallow submarine hydrothermal system off Milos and propose this unique system as a natural laboratory to study the interactions between the geosphere and biosphere.

4.2 Introduction 4.2.1 Background Microbes are by far the most abundant, most diverse and widespread lifeform on Earth, due to their ability to resist extreme conditions and because they harbor a vast metabolic versatility. Despite their crucial role in major biogeochemical cycles, our understanding of their diversity and function in extreme environments is still https://doi.org/10.1515/9783110493672-004

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limited. One reason is the inaccessibility of many extreme environments, like deep sea hydrothermal vent systems which, despite recent advances, hampers systematic investigations of the linkages between geochemistry and microbiology. The historic discovery of deep sea hydrothermal vents at the Galapagos Rift took place in 1977 [1]. During this expedition, the scientific team came across cracks in the ocean floor emitting warm, mineral rich fluids, colonized by rich and diverse animal communities, later shown to be supported by chemosynthetic microorganisms using the chemical energy stored in the reduced chemicals generated deep within the Earth [2–4]. Microorganisms gain energy from these reduced chemicals by coupling their oxidation with a terminal electron acceptor, and use it to assimilate inorganic carbon into biomass. Studies of extremophiles from these hydrothermal vents and other extreme environments have shown that life on Earth is far more diverse, widespread and resistant to ‘hostile’ environmental conditions than previously thought possible. Microfossils have been observed in 3.2 Ga-old sulfide deposits, making hydrothermal systems one of the oldest continuously existing ecosystems on Earth [5]. Thus, the extant organisms in these systems are living records of the changes that occurred over geological time, putting these environments into the spotlight for the study of the origin and early evolution of life, and positioning them as potential analogs for extraterrestrial environments.

4.2.2 Shallow marine systems: overview Hydrothermal vent systems in shallow waters represent a ‘high energy’ environment, where microbial metabolism can be fueled by different energy sources: geothermally generated reducing power and sunlight. Hydrothermalism in deep and shallow waters are both driven by oxygenated seawater entering the seafloor through cracks in the crust in tectonically or volcanically active areas. Subsequently, the seawater gets heated up and through seawater-rock interactions, transformed into a reduced hydrothermal fluid, the composition of which depends on the geological setting and the depth of the reaction zone (󳶳 Fig. 4.1). In the deep sea, the enormous pressure allows fluids of up to 400 °C to be emitted at the seafloor (e.g. [6]). Metal sulfides precipitate when they come in contact with the oxygenated and cold seawater, forming the well-known black smoker chimneys. Deep sea vents are usually inhabited by a rich, endemic fauna living in obligate symbiosis with chemoautotrophic microbes [7]. The degree of the dependence of fauna on symbiotic associations at hydrothermal vents changes at the depth of approximately 200 m [8], due to the availability of light and larger amounts of organic matter, eliminating the need for chemoautotrophy to provide the nutrition to the host. In many shallow vent systems, the hydrothermal fluids diffuse upwards through sediments, allowing the evolution of multiple redox interfaces, with distinct ecological niches, all within the known temperature limit of life [9, 10], because the maximum temperatures within the shallow water vents usu-

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Fig. 4.1: The main differences between hydrothermal venting in deep versus shallow waters

ally do not exceed 120 °C [8]. Marine shallow water hydrothermal systems are characterized by high temperatures, extreme pH, and highly elevated concentrations of substances potentially toxic to most metazoans, such as arsenite or hydrogen sulfide. They are also thought to resemble sites of the ‘origin of life’, due to their unique combination of geochemical conditions together with light and high temperatures [11]. Well-known examples of shallow water hydrothermal systems are found off Papua New Guinea [12–14], the Aeolian Islands [15, 16], the South Aegean [17, 18], the Caribbean [19, 20], and the Kurile–Kamchatka island arcs [21, 22]. Although only recently investigated in greater detail, shallow hydrothermal vents were already known from the Mediterranean Sea a long time ago, when Dumas described the gaseous manifestations in the Caldera of Panarea (Aeolian Islands, Italy) [23]. In general, habitats with temperatures above 60 °C are deprived of eukaryotes, and are usually the realm of archaea and bacteria. While many studies have shown that the diversity of archaea is high in terrestrial and marine hydrothermal systems (e.g. [24–26]), it appears that bacteria dominate the microbial communities in most cases [e.g. [26–30]). However, while shallow water hydrothermal systems have been studied regarding their microbial community composition, few attempts have been made to systematically study the linkages between microbial activity and biogeochemical processes, like carbon cycling and its linkages to other element cycles. Investigations at shallow vent sites can improve our understanding of how microbial communities at deep sea vents function, as similar biogeochemical processes and microbial communities are known to occur in both environments. However, to sample and study deep sea sites at the level of detail possible at shallow vent systems

88 | 4 The shallow submarine hot vent system off Milos (Greece)

would be nearly impossible and certainly much more expensive. Data generated and conclusions reached at shallow sites will help in the construction of a framework, which can lead to hypotheses that might be tested at the less accessible, deep sites. However, at the same time the study of marine shallow water hydrothermal systems has been hampered by researchers having fewer resources at their disposal and by a lack of dedicated funding opportunities. In our opinion, the field of hydrothermal geomicrobiology can advance tremendously using marine shallow water hydrothermal systems as model sites, following multidisciplinary approaches and targeting three major goals as follows:  to determine the drivers of microbial community composition and function along physicochemical gradients,  to assess the response and adaptation of microbes to extreme conditions within a shallow water hydrothermal system, which includes their use as a perturbation system analog for climate change, ‘ to design and implement novel technologies to study key processes that might also prevail in deep sea vent habitats. The expected achievements arising from research addressing these goals have the potential to transform our knowledge from largely isolated monodisciplinary results towards a global understanding of microbial life strategies and evolution. One exceptional opportunity to study ecosystem response to elevated temperatures and high concentrations of toxic chemicals is the shallow water vent system off the volcanic Greek island Milos (Cyclades, Aegean Sea).

4.3 Milos – an extreme environment as a model system The submarine hydrothermal vent areas off Milos are widely distributed around the island [17], with the most active venting recorded for Palaeochori Bay and Spathi Bay (󳶳 Fig. 4.2b [31]). Several aspects of the hydrothermal system in Palaeochori Bay have been intensively studied as part of several multidisciplinary projects, including its geology, seismology, (bio)geochemistry, microbiology, and biodiversity, making this site arguably one of the best-studied marine shallow water hydrothermal vents in the world (e.g. [17, 32–39]). The marine hydrothermal field off the coast of Milos (Greece) is fed by hot, reduced fluids from a larger reservoir, percolating upwards through fissures. This creates defined venting zones in the sandy sediment with well-defined thermal gradients in two dimensions. This provides an ideal natural laboratory to test how temperature and other factors associated with venting drive changes in microbial diversity and biogeochemical processes.

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Fig. 4.2: Areal map of Milos within the Eastern Mediterranean Sea (a) and the two main study sites of hydrothermal activity in the southeastern tip of Milos (b) (from Google Maps)

4.3.1 Geophysical and geochemical conditions Volcanism in the Hellenic Volcanic arc is generated by the subduction of the African Plate beneath the Aegean microplate. The volcanism started during the Early to Middle Pliocene, while the last eruption occurred 90,000 years ago. The high seismic activity in the area is associated with important geothermal gas venting, the major systems being found in relatively shallow waters (1–300 m depth) along the Volcanic Arc at Methana, Milos, Santorini and Kos [40]. Due to continuous tectonic activity, volcanism is still present and subsequently, hydrothermalism occurs onshore as well as in the shallow waters off the coast of Milos. The strong geothermal gradient in the area creates a robust hydrothermal circulation with a large amount of CO2 generated from degassing processes in the subducted slab coupled with mantle and magmatic contributions and thermal decomposition of the marine carbonates entrained during subduction ( [31] and references therein). The δ2 H and δ18 O isotopic compositions of the fluids are far off from the estimated Mediterranean Sea meteoric water line, indicating that the hydrothermal fluids are predominantly fed by recirculating seawater (󳶳 Fig. 4.3, [39]). The SE coast of the island of Milos in the middle of the Volcanic Arc with some 35 km2 of geothermally active seabed is characterized by high temperature fluids, reaching over 115 °C [35] and fluctuating in response to tidal frequencies [41]. The average gas composition for the vents is 80.5% CO2 , 1.2% H2 S, 0.8% CH4 and 0.4% H2 [40]. The stable carbon isotopic composition of the CO2 (near 0‰) indicates an inorganic carbon source (dissociation of underlying marine carbonates) [42]. The total estimates for CO2 emissions from the area are from 9 to 35 Mt/year, which is considered to be influential for the global geochemical cycle [33, 40]. The most intense submarine venting off Milos occurs in Palaeochori Bay and Spathi Bay (󳶳 Fig. 4.2b), where hydrothermal fluids and gases escape through the marine and volcanoclastic sediments of varying thickness, consisting predominantly of sand with variable contributions of mud and gravel [33].

90 | 4 The shallow submarine hot vent system off Milos (Greece)

Fig. 4.3: Diagram of δ 2 H and δ 18 O isotopic values of shallow hydrothermal fluids (yellow) and surface seawater (blue) samples. Both δ 18 O and δ 2 H are expressed relative to Vienna Standard Mean Ocean Water. Figure adapted from Gomez-Saez [39], Mediterranean Sea meteoric water line from Price et al. [37]

The vents discharge acidic (pH ≈ 5), hot (up to 115 °C) fluids, which are highly sulfidic (up to 3 mM H2 S) [35, 37, 43]. Two distinct vent types are reported in the sediment hosted system at Palaeochori Bay, Milos: a high-Cl fluid (enriched by up to 47% Cl relative to seawater), which is depleted in Mg and SO4 but enriched in Na, Ca, K, Cl, B, Br and SiO2 ; and a low-Cl fluid (depleted in Cl by up to 66% relative to seawater) that is enriched in Mg and SO4 but depleted in Na, Ca, K, Cl, B and Br [35, 37]. Price et al. reported extremely high concentrations of arsenic in the fluids, with concentrations reaching 78 µM [37]. This is equal to 3,000 times the concentration in local seawater and is moreover the highest value reported for any marine hydrothermal system so far. The formation of brine and vapor phase enriched fluids is constrained by subseafloor phase separation processes [17, 44]. Some of the metals in the fluids are biologically essential (such as Fe), while others may be toxic (like Cu). Molecular insights into the dissolved organic matter (DOM) expelled with the vent fluids was provided by Gomez-Saez et al. through the use of Fourier transform ion cyclotron resonance mass spectrometry (FT-ICR-MS) [39]. The authors studied the sources and fate of dissolved organic sulfur (DOS), revealing that 93% of all assigned DOM formulas exclusively present in the fluids contained sulfur. This study shows extensive hydrothermal reworking of OM in the system, and that hydrothermal systems can act as significant sources of sulfur-containing organic carbon compounds to the ocean.

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4.3.2 Seabed features The hydrothermal activity leads to the occurrence of a distinct zonation on the sediment surface, with yellow-orange patches in the center surrounded by white and brown areas, often occurring in concentric circles of about 2 m in diameter ([32]; 󳶳 Fig. 4.4). The white mats display the largest area and are characterized by midrange temperatures (∼45–85 °C). Yücel et al. showed that these areas are rich in dissolved Fe2+ and Mn2+ and free sulfide using multianalyte voltammetric profiling in retrieved sediment cores [38]. The white patches are predominantly made up of sulfur (up to 50–80%) and silica, and presumably formed by microbial processes [32, 33, 45]. In the lower temperature range (∼30–35 °C) on the outer rim of the white patches, browncolored manganese and iron oxide deposits are present. These areas are characterized by higher signals of FeSaq whereas the concentrations of Fe2+ and Mn2+ were lower than in the white areas [38]. The hottest areas (≥85 °C) are characterized by yelloworange precipitates, which are mainly composed of amorphous orpiment-like Assulfides (As2 S3 ) [46]. Here the oxygen penetration depth is deeper and the pore water lacks dissolved Fe2+ and Mn2+ [38]. These seafloor features strongly depend on hydrodynamic conditions, being influenced by tides and subseafloor microearthquakes [47, 48]. High frequency temperature series and continuous in situ H2 S measurements with voltammetric sensors showed substantial variability in temperature and total sulfide in the upper sediment layer, indicating intermittent mixing conditions due to tidal and wind forcing [38]. Periods of disappearance of the white mats due to wave action are followed by rapid reformation under calm conditions [32, 38].

Fig. 4.4: Morphological seabed features of the hydrothermal sediments in Palaeochori Bay: (a) underwater photograph of a typical hydrothermal patch; (b) schematic drawing of a patch with the simulated temperature profiles along the gradient (c) [69]

92 | 4 The shallow submarine hot vent system off Milos (Greece)

4.4 General aims in hydrothermal geomicrobiology and how they could be addressed by using Milos as a natural laboratory  to determine the drivers of microbial community composition and function along physicochemical gradients The shallow water hydrothermal vent system of Milos is extreme in many respects, such as high temperature, high salinity fluids, as well as low pH conditions, and therefore has attracted many microbiologists over the last 20 years. Early studies often used cultivation as a tool to isolate and identify thermophilic organisms, such as hyperthermophilic archaea [45, 49, 50]. Two giant sulfur-oxidizing bacteria, i.e. Thioploca and Achromatium volutans, have been morphologically identified during studies conducted in the 1990s [45]. Isolates of the chemolithoautotrophic sulfur-oxidizing bacterium Thiomicrospira spp. were obtained from higher dilutions of a most-probablenumber (MPN) approach, indicating their abundance [51]. In addition, other sulfur oxidizers were isolated, some closely related to symbionts of various metazoan hosts occurring in other systems [52–55]. In addition, the thermophilic sulfate reducer Desulfacinum hydrothermale and the iron reducer Deferrisoma palaeochoriense have been isolated [56, 57]. Previously, a combination of cultivation-dependent and cultivation-independent methods was used to describe the microbial community of Milos, revealing the importance of chemolithoautotrophic sulfur-oxidizing bacteria within the system, with Thiomicrospira spp. among the predominant populations [30, 32, 58]. In addition, Cytophaga-Flavobacterium and Acidobacterium were the most frequently retrieved bacterial groups, indicating the importance of heterotrophic processes fueled by photoand chemosynthetically derived organic matter. The relatively large numbers of iron reducers in the mesophilic zone further indicated that in addition to sulfate reduction, ferric iron reduction is an important pathway for the anaerobic degradation of organic matter at the vent site. Highest total cell numbers were found in a transition zone from the hydrothermally influenced sediment to background conditions, indicating the stimulatory effect of hydrothermal activity on the resident microbial community [32]. Systematic studies using the well-defined physicochemical gradients to investigate changes in bacterial diversity and community composition were first conducted using denaturing gradient gel electrophoresis (DGGE) as a fingerprinting technique [30, 32]. These studies revealed a trend of decreased diversity with increasing temperature and also documented the effects of physical disturbances, such as sediment resuspension due to storm activity, on the bacterial community composition. Recently, these early studies were followed up with more fine-scaled analyses, such as automated ribosomal intergenic spacer analysis (ARISA), providing a rapid estimation of the microbial diversity and community structure based on the variability of the 16S-23S intergenic spacer region. A study by Santi systematically investigated

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Fig. 4.5: Nonmetric multidimensional scaling (NMDS) of the bacterial community between the different sampling sites of Spathi Bay. The Bray–Curtis dissimilarity index was used. Symbols are colored according to the different sites: S1 (black), S2 (red), S3 (green), S4 (blue) and S5 (pink). Open symbols represent the top sediment layers (0–2 cm); all other sediment layers (2–20 cm, in 2 cm depth intervals) are depicted with closed symbols. At S5 amplification of bacterial DNA for deeper sediment layers was not possible [59]

a transect from the vent towards nonhydrothermally influenced sediments in 2 cm depth intervals, and revealed comparable community structures within the white areas (S3 and S4) and clearly separated communities in the brown area (S2) and the background (S1) (󳶳 Fig. 4.5) [59]. Other studies also assessed the diversity by directly sequencing the 16S rRNA, either by cloning or more recently through next generation sequencing approaches, providing information on the identity of the bacteria inhabiting the sediments. Price et al. investigated the microbial diversity at two sites strongly differing in composition of the hydrothermal fluids due to phase separation using 16S rRNA gene clone library analyses [18]. The bacterial and archaeal communities of the surface sediments of the high salinity site were dominated by Epsilonproteobacteria, in particular Arcobacter, and Archaeoglobales and Thermococcus respectively, whereas the low salinity site showed a predominance of Bacteroidetes (Flavobacteria) and Thermoplasmatales in the surface sediments, with a transition to Epsilonproteobacteria and Thermoproteales in the deeper layers. Previously, Arcobacter have also been implicated in the formation of the white mats forming on the sediment surface [32, 60]. A study by Giovanelli et al., again using a transect along a thermal gradient of a vent site in Palaeochori Bay, revealed the overarching importance of Epsilonproteobacteria (60% of all sequences), in particular sequences related to the chemolithoautotrophic genus Sulfurovum [61]. Gammaproteobacteria were found to increase in relative abundance with increasing distance from the vent center. More recently, a detailed study using next generation sequencing of 16S rRNA tags investigated changes in bacterial and archaeal diversity along a thermal gradient at a finer scale, again identifying distinct zones corresponding to differences in community composition for both bacteria and archaea (󳶳 Fig. 4.6). As described previously by Sievert et al. based on fewer data points, temperature is a critical parameter in determining this clustering, in particular for the archaea [32]. The fact that this clustering of microbial communities along the thermal gradient exists despite frequent disturbances of the upper sediment layers is suggestive of rapid growth rates, allowing the communities to quickly establish themselves. Alternatively, it could mean that conditions in the sediments are more stable than currently thought. Epsilonproteobacteria were identified as the dominant microbial community members in the areas more heavily affected by

94 | 4 The shallow submarine hot vent system off Milos (Greece)

Fig. 4.6: Taxonomic composition and a tree based on hierarchical clustering using UPGMA (Unweighted Pair Group Method with Arithmetic mean) of (a) the bacterial community structure and (b) the archaeal community structure based on normalized weighted UniFrac distances. Bootstrap values obtained from 100 iterations are depicted as circles on the nodes (black: > 80%; white: > 50%). Sample names are colored according to the distance from the hydrothermal source. Sediment temperature is represented as a gradient from 20 °C (white) to 83 °C (red). Distances in cm from the center of the vent (0 cm, 50 cm, 100 cm, 150 cm, 200 cm, and 300 cm) and sediment layers (0–2 cm, 2–4 cm, 4–6 cm, 6–8 cm, and 8–10 cm) are provided. 300 cm corresponds to the ambient sediment with a temperature of around 20 °C

hydrothermal venting, roughly corresponding to the area of the white precipitate, and were also predominantly confined to the upper sediment layers. These results confirm the importance of Epsilonproteobacteria for chemoautotrophic production, similar to what has been found at deep sea hydrothermal vents. Here it has been proposed that they maximize overall ecosystem function as a result of their high growth rates, rapid adaptations to changing geochemical conditions, and metabolic versatility [62, 63]. They characteristically exhibit chemolithotrophy, meeting their energy needs by oxidizing reduced sulfur, formate, or hydrogen coupled to the re-

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duction of nitrate or oxygen; conditions that also exist at the Milos hydrothermal system. Building on this work, Sievert et al. (unpublished) recently performed in situ stable isotope probing experiments with 13 C-bicarbonate using a device originally designed for deep sea applications [64] to identify the metabolically active microorganisms carrying out chemoautotrophy. By combining the determination of the uptake of 13 C into microbial lipids with rRNA-based analysis, it was revealed that Epsilonproteobacteria dominate dark carbon fixation (Sievert et al., unpublished). These studies further confirmed the essential role of fluid circulation in driving chemoautrophy, as the restriction of flow by closing the top of the incubation device led to significantly reduced levels of incorporation as compared to leaving them open. Milestones Future work can build on this body of work through an interdisciplinary research network to explore the connections between microbial diversity, function, rates and geochemistry. This would include the application of omic techniques and manipulation experiments to fully realize the potential of the vents in Paleochori Bay as a natural laboratory to study chemoautotrophy at extreme conditions and the linkages between the geosphere and biosphere. 

to assess the response and adaptation of microbes to extreme conditions within a shallow water hydrothermal system, which includes their use as a perturbation system analog for climate change

Submarine hydrothermal vents are known for the extreme geochemical conditions that they impose on their inhabiting microbiota. Important selecting factors are low pH, elevated temperatures and high metal and sulfide concentrations (e.g. [65]); all factors that also impact the indigenous microbial communities of the vent sites in Palaeochori Bay. Microorganisms contained in hydrothermal fluids and sediments mainly belong to the group of thermophiles and hyperthermophiles [66], whose genomes and biomolecules have been adapted to the extreme conditions, and with a general community distribution and structure that is strongly affected by temperature (e.g. [32, 67]), trace metals, oxygen concentration and substrate availability [68]. Generally, the efflux of gas and hot water at hydrothermal vent systems leads to a local heterogeneity of the hydrostatic pressure, which is followed by a convective circulation of seawater through the permeable sediment. Convective and advective flow may enhance biogeochemical processes by rapidly supplying oxidants as well as organic substrates, and efficiently removing metabolic end-products from the flushed layers [68]. Lipids are an important component allowing microorganisms to adapt to extreme conditions. A recent study assessed the relative contribution from archaea and bacte-

96 | 4 The shallow submarine hot vent system off Milos (Greece)

ria using a lipidomic approach along a thermal transect [69] (󳶳 Fig. 4.7). The authors detected a high diversity of bacterial and archaeal lipids, with archaeal lipids generally dominating over bacterial lipids in deeper layers with a stronger hydrothermal influence and thus increased temperature. This is in line with previous work showing an increased proportion of archaea with increasing temperature based on quantifying 16S rRNA, although bacteria were dominant up to 82 °C; the highest temperature were RNA could still be quantified [58]. The permeability of biological membranes increases with temperature, which can affect energy-transducing proton and sodium coupling reactions, for example [70]. Based on intact polar lipids (IPL) archaea were found to dominate over bacteria in sediments with temperatures above 50 °C, reflecting the low permeability of their ether-linked isoprenoidal lipids, which makes them more resistant to elevated temperatures [69]. The reasons for the somewhat contradicting results between lipid based and RNA based approaches studying the relative importance of bacteria versus archaea along a temperature gradient remain to be identified. Besides temperature there are other stressors that impact microbial communities in hydrothermal systems. High metal concentrations and low pH are often associated with the venting fluids. Organisms living in and around shallow water hydrothermal vent sites in the photic zones are exposed to these high metal concentrations, with some of the metals being biologically essential and serving as micronutrients (e.g. Fe) while others, such as Cu may be both essential and toxic. Three basic processes can act on the dissolved metals under hydrothermal conditions: (1) oxidation and precipitation, (2) complexation by inorganic ligands, (3) complexation by organic ligands. Studies at Milos have shown process (1), leading to the formation of amorphous Assulfide precipitates around the vents [46]. A study by Kleint et al. indicated organic complexation of Cu in the hydrothermal vents off Milos [71]. Organisms living in these habitats produce small organic molecules – ligands that are able to form complexes with different metals – to either enhance their bioavailability or to decrease their toxicity. Iron is an essential micronutrient for all marine organisms that has been shown to have the ability to travel thousands of km within a deep sea hydrothermal plume due to dissolved-particulate exchange processes [72]. The transformation and distribution of soluble and complexed iron in shallow vent areas is a major gap in knowledge, which needs to be addressed as these compounds directly impact the photic zone of coastal oceans. Phase separation is a ubiquitous process in hydrothermal systems, creating a large range of salinities. Arsenic partly partitions into the vapour phase, and thus can be enriched in different types of fluids with high and low salinity. Extremely elevated concentrations of arsenic have been reported in fluids venting through sediments in Paleochori Bay [37], providing an exceptional natural laboratory to study the role of microbial communities in driving As-redox chemistry. As(III) oxidizers recently attracted a lot of scientific interest because they mediate an environmentally important process by converting the extremely toxic arsenite to the less noxious arsenate, which is also much more effectively removed than arsenite by coagulation with Fe(III) [73]. This

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Fig. 4.7: Sediment profiles of concentration of total polar lipids (dots) and distribution of archaeal versus nonarchaeal lipids (horizontal bars) in Spathi Bay. Adjacent vertical bars indicate the thermal gradient of each station with upper, middle and bottom sediment temperatures in °C. (n.d. = no data) [69]

98 | 4 The shallow submarine hot vent system off Milos (Greece)

group also includes chemolithoautotrophs, which can gain metabolic energy from this reaction to fix CO2 to form biomass [74–77]. The analyses of the aioA-like functional genes (AFGs) suggested arsenotrophy as an important metabolism in the Milos hydrothermal vent system [18]. Similar processes as for As can be expected for other redox sensitive metals such as Sb, Se, Cr, and Hg, which can also be significantly enriched in shallow vent systems. Carbon dioxide (CO2 ) is the primary greenhouse gas emitted through human activities and the ocean is the major sink of atmospheric CO2 . Ocean chemistry is capable of buffering slow changes in CO2 concentration by carbonate chemistry, however, abrupt changes, like the anthropogenically induced increase in CO2 , can cause changes in ocean pH. There is an emerging need to study natural systems experiencing a gradient in pH, like the hydrothermal system off Milos. This approach is suitable for addressing questions on community and chronic effects, biological acclimatization and even adaptation. There are only a limited number of microbiological investigations, most of which reported the differences in prokaryotic community composition relative to the pH gradients [78–82]. Given its accessibility and the seawater pH range between 8 and 5 [32, 34, 69], the hydrothermal field of Milos provides an excellent opportunity for researchers to study the biogeochemical status of an ecosystem acclimatized to an acidified ocean. Milestones To better assess the consequences of climate change, high temperature and CO2 perturbation, systems like the hydrothermal vents off Milos should be used to develop scenarios for a warmer-than-present world. ‘ to design and implement novel technologies to study key processes that might also prevail in deep sea vent habitats To date, hydrothermal systems occurring in shallow and deep water have been treated as independent, seemingly unrelated entities, with the distinction based on the occurrence of vent obligate fauna, which is mainly restricted to vents deeper than 200 m [8] where no sunlight is available. The hydrothermal system off Milos shows venting activity from shallow to deeper waters, with fumarole sites on the beach [83], sites in the shallow subtidal from 4 to 17 m in Palaeochori and Spathi Bays [37] and with sites in deeper waters (e.g. 110 m, [40]). It is expected that deeper sites exist as well, but they have not been explored yet. Here, we suggest exploring the hydrothermal system along the Hellenic Arc in order to investigate the geochemistry and microbiology of marine hydrothermal systems in a transect from the shallow, near shore photic zone to the deep offshore aphotic zone. It is expected that the transition to obligate symbiotic associations occurs at relatively shallow depths, due to the oligotrophic nature of

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the photic zone of the eastern Mediterranean Sea, supplying low amounts of organic matter to the sediments. Many of the thermophiles from shallow vent systems have furthermore been characterized as facultative or obligate heterotrophs [30, 84]. Heterotrophs gain energy by electron transfer between electron donors (organic carbon) and electron acceptors such as oxygen, nitrate, or oxidized sulfur and iron compounds in oxidation-reduction reactions. Thereby, aerobic as well as anaerobic oxidation of organic carbon compounds has been detected at shallow water hydrothermal vents (e.g. [68]). Energyyielding substrates utilized during heterotrophic thermophilic production at various shallow water hydrothermal vent systems include carboxylic acids (short-chain saturated acids, tricarboxylic acids and amino acids), peptides and carbohydrates (mono-, di- and polysaccharides as well as sugar alcohols; [66, 68]). Even biologically labile organic matter might embody a sufficient carbon source for heterotrophs [85]. Dissolved amino acids (AAs), the building blocks of proteins, are of great interest as they are common sources of carbon as well as nitrogen for microorganisms [66] and are thus turned over most rapidly in practically all organisms [86]. Microbial sulfate reduction (SR) was found to be an important process of organic matter mineralization in the hydrothermal sediments of Milos, using incubation experiments with radioactive 35 S, indicating a natural community that is adapted to low pH values and high pCO2 [43], even though a study by Gilhooly et al. failed to identify an isotopic signal for SR by determining the δ34 S of H2 S, concluding that any biological sulfur cycling within the system is probably masked by abiotic chemical reactions [87]. Despite the knowledge of the important mechanistic coupling of microbes and dissolved organic matter within the microbial loop, no further studies specifically targeted heterotrophs within hydrothermal systems. As a consequence, their role in hydrothermal carbon cycling remains to be elucidated. Milestone To quantify the relative importance of auto- versus heterotrophic carbon assimilation in hydrothermal ecosystems and to assess these processes on a transect from light exposed shallow water systems towards aphotic deep sea systems with symbiotic vent communities. The hydrothermal vents off Milos thus provide an exceptional opportunity to study the essential differences between, and the transition from, submarine venting in shallow near shore environments to deeper waters and the effect of venting on biological processes and diversity.

100 | 4 The shallow submarine hot vent system off Milos (Greece)

4.5 Conclusions and future directions The shallow water hydrothermal vents off the Greek island Milos offer a unique opportunity to study the linkages between hydrothermal venting and microbial diversity and function along well-defined gradients of temperature, salinity and pH (󳶳 Fig. 4.8) as well as concentrations of various elements contained in the fluids, such as arsenic. The system thus provides a unique setting to study adaptations of natural communities to changes in environmental conditions, and could therefore be a useful model system to obtain insights into early Earth conditions as well as to test future scenarios of an ocean that gets warmer or more acidic. Shallow water hydrothermal systems also act as as test beds for the development of new technologies to be used in deep sea systems. We argue that research on marine shallow water vent systems deserves more attention than it has received in the past. At present, funding for research on shallow water marine hydrothermal systems lags behind the less accessible, but probably more charismatic deep sea vent systems, severely limiting the progress that could be made by taking full advantage of the opportunities presenting themselves at shallow submarine hydrothermal systems. Acknowledgment: We thank all the members of previous Milos expeditions that we have been part of, especially our joint expedition in May 2012 (J. P. Amend, D. Giovanelli, S. Häusler, C. Kleint, N. Le Bris, I. Pérez-Rodríguez, T. Pichler, P. Pop Ristova, R. E. Price, M. Sollich, C. Vetriani, D. Foustoukos, M. Yücel), for their collaboration,

MetL MetNP

Fe2+/3+ S2As(III)/As(IV)

rich in FeS

Fig. 4.8: Schematic drawing of the different zones of the Milos hydrothermal system, including microprofiles from Sollich et al. [69]. The 3 cm dashed line indicates the dynamic mixing zone within the hydrothermally influenced sediments [38], which is characterized by the copresence of electron acceptors and donors. (MetL : metals bound to ligands; MetNP : metals in nanoparticulate form)

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help, support, stimulating discussions, and good company. Thanks to the General Directorate of Antiquities and Cultural Heritage in Athens for granting us permission for sampling and processing. Special thanks to Gonzalo V. Gomez-Saez and Miriam Sollich for reproduction of figures from their work and to GCA and Frontiers for agreeing to their use, and to François Thomas for analyzing the 16S rRNA amplicon data shown in 󳶳 Fig. 4.6. SMS is grateful to Jan Küver, Wiebke Ziebis, and Paul Dando for giving him the opportunity to work on the shallow submarine vents in Palaeochori Bay for his PhD, which opened many doors. We also thank Athanasios Godelitsas and Antonios Vichos and the Artemis Deluxe Rooms for logistical support in Athens and on Milos. Funding for writing this chapter and for some of the more recent research described herein was provided through the DFG Emmy Noether Grant BU 2606/1-1 to SIB and NSF grant OCE-1124272 and Investment in Science Funds at WHOI to SMS.

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Arab H, Volker H, Thomm M. Thermococcus aegaeicus sp nov and Staphylothermus hellenicus sp nov., two novel hyperthermophilic archaea isolated from geothermally heated vents off Palaeochori Bay, Milos, Greece. International Journal of Systematic and Evolutionary Microbiology 2000, 50:2101–2108. Brinkhoff T, Sievert SM, Kuever J, Muyzer G. Distribution and diversity of sulfur-oxidizing Thiomicrospira spp. at a shallow water hydrothermal vent in the Aegean Sea (Milos, Greece). Applied and Environmental Microbiology 1999, 65:3843–3849. Sievert SM. Microbial communities at a shallow submarine hydrothermal vent in the Aegean Sea (Milos, Greece). Bremen, Germany: University of Bremen; 1999. Sievert SM, Heidorn T, Kuever J. Halothiobacillus kellyi sp nov., a mesophilic, obligately chemolithoautotrophic, sulfur-oxidizing bacterium isolated from a shallow water hydrothermal vent in the Aegean Sea, and emended description of the genus Halothiobacillus. International Journal of Systematic and Evolutionary Microbiology 2000, 50:1229–1237. Kuever J, Sievert SM, Stevens H, Brinkhoff T, Muyzer G. Microorganisms of the oxidative and reductive part of the sulphur cycle at a shallow water hydrothermal vent in the Aegean Sea (Milos, Greece). Cahiers De Biologie Marine 2002, 43:413–416. Goffredi SK, Waren A, Orphan VJ, Van Dover CL, Vrijenhoek RC. Novel forms of structural integration between microbes and a hydrothermal vent gastropod from the Indian Ocean. Applied and Environmental Microbiology 2004, 70:3082–3090. Sievert SM, Kuever J. Desulfacinum hydrothermale sp nov., a thermophilic, sulfate-reducing bacterium from geothermally heated sediments near Milos Island (Greece). International Journal of Systematic and Evolutionary Microbiology 2000, 50:1239–1246. Pérez-Rodríguez I, Rawls M, Coykendall DK, Foustoukos DI. Deferrisoma palaeochoriense sp. nov., a thermophilic, iron(III)-reducing bacterium from a shallow water hydrothermal vent in the Mediterranean Sea. International Journal of Systematic and Evolutionary Microbiology 2016, 66:830–836. Sievert SM, Ziebis W, Kuever J, Sahm K. Relative abundance of Archaea and Bacteria along a thermal gradient of a shallow water hydrothermal vent quantified by rRNA slot-blot hybridization. Microbiology-Uk 2000, 146:1287–1293. Santi I. Microbial diversity and biogeochemistry of Milos (Aegean Sea, Greece) hydrothermal vents. Bremen, 2013. Sievert SM, Wieringa EBA, Wirsen CO, Taylor CD. Growth and mechanism of filamentoussulfur formation by Candidatus Arcobacter sulfidicus in opposing oxygen-sulfide gradients. Environmental Microbiology 2007, 9:271–276. Giovannelli D, Grosche A, Starovoytov V, Yakimov M, Manini E, Vetriani C. Galenea microaerophila gen. nov., sp. nov., a mesophilic, microaerophilic, chemosynthetic, thiosulfateoxidizing bacterium isolated from a shallow water hydrothermal vent. International Journal of Systematic and Evolutionary Microbiology 2012. Nakagawa S, Takai K. Deep sea vent chemoautotrophs: diversity, biochemistry and ecological significance. FEMS Microbiology Ecology 2008, 65:1–14. Sievert SM, Vetriani C. Chemoautotrophy at Deep Sea Vents Past, Present, and Future. Oceanography 2012, 25:218–233. Wirsen CO, Jannasch HW. Microbial transformations in deep sea sediments: free-vehicle studies. Marine Biology 1986, 91:277–284. Jannasch HW, Mottl MJ. Geomicrobiology of deep sea hydrothermal vents Science 1985, 229:717–725. Svensson E, Skoog A, Amend JP. Concentration and distribution of dissolved amino acids in a shallow hydrothermal system, Vulcano Island (Italy). Organic Geochemistry 2004, 35:1001– 1014.

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Matthew O. Schrenk

5 Life in serpentinite hosted alkaline springs 5.1 Abstract The aqueous alteration of ultramafic rocks is a ubiquitous phenomenon that occurs wherever reducing, iron bearing mineral assemblages and water interact through a process known as serpentinization. Serpentinization produces fluids with unique chemical signatures, including high pH, low oxidation-reduction (redox) potential, and which contain abundant volatile compounds. Fluids are sustained in groundwater aquifers in serpentinite settings, and emerge onto the surface as springs. Microbial life in serpentinization influenced habitats is closely linked to geochemical processes, including physiological adaptations to high pH and the ability to utilize reduced gases such as H2 and CH4 . Although to date very few pure microbial cultures have been obtained from serpentinites, a great deal of information has been gleaned from cultureindependent studies using DNA sequencing approaches. Taken together, these data suggest a functionally common serpentinite microbiome, whose members use the reduced products of serpentinization to drive their metabolism. These data also highlight how site-specific differences in chemistry and tectonics impact the resident microbial communities. Understanding the microbiology of serpentinite habitats will provide insights into the functioning of rock hosted subsurface ecosystems and is potentially important to applied topics such as geological carbon sequestration and landscape management in serpentinization influenced settings.

5.2 Introduction to serpentinization influenced ecosystems Serpentinization is a geochemical process whereby ultramafic rocks, in equilibrium with high-pressure reducing conditions characteristic of the Earth’s upper mantle, are brought into the near surface environment and hydrated by circulating water, creating serpentine minerals and distinctive geochemical characteristics. The reduced compounds released through serpentinization can serve as electron donors or ‘fuels’, to sustain microbial metabolism [1, 2]. Serpentinization processes can impact both surface and subsurface carbon pools by producing methane and other low molecular weight hydrocarbons through Fischer-Tropsch-type (FTT) reactions, microbial chemosynthetic processes, and the precipitation of carbonate minerals. Serpentinization also imposes unique challenges upon resident microbial communities, including convoluted fluid circulation pathways, pH at the upper limits of life, and limiting concentrations of dissolved inorganic carbon.

https://doi.org/10.1515/9783110493672-005

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Fig. 5.1: Projection of the global distribution of serpentinite and ophiolites, showing the ubiquity of sites worldwide. Note that most of the terrestrial sites are associated with continental margins. Reproduced with permission from Oze, et al. [4]

5.2.1 Serpentinization is common Although their parent materials originate in the upper mantle, 5–70 km beneath the Earth’s surface, serpentinization influenced terranes are relatively common near the surface of the modern Earth. They occur as uplifted blocks of faulted material associated with tectonic processes at slow spreading mid-ocean ridge (MOR) spreading centers [3], as ophiolite complexes along continental margins [4], and even as ultramafic intrusions in continental crust [5, 6]. As ultramafic rocks were likely more common early in Earth’s history, before the differentiation of continental and oceanic crust and the onset of plate tectonics, serpentinization influenced processes may have played important roles in the origins and early evolution of life on Earth [7, 8].

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5.2.2 Chemical reactions associated with serpentinization Serpentinization occurs when water reacts with minerals (specifically olivine and pyroxene) in iron rich, ultramafic rocks, producing hydrogen and releasing hydroxyl ions, magnesium and calcium (Reaction 1). Experimental evidence suggests that this reaction is sluggish at low temperatures (< 350 °C) and pressures [1]. Surprisingly, however, there are numerous low temperature sites on the modern Earth that bear the geochemical signatures of active serpentinization (e.g. high pH and low redox potential). This enigma may be partly explained by the presence of heterogeneous catalysts in the reacting rock, but the setting and mechanisms governing active low temperature serpentinization processes remains an area of active investigation [1, 9]. Generalized reaction for serpentinization: Reaction 1:

olivine + pyroxene + H2 O → serpentine ± brucite ± magnetite ± H2

The production of hydrogen is directly linked to the oxidation of ferrous iron in olivine and pyroxene with water. A secondary consequence of serpentinization in some settings is the production of hydrocarbons through FTT syntheses. High concentrations of hydrogen in the presence of carbon (CO2 ) can produce methane. This reaction is favored in the presence of NiFe alloys such as awaruite, which is found in trace amounts in some serpentinites. A range of catalysts have been explored including chromite, magnetite, and Fe-Ni sulfide, but have so far produced only small quantities of methane. To date, none of these studies have completely closed the gap between field observations and experimental results. Production of CH4 in laboratory experiments and the relatively enriched δ13 CH4 signatures of methane from serpentinites have led many to conclude that the methane is formed through abiogenic reactions [10]. The carbon may also originate from “dead” 14 C sources, coincident with 3 He and other tracers of mantle sources [11, 12]. Confusingly, several reports using both phylogenetic and experimental evidence have suggested that both biological methane production and consumption may occur in serpentinite settings [13]. Recent work has suggested that biogenic processes can also contribute to 13 C enriched methane, or that the process can be a combination of methane production and consumption [14, 15]. Recent work using methane isotopologues applied to serpentine settings has begun to make headway in resolving this conundrum by showing that specific aspects of the environmental setting can influence the methane signatures, even in geographically proximal locations [16]. In some submarine serpentinites, small organic acids and short-chain hydrocarbons may also be produced by FTT reactions [12, 17]. Formate is the most commonly detected molecule, with the most conclusive abiotic origin [17]. However, hydrocarbons have also been determined to be enriched in ultramafic influenced hydrothermal sites along the Mid-Atlantic Ridge, such as the Rainbow vent field, which may also be associated with abiogenic processes [18]. The dissolved organic matter content in downwelling water and its alteration must also be considered. The circulation

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and alteration of seawater derived organic matter through the Atlantis Massif was explored using samples from Integrated Ocean Drilling Program (IODP) site 1309D [19]. The authors concluded that hydrocarbons in the core material, including biomarkers, were the product of seawater circulation through the massif. Raman microspectro-

Fig. 5.2: Images of carbonate precipitates and travertine terraces. (a) The IMAX hydrothermal chimney at the Lost City hydrothermal field, a three-storey high carbonate chimney percolating with 55 °C, alkaline vent fluids (reproduced with permission Kelley, et al., 2007 [61]). (b) Actively depositing travertine in the Samail Ophiolite of Oman (reproduced with permission from Kelemen and Matter, 2008 [23]). (c) Hyperalkaline spring from Oman showing extensive and active carbonate precipitation as hyperalkaline calcium rich groundwater absorbs CO2 from the atmosphere (reproduced with permission from Chavagnac, et al., 2013 [75])

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scopic studies of organic matter distributions associated with hydrogarnets provided evidence for microbial production in the serpentinization influenced subseafloor [20]. A recent report of a fossilized hydrothermal system near the Iberian margin found evidence of 13 C depleted organic carbon and biomarkers at 65 m below the seafloor that bore similarity to that of well-studied carbonate chimneys at the Lost City Hydrothermal Vent Field [21]. Both of these studies indicate the potential for an active subseafloor biosphere associated with serpentinization in marine hydrothermal systems. The serpentinization-influenced subseafloor biosphere is not just constrained to MORs. A recent study of complex organic matter in clasts from serpentine mud volcanoes in the Izu-Bonin-Mariana subduction zone suggests a deep active biosphere 110 m beneath the seafloor [22]. Modeling studies predict that life in deep subduction zone environments may even exist up to 10 km beneath the seafloor. The majority of dissolved CO2 in downwelling waters is scrubbed from the system via the process of circulation, precipitation as calcium carbonate veins, and subsequent feeding of reductive processes. The mixing of serpentinite springs with meteoric water, either in the shallow subsurface or at the surface, can lead to extensive calcium carbonate deposits (or travertines) [23, 24]. Two prominent examples of this phenomenon are the massive carbonate chimneys found at the Lost City hydrothermal field (LCHF) near the Mid-Atlantic Ridge [25], and the Samail Ophiolite in the Sultanate of Oman. These systems sequester CO2 from the atmosphere in substantial quantities over relatively short timescales, as is evident in the films that form at the surface of stagnant pools. Microbial biofilms can also form as nucleation sites or scaffolds for carbonate precipitation. In some cases the organisms may serve to mobilize solid carbonate phases either through direct or indirect processes.

5.2.3 Inorganic geochemistry of serpentinites Strongly alkaline conditions are a hallmark feature of low temperature serpentinization sites [26]. The alkaline pH is driven by titration of dissolved acids by hydroxyl ions that are a product of serpentinization. Hyperalkaline (or ultrabasic) conditions are further driven by near complete removal of dissolved inorganic carbon (in the form of CO−2 3 ) [27]. As a consequence, pH values in the range of 9–12 are commonly reported in continental serpentinites [26], including some of the highest pH values ever recorded in natural environments [5, 28]. As mentioned earlier, the alkalinity of serpentinite fluids, especially in the presence of the dissolved cations Ca2+ and Mg2+ , not only has a strong control upon the availability of dissolved carbonate, but can also impact the speciation and mobility of other metals, even under reducing conditions. Beyond the phases and mobility of carbon compounds, there are important differences in the geochemistry of serpentinite springs that appear to have major influences on resident microbial communities. As many serpentinites originate on the seafloor, dissolved ions in seawater play an important role in rock alteration, potentially pro-

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Tab. 5.1: Comparison of physical-chemical characteristics of two serpentinites in the California Coast Range

Location (Lat., Long.) pH ORP (mV) Conductivity (mS/cm3 ) DIC (µM) DOC (µM) H2 CH4 Cells/ml a

The Cedarsa “GPS1”

CROMOb “CSWold”

N38°37󸀠 14.84󸀠󸀠 , W123°08󸀠 02.13󸀠󸀠 11.9 −700 3 35 170 50.9% by vol. 7.4% by vol. < 10

N38°51󸀠 42.732󸀠󸀠 , W122°24󸀠 50.04󸀠󸀠 9.8 −277.5 11.2 42 93 < 0.003 µM 1983 µM 2 × 104

data from Morrill, et al., 2013 [32]; b data from Crespo-Medina, et al., 2014 [34]

viding oxidants for sulfate-reducing bacteria, and creating lasting biosignatures in the form of sulfur minerals and their isotopic compositions [29, 30]. Reduced sulfur compounds, including HS− and intermediates, can also serve as ‘fuels’ for further redox reactions, particularly at interfaces. As seafloor serpentinites accrete onto continental margins, some of the marine character of these systems may be sustained, including elevated salinity and the availability of buried marine organic matter. As meteoric water circulates through serpentinites and recharges subsurface aquifers, the character can change from highly saline to freshwater. The contrasting character of these sites is evident through comparison of the well-studied serpentinites of the California Coast Range (󳶳 Tab. 5.1). Groundwater from the Cedars is freshwater and extremely reducing, and contains some of the lowest cell concentrations of any groundwater observed [31, 32]. In contrast, groundwater at the Coast Range Ophiolite Microbial Observatory (CROMO) shows a clear influence of seawater in terms of salinity and aqueous geochemistry. These phenomena are reflected in the higher biomass observed at this site and the functional character of resident microbial communities [33, 34]. Dissolved solutes such as sulfate and intermediate sulfur species may serve as alternative electron acceptors in anoxic fluids. Redox chemistry in fluids of varying salinity, and in the presence of dissolved sulfur, could also influence the mobility of trace metals that may be important both in biochemistry and toxicity. Recent work has shown that oxidized iron compounds in the solid phase can serve as oxidants in the serpentinite environment in both dissolved and solid phases. This is especially important where other oxidants such as O2 , NO−3 , or SO2− 4 are unavailable. Chromium and nickel can also be mobilized and are important for environmental remediation strategies. Serpentinites in terms of laterites have been some of the richest sources of these heavy metals [35]. The release of these metals to the environment has been detected in groundwater associated with serpentinites [4]. Serpentinite soils are

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well known for their endemic plant species, and it is likely that heavy-metal resistant microorganisms are important to the sustenance of these plants [36, 37].

5.2.4 Hydrogeology of serpentinite springs Central to the geochemistry, energy, and nutrient availability in serpentinite influenced ecosystems are the time frame and origin of circulating fluids. In submarine systems fluid flow is associated with thermal gradients that serve to generate hydrothermal flow, the most prominent of which are observed near MOR spreading centers. Ultramafic exposures are more common along slow spreading MORs and are associated with tectonic uplift and emplacement of peridotite massifs [3]. Exothermic heat released by serpentinization may contribute to the thermal budget of these sites, although latent heat from uplift and potential magmatic influences are also important contributors [38]. Tectonic uplift is associated with volume expansion in the parent material, which can facilitate fracturing and exposure of fresh, reactive mineral surfaces. Hydrothermal flow associated with seafloor ultramafic rocks can result in high discharge rates and substantial quantities of dissolved H2 and CH4 expelled in the deep ocean, even detectable as plumes above the hydrothermal vent sites [39, 40]. In contrast, at subduction zones serpentine mud volcanoes are the products of the hydration of ultramafic rocks and the density contrast between the serpentinites and host material [22, 28]. In terrestrial settings, serpentinite springs are found in areas of steep elevation changes, where the water table intersects with the land surface, commonly with only minor differences in groundwater temperature. It is likely that subsurface aquifers with variable storativity feed these fluid seeps. At the Santa Elena Ophiolite in Costa Rica, strong seasonal wet-dry cycles lead to short mean residence times of just a few years [41]. Isolated deep fluids have been reported in wells, such as the Cabeço de Vide in Central Portugal, with residence times of thousands of years [6]. It is important to note that the hydrologic tracers provide a different story with regards to the sources and residence times versus macronutrients, such as carbon [42]. Active serpentinization can even be sustained in very old ultramafic materials. The Tablelands Ophiolite in Newfoundland, Canada was emplaced nearly 500 Ma and continues to seep hyperalkaline, gas rich fluids [43]. The permeability of heavily serpentinized rock is sustained in a fractal network, facilitated by volume expansion and fracturing during the serpentinization reaction itself, and carbonate infilling, which serves to seal fractures leading to discontinuous flow pathways. The flow dynamics can be influenced by numerous factors, including the presence of vegetation and well-developed soil horizons, the magnitude and timing of rainfall events and isostatic processes impacting topography [41, 44]. These processes impact mixing between serpentinization influenced fluids and other water sources, as well as the extent to which these processes approach chemical and redox equilibrium.

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5.2.5 Interfaces associated with serpentinization As serpentinization imparts a strong reducing character upon circulating fluids, when these fluids emerge as springs or mix in subsurface aquifers they generate stark physical and chemical gradients. As with other gradient environments, such as high temperature hydrothermal systems, microbial communities thrive at these interfaces. Gradients in temperature, salinity, and redox state can exist and ultimately challenge and sustain the microbial communities. The bioenergetics of microbial processes in both marine and terrestrial serpentinizing systems have been explored. In the marine environment, the energy yield and favored metabolisms in mafic versus ultramafic hosted hydrothermal systems was studied at various degrees of mixing between hydrothermal fluid and seawater [45]. While some of the results were expected, such as the favorability of metabolic pathways involved in hydrogenotrophy, a higher diversity of energetic reactions was favorable in the ultramafic systems. Curiously, the ultramafic systems also sustained conditions that favored and facilitated anabolic pathways, which may make the energetic demands in peridotite hosted systems less expensive [45]. In terrestrial systems, interfacial environments with mixing between deep and shallow fluids sustain a range of metabolisms including hydrogenotrophy and methanotrophy [46, 47]. The relationship between fluid flow, energy availability and microbial load is readily apparent in natural serpentinizing systems. Highly reducing groundwater representative of ‘endmember’ ultrabasic fluid at The Cedars site and the deep Cabeço de Vide aquifer in central Portugal have some of the lowest cellular abundances in groundwater reported to date( 100 t of CH4 were produced per year at that site alone [51]. As described earlier, ongoing work with methane isotopologues is helping to resolve site specific differences in carbon sources and geologic context to better incorporate these data into comprehensive models [16]. One poorly constrained portion of the methane cycle in serpentinites is the extent to which methane is consumed, dampening the net methane flux from serpentinites, akin to gas hydrate ecosystems. Geological processes can also sequester carbon in the form of CO2 precipitated as carbonate minerals, as evident from the extensive veining networks in heavily serpentinized materials [23]. Circulating organic matter from seawater or meteoric waters can contribute to subsurface biogeochemistry and potentially be metabolized by serpentinite hosted organisms [19]. Biological activities superimposed upon these processes can have an impact upon the rate and character of carbon exchange. Chemoautotrophic reactions can utilize the free energy at interfaces between serpentinites and surface waters to drive primary production. However, due to the precipitation of carbonate minerals, these may be limited to situations for which the flux of carbonate is sufficiently fast to exceed the precipitation kinetics. Alternatively, microbial activities at interfaces between the serpentinizing fluids and surface water may allow for the availability of CO2− 3 such as that found in carbonate chimneys at the LCHF and in travertine deposits [47, 52, 53]. Microorganisms can consume organic compounds from both biogenic and abiogenic sources, producing CO2 and other soluble metabolites [43, 54]. Microorganisms can also generate methane from both organic and inorganic sources [16, 54, 55]. Finally, two little discussed carbon compounds in serpentinites are carbon monoxide and organic acids. Carbon monoxide has been detected in serpentinites and CO oxidation and assimilation genes have been detected [33, 57]. The widespread candidate genus Serpentinomonas can use CO as a carbon source [56]. Genetic signatures of carbon monoxide oxidation are evident in metagenomic data from the Tablelands Ophiolite and in the Coast Range Ophiolite [33, 57]. Carbon monoxide assimilation genes are also evident within the metagenomic data corresponding to the Clostridiales [33]. In these systems, CO may be serving as both an energy source and a readily metabolizable source of carbon in the absence of CO2 . Both formate and acetate have been detected in circulating fluids from serpentinizing hydrothermal systems [17]. Formate carries signatures of being abiogenic, whereas acetate appears to be a product of biological processes [17]. These same small organic acids are abundant in terrestrial serpentinizing systems [33]. Whether they are derived from biotic or abiotic sources, these small organic acids may play a critical role in providing soluble carbon to subsurface microbial communities. In fact, recent evidence suggests that acetogenic bacterial taxa may contribute to the budget of soluble small organic acids at these sites, and may serve as the base of these ecosystems [54, 58]. The utilization

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of both inorganic and organic carbon sources is a hallmark of many of the organisms commonly found in serpentinites: versatility is key.

5.3 Where do serpentinization processes occur? As mentioned earlier, the process of serpentinization can take place in a range of different settings, from the exposure of relatively unaltered material (or peridotites) near MOR spreading centers, to the hydration of overriding mantle material in forearc environments associated with subduction, and to their eventual emplacement onto continental margins. At the same time the character of serpentinization changes in terms of alteration temperatures, the salinity of the circulating fluids, the degree of alteration of the host rocks and other factors. Each of these considerations can influence the character of extant microbial communities, their evolution and resulting biogeographic patterns. In the following section, we will discuss the physical, chemical, and biological properties of some of these sites as a means of broadly considering serpentinization influenced systems. A comprehensive review of chemical and physical aspects of serpentinization can be found in an earlier review on the topic [26].

5.3.1 Submarine systems Prior to 2001 there were very few considerations of the microbiology of serpentinites. The microbiology of serpentine soils had been studied from the standpoint of metal hyperaccumulation and plant-microbe interactions in the rhizosphere. However, at deeper sites, and in submarine deposits, little data existed. Studies of sulfide minerals in ocean peridotites showed evidence of 34 S depletion indicating a role for biological sulfate reduction [29]. A survey of the Samail Ophiolite of Oman reported the trace element geochemistry of high pH springs and used enrichment culture approaches to report on readily cultivable taxa [59]. As studies of both the seafloor MOR spreading centers and the subseafloor biosphere accelerated in the late 1990s and early 2000s it was predicted that hydrothermal systems hosted in serpentinite should exist along MORs. However, the discovery of the Lost City hydrothermal field in late 2000 during a survey of the Atlantis Massif fundamentally shifted our view of the diversity of deep sea hydrothermal vent systems [60, 61]. At the Lost City, high pH fluids vent through tall (up to 60 m off of the seafloor) carbonate edifices. Fluids from the LCHF range from pH 9–11 and temperatures from 40–91 °C and contain copious H2 and CH4 (up to 15 and 2 mM respectively) [26, 60, 62]. Microbiological analyses of the chimneys documented the occurrence of biofilms of archaea related to the order Methanosarcinales, which made up a high percentage of the total microbial populations and which had low phylogenetic diversity [2, 62]. Followon studies sought to resolve whether the Lost City Methanosarcinales (LCMS) were

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producing or consuming methane, only to lead to more enigmatic results. Stable isotope tracer experiments with Lost City chimneys showed evidence for both methane production and oxidation, even in the presence of abundant H2 [13]. Sequence based analyses targeting different gene loci (e.g. nifH or mcrA) showed that the biofilms harbored multiple copies of these functional genes. Shotgun metagenomic analyses of biofilms from LCHF showed that the communities contained the highest abundances of transposases known to date, which may play a role in physiological and functional diversity [63]. Compound-specific stable isotope analyses of lipid biomarkers further complicated this scenario, showing an extraordinary 13 C enrichment in biomarkers associated with methanogenic archaea and sulfate-reducing bacteria [64]. These data presumably indicate near complete consumption of available carbon in interior portions of the hydrothermal chimney habitat. A second major functional group at the LCHF concerns the sulfur cycle. The exterior of carbonate chimneys features prolific gammaproteobacterial communities related to the genus Thiomicrospira, similar to other deep sea hydrothermal vent systems [65]. Patterns in sulfate and sulfide occurrence were observed across the LCHF, suggesting the potential for sulfate reduction [11]. Studies by Russian microbiologists confirmed the potential for sulfate reduction by Desulfotomaculum-like bacteria through activity assays and dsrAB sequencing, which were largely consistent with previous 16S rRNA sequencing analyses [66, 67]. Carbon isotopic budgets, when integrated with sequence and lipid data, suggest a cycle of sulfate reduction coupled to abiogenic or “dead” formate and associated production of CO2 , which is subsequently metabolized to methane by LCMS [11]. However, it will be necessary to obtain isolates to confirm these results. A successful drilling expedition to the Atlantis Massif in 2015 (IODP Expedition 357) sampled basement rock near the LCHF and will almost certainly yield new data to improve our understanding of the vent habitats [68]. Fossil Lost City-type systems have been discovered on the Iberian Margin [21] and elsewhere along the Mid-Atlantic Ridge [69], and may be globally abundant along slow spreading MORs [3]. However, because they lack the strong thermal and chemical signatures of high temperature black smoker type systems, low temperature serpentinizing systems are much more difficult to detect. It should be noted that a number of deep sea vent systems exist where high temperature magmatically influenced fluids interact with ultramafic rocks and show chemical signatures of serpentinization. Fluids venting from the Rainbow hydrothermal site along the Mid-Atlantic Ridge exhibited some of the highest volatile concentrations of any hydrothermal system observed to date [70]. The Mid-Cayman Rise in the Caribbean Sea contains two types of hydrothermal systems, a deep site hosted in gabbros, and a shallow site (∼2,000 m) hosted in ultramafic rocks [71]. The serpentinite hosted system contains high abundances of methanogenic archaea, presumably stimulated by the high H2 content of the site [72]. As these sites show a strong high temperature hydrothermal character, a comprehensive discussion of their occurrence is beyond the scope of this review.

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A second major site where serpentinization occurs in deep sea environments is at subduction zones. Volatiles from the subducting lithosphere hydrate the overlying mantle facilitating serpentinization and the ascension of fluids and clasts of serpentinized material to the seafloor [28]. Studies of the Mariana forearc mud volcanoes showed abundant methanogenic archaea related to the order Methanosarcinales. A drilling expedition in late 2016–early 2017 (IODP Expedition 366) aimed to study the subsurface biosphere in the Mariana forearc and will provide new data about the identity and activity of microorganisms in this little-studied serpentinite ecosystem. Recently, microbiological studies of a shallow sea serpentinizing system in New Caledonia in the South Pacific have uncovered fascinating parallels to the LCHF. Carbonate chimneys in the Prony Bay hydrothermal field vent fluids derived from meteoric water occur at depths < 50 m [73, 74]. Onshore and offshore seeps at Prony Bay contained Methanosarcinales related archaeal sequences, with a high degree of similarity to communities at the LCHF and The Cedars [74]. The Prony hydrothermal field (PHF) has characteristics of both terrestrial seeps and submarine springs, possibly representing a transition between the two primary types of serpentinite ecosystem. Interestingly, a study of juvenile carbonate edifices at PHF provided insight into biological succession in serpentinizing hydrothermal systems [58]. Dethiobacter-like organisms were associated with minerals in the young hydrothermal chimneys, followed by Methanosarcinales-like communities in later stages of the chimney growth. These studies point to the potential importance of interactions between microbial populations for the ecology of serpentinizing environments.

5.3.2 Continental serpentinite springs In comparison to the relatively few known marine serpentinite springs, studies of continental systems are widespread and have accelerated in the past decade. The first studies of the microbiology of serpentinite springs were undertaken in the Sultanate of Oman and published in 1987 [59]. The Samail Ophiolite in Oman represents the world’s largest ophiolite complex (> 350 km long × 40 km wide × 5 km thick), and contains a terrific record of the crustal architecture [23]. Fluids emerge through hyperalkaline springs rich in dissolved hydrogen and methane [75]. Due to its lateral extent, carbonation of the Samail Ophiolite is being explored as a means of geologic carbon sequestration [23, 24]. An ongoing project by the International Continental Drilling Program (ICDP) aims to target the recharge, reaction, and discharge zones of the Samail ophiolite and should provide unprecedented insight into geochemical processes and associated subsurface microbiology. The aqueous geochemistry and stable isotope composition of volatiles has been an important topic in the study of serpentine spring since at least the late 1960s. Work by Barnes and O’Neill first documented high pH springs in these settings and presented scenarios for the formation of ultrabasic fluids through active serpentiniza-

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tion [76, 77]. Work at the Zambales Ophiolite in the Philippines showed the possibility of abiogenic methane production [78, 79]. Recent studies of outflow from the Zambales Ophiolite showed an increased abundance of hydrogen-oxidizing bacteria with increasing redox potential, and the presence of methanogenic archaea in travertine deposits in the near surface environment [47]. Ultrabasic fluids emanating from The Cedars peridotite body in northern California are amongst the most extreme of any terrestrial serpentinite [32]. The springs produce highly reducing (−585 to −656 mV), low salinity water at pH values from 11–12. As a consequence, microbial populations in endmember spring fluids at The Cedars harbor exceedingly low cell abundances, in some cases < 10 cells/ml. Ongoing work at The Cedars has also led to the isolation of one of the few cultivars from serpentinites, of the proposed betaproteobacterial genus Serpentinomonas, which grows optimally at pH 11 and is extremely versatile in terms of its utilization of carbon and energy sources [56]. Additionally, work at the Tablelands Ophiolite in Newfoundland, Canada made the most definitive linkages between subsurface geochemical processes and microbial biodiversity and activities. The Tablelands yielded the first culture-independent data from a terrestrial serpentinite spring, and provided insight into both taxonomic and functional diversity through shotgun metagenomics. The work showed an abundance of Betaproteobacteria, particularly in mixing zones [80]. In more endmember ultrabasic portions of the system, taxa related to the genus Erysipelothrix were found to be abundant. Shotgun metagenomic analysis showed genes involved in hydrogen consumption and production, as well as those associated with carbon monoxide metabolism and CO2 fixation [57]. Phylogenetic comparison of these genes to previously obtained sequences from LCHF showed a high degree of similarly, pointing towards an intriguing unity of processes in marine and terrestrial serpentinite systems. The preceding sites represent only a few prominent examples for the numerous continental serpentinite springs being examined through ongoing studies. Work at the Santa Elena Ophiolite in Costa Rica [15, 41], the Tekirova Ophiolite in Turkey [51, 81], the Voltri Massif in Italy [52, 55] and elsewhere encompass a range of geochemical and hydrodynamic conditions that are providing color to the diversity of serpentinite spring ecosystems on the modern Earth.

5.3.3 Wells in serpentinization influenced aquifers In some cases, more direct access to serpentinizing subsurface aquifers has been obtained through existing or newly drilled wells. Through adequate well rehabilitation and flushing, it is expected that in situ conditions within the aquifer can be sampled. Samples from the AC3 well at 130 m depth at Cabeço de Vide, Portugal represented some of the first microbiological studies of serpentinization influenced wells [48]. Low cell abundances of < 100 cells/ml were observed, and cultivation based techniques

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Fig. 5.3: Global distribution of Desulforudis-like sequences in deep subsurface environments. Note the occurrence of similar taxa in both marine (LCHF) and terrestrial serpentinite environments. Reproduced with permission from Jungbluth et al., 2017 [84]

led to the isolation of a range of alkalitolerant organisms. A follow-on study used molecular techniques to sample the same site. An interesting outcome of the study was the stability of populations in the groundwater across multiple time points, and the prevalence of organisms related to methane and sulfur cycling [82]. One of the most abundant organisms in the sequence based study was related to the candidate species Desulforudis audaxviator, also found in deep subsurface environments of South Africa [83]. This same species has now been reported in numerous rock hosted habitats in terrestrial and marine environments [84]. The genetic and physiological differences between these disparate ecosystems will be a fascinating point of further consideration. A drilling project in 2011 sought to install a network of wells for microbiological studies at the McLaughlin Natural Reserve, near Lower Lake, California, USA [85]. This site, termed the Coast Range Ophiolite Microbial Observatory (CROMO), created eight new wells drilled using microbiological validated methods, in addition to four preexisting monitoring wells, which allowed coupled microbiological-geochemical studies. Work at CROMO has identified abundant populations of Serpentinomonas-like organisms, which are enriched near regions of mixing, similar to observations made in previous studies of serpentinite springs [31, 33]. In deeper, more reducing regions, Clostridiales-like organisms were observed that appear to play a role in sulfur cycling

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and potentially acetogenesis. A comprehensive study of microbial community composition at CROMO found that the predominance of key taxa was related to factors such as pH, CH4 , and carbon monoxide concentration [33]. Recently, work in deep wells of the Samail Ophiolite in Oman has provided further information about microbiological processes taking place in deep aquifers [14, 86]. The authors of this work suggest that biological methanogenesis may contribute to the methane flux at the site, although this result is controversial due to the relatively enriched methane stable isotope signature at the site [87]. Ongoing work coupling geochemical and microbiological processes aims to resolve this discrepancy. A follow-on study in Oman looked at microbial community compositions compared to different physical and chemical parameters and found intriguing differences in microbial diversity and distribution in different settings within the subsurface aquifer systems. The highest diversity was found at interfaces between mafic and ultramafic rocks. In the most alkaline portions of the system diversity was significantly reduced [86].

5.3.4 Serpentine cores In contrast to fluid and gas sampling, direct analyses of native, rock associated microbial communities in serpentinite cores are almost nonexistent. A study of a relatively shallow (∼1 m) core from the Leka Ophiolite in Norway found sequences related to known hydrocarbon oxidizers, as well as hydrogen and ammonium oxidizing microorganisms in fracture-coating material from up to 50 m below the land surface [88]. As part of the establishment of CROMO, two cores were collected up to 50 m in depth below the land surface, which provided an interesting transition between the serpentine soil microbial communities, groundwater and the native aquifer. The cores harbor low microbial abundances (∼103 cells per gram) but showed an interesting peak in methanogenic archaea, which was not observed in associated groundwater [33]. Perhaps the subsurface environment sustains niches that allow methanogenesis to persist, either through the sustenance of CO2 in some form, more stable reducing conditions, or greater quantities of organic matter. A drilling project at the Samail Ophiolite in Oman in 2017–2018 aims to obtain core from up to 400 m below the land surface from different portions of the hydrogeological network, providing one of the best opportunities to look at microbial processes within the ‘root zone’ of continental serpentinizing systems. It is important to note that serpentinization is not only found in continental ophiolites, but may impact microbial communities in other areas of the continental subsurface biosphere. Although it is near a continental margin, the Cabeço de Vide aquifer in central Portugal is a serpentinizing ultramafic intrusion [6]. The Outokumpo borehole in Finland shows a distinctive chemistry associated with a particular serpentinization influenced layer at 1,500 m depth and interestingly shows similar species to those observed in the ophiolite system [89]. One of the more fascinating examples of this are the Kirkland Lake and Timiskaming kimberlite pipes in Ontario, Canada [5].

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These kimberlite pipes are igneous rocks that originated during violent, volatile rich events associated with deep subduction of oceanic lithosphere, but have similar mineral chemistry to seafloor peridotites. The alteration of these materials results in high pH, reducing, volatile rich fluids. As the kimberlites are 100 s of kilometers from the continental margins, it is unclear what biogeographic barriers influence the character of their microbial communities in the groundwater.

5.3.5 Serpentine soils In comparison to the relatively recent interest in deep subsurface serpentines, it has long been noted that the weathered products of serpentinization, serpentine soils, harbor unique biological communities. As far back as the 1920s, microbiologists noted the environmental stress imparted by serpentine soils and low microbial abundances [90]. Serpentine adapted plants are found in many settings, and their associated microorganisms may play roles in nutrient acquisition, and Ni and Cr tolerance. These systems have been noted for their remarkable carbon dioxide uptake ability, albeit in short-lived pools [91]. An important area for research in future years will be coupling the study of extreme serpentinite ecosystems with their near surface counterparts to provide a whole system picture of serpentine ecology and its influence upon the fluxes of carbon and greenhouse gases.

5.4 Constraints upon the microbial ecology of serpentinite habitats Despite the diverse settings in which they are found, ranging from the deep sea to the continental shelf, and dispersed throughout the world, there are several important commonalities among the microbial ecologies of serpentinization systems. First, the systems all exist at alkaline pH, commonly in the 10–12 range. Second, they are reducing, attributed to oxidant consumption and the generation of reduced gases. Third, they are enriched in volatiles, either H2 , CH4 , or both [26]. From a microbiological standpoint, there are other commonalties. The systems in both marine and terrestrial environments harbor extremely low taxonomic diversity. For example, biofilms within carbonate chimneys of the LCHF harbored nearly identical archaeal communities (> 99.9% similarity in 16S rRNA genes) in different microhabitats and over multiple years [65, 92]. At the terrestrial sites, there are commonly only two or three major taxa, with clear evidence that diversity increases as pH decreases [33]. Additionally, while not surprising, a substantial proportion of organisms in these systems appear to use the reduced products of serpentinization, containing hydrogenases, for example, and carbon monoxide assimilation genes.

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Fig. 5.4: Alpha diversity of microbial 16S rRNA in serpentinite spring ecosystems. (a) The high temperature, high pH environment in hydrothermal chimneys at Lost City contained populations of nearly identical (> 99.9% similarity in 16S rRNA genes) archaea related to the Methanosarcinales (b) over multiple years and samples (reproduced with permission from Brazelton, et al., 2010 [92]). (c) Studies of continental serpentinite springs have revealed similarly low taxonomic diversity, comprised of a few major taxa at the most extreme sites. Bacterial populations in groundwater at QV1.1 (blue) and CSW1.1 (red) at the Coast Range Ophiolite Microbial Observatory are comprised of only a few major taxa based upon tag sequencing analyses of > 100,000 amplicons (reproduced with permission from Twing, et al., 2017 [33])

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While alpha diversity of microbial communities is low, there are also indications that physiological and metabolic diversity is not clearly represented in the 16S rRNA gene sequence data. The Lost City hydrothermal field biofilms contain abundant transposases, potentially involved in gene duplication and exchange. They contain numerous distinct copies of the nifH operon and mcrA gene, and this functional diversity may allow these seemingly homogenous microbial communities to occupy several roles [13]. Physiological versatility is also reflected in closely related strains of the cosmopolitan genus Serpentinomonas, which can metabolize multiple carbon sources including CO, CO2 , and acetate, and use O2 and thiosulfate as oxidants [56]. From an evolutionary standpoint, the juxtaposition of low taxonomic diversity and high functional diversity presents an interesting case. Diversification of microbial communities is likely constrained by the extreme environmental conditions, but within this niche diversity and complementarity of multiple independent lineages can coexist. In the subsequent sections, we discuss the roles that high pH, low redox potential, and low dissolved inorganic carbon concentrations may play in constraining these ecosystems.

5.4.1 Challenges of hyperalkaliphily A substantial literature exists on microbial adaptation to high pH, particularly in alkaliphilic strains of the genus Bacillus, as a product of exploiting alkaliphiles for biotechnological applications [93, 94]. Several important challenges to life at high pH include its impact upon RNA stability, membrane fluidity and maintenance of a proton motive force. Alkaliphilic adaptations include maintaining a basic cytoplasmic pH, the use of alternative ATP synthases, or integration of carotenoids or other modifications of lipids in cytoplasmic membranes [95]. Particularly in low salinity, freshwater serpentinite springs, the availability of dissolved ions to use as alternatives to H+ in crossmembrane transport processes, as has been observed in alkaline soda lakes, may be severely limited. Substrate-level phosphorylation may also be used as an alternative to cross-membrane electron transport and ATP generation. Interestingly, studies to date have not revealed unique features of the ATP synthases in serpentinite hosted organisms [56]. Coincident with the need for isolation of more organisms from serpentinite springs is the characterization of the molecular underpinnings for their high pH adaptations.

5.4.2 “Plenty to eat, nothing to breathe . . . ” The supply of oxidized chemicals to react with electron donors may be another confounding factor for biological processes in some serpentinizing systems. The most powerful oxidants may be removed from the system during fluid recharge. Alterna-

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tively, in the upflow zones, biological processes may effectively remove the oxidants in interfacial environments. These systems are impacted by the rates and trajectories of fluid recharge. In the absence of O2 alternative electron acceptors may be utilized, as has been observed in other groundwater ecosystems. When available, nitrate may be a potent oxidant, although in the few reported studies of these systems most dissolved nitrogen exists as reduced species, such as ammonium [34]. Recent work in the vigorous Santa Elena Ophiolite in Costa Rica suggested that perhaps anaerobic methane oxidation could be driven by nitrate reduction, although the system, in general, appears to be oxidant limited [15]. In lieu of nitrate, variably oxidized sulfur compounds may play important roles. Sulfate derived from seawater is present in some systems in nonlimiting quantities. In marine systems, with high flow rates and copious dissolved sulfate, conventional sulfate reduction is believed to play a role [11]. Work from continental serpentinites also suggests that intermediate sulfur compounds such as thiosulfate or polysulfides could play an important role as oxidants [34]. Such processes may be important in serpentine seeps where sulfur is provided by paleoseawater, reaction with buried sulfide minerals, or the dissolution of evaporite deposits [6, 30]. Clear evidence of serpentinization associated sulfate reduction is observed in marine serpentinites like the LCHF [29, 67] and even in obducted ophiolites [30]. At the Eh-pH range of terrestrial serpentinites, intermediate sulfur species may be an abundant and effective resource to support redox processes. Work by Crespo-Medina et al., showed that fluid from CROMO, which are amongst the most saline of all terrestrial serpentinites, showed an abundance of Dethiobacter-like organisms when thiosulfate was added [34]. Dethiobacter-like sequences were also observed in sulfur rich springs at CVA and the hybrid marine-terrestrial system of the Prony hydrothermal field [58]. Although detailed physiological studies of these organisms have not yet been reported, it is likely that they use small organic acids, and potentially carbon monoxide, while reducing intermediate sulfur compounds. The microbes may also use alternative electron acceptors such as oxidized iron. Iron can be potentially stabilized in colloidal forms or present in a dissolved form in dilute quantities [46]. More broadly, solid phases may allow greater versatility as oxidants, since diverse oxidized iron minerals are a product of serpentinization. Recent work at The Cedars site led to enrichment of Pseudobacillus strains using artificial electrodes [96]. These isolates were subsequently shown to couple growth to the reduction of magnetite, a common product of serpentinization, albeit at moderate pH values (∼9). Overall, the coupling of reduced compounds in serpentinites to the reduction of solid iron phases may be much more widespread than previously appreciated. Ultimately, water chemistry may dictate the availability of these ions. Interfacial environments with mixing of surface and serpentinite endmember fluids harbor relatively abundant microbial communities, forming visible macroscopic biofilms in some cases. In less vigorous systems, alterative electron acceptors play more predominant roles.

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5.4.3 Carbon availability Absent from the previous discussion is mention of the use of carbon dioxide as an oxidant. It is likely that dissolved inorganic carbon (DIC) availability plays a particularly significant role in limiting biological productivity in serpentinite systems. The enormous quantities of CO2 precipitated in these systems are evident in the serpentinite ‘stockwork’ observed in serpentinized peridotites, in the carbonate towers at Lost City, and the travertine terraces in Oman and the Cedars. Endmember fluids are almost completely DIC free. The availability of DIC ia likely a strong control on autotrophy in serpentinite systems. One physiological solution to the problem may be to mobilize carbonate through the production of acids to locally acidify environmental conditions, or to form biofilms to directly mobilize carbon from the solid substrates, as has been suggested recently in the study of novel Serpentinomonas isolates [56]. An interesting alternative to either of the above is the direct microbial utilization of dissolved abiogenic carbon compounds. Copious methane is formed by serpentinization, a majority of which appears to be abiogenic (either thermogenic or formed through FTT reactions). Short-chain hydrocarbons have been observed in marine settings at both the Rainbow and Lost City hydrothermal fields [97]. Work at CROMO has shown that methane and carbon monoxide are highly correlated with the most abundant organisms, and that organisms are stimulated by the addition of methane [34]. At the Santa Elena Ophiolite in Costa Rica, there is evidence of both biological methanogenesis and methanotrophy in different regimes, possibly stimulated by the dynamic environment. Work with samples from The Cedars showed the potential for methanogenesis, but primarily from organic acid substrates [54]. Even if CH4 is not a favorable carbon source, the incorporation of small organic acids and their byproducts, such as formate, acetate, and other small organic molecules, can be generated in short-term pools through both biogenic and abiogenic processes and used by resident microbial communities [17]. An interesting consideration, which goes back to some of the earliest studies of subsurface lithoautotrophic microbial ecosystems (SLiMEs) is that homoacetogenesis may serve as the base of the rock hosted ecosystem in serpentinites [98, 99]. Acetogenesis, presumably catalyzed by Clostridiales-like organisms variably using CO, CO2 , or organic carbon sources may serve to feed the remainder of the ecosystem, at least in terms of carbon flow. For systems with connectivity to the surface environment, meteoric waters and surface derived organic matter can be transported by soils and plant exudates to sustain subsurface microbial communities. Alternatively, the copious diffuse gas fluxes through serpentine soils may stimulate bacterial production in mixing zones. Finally, marine sedimentary organic matter and its thermogenic products may impact serpentinite hosted microbial communities [16, 32]. In these situations, the complex organic molecules may serve as fermentative substrates that can be metabolized by the ubiquitous Clostridiales in the absence of exogenous oxidants.

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5.5 Are there endemic species in serpentinite hosted systems? The discovery of the LCHF in late 2000, and the application of culture-independent molecular methods to chimney samples shortly thereafter, led to a great deal of curiosity over the role of the predominant methane-metabolizing species, their link to the high, presumably abiogenic methane fluxes, and the interpretation of the confusing stable isotope values of archaeal biomarkers and total organic carbon in the hydrothermal chimneys [64]. The LCHF has been used as an analog to early-Earth type systems and LCMS-like physiologies have been implicated in some origins-of-life scenarios [100]. The relatively recent reports of LCMS-type organisms at the Prony hydrothermal field only serve to further its intriguing placement [74]. However, the LCMS does not appear to carry the strong 14 C signature of the dead subsurface methane at Lost City [12]. Recent work has suggested that sulfate reducing Clostridiales may serve as the base of these hyperalkaline marine hydrothermal ecosystems and be more closely coupled to subsurface serpentinization reactions [11, 17]. The premise is that sulfate-reducing bacteria are the first colonizers during the growth of nascent chimneys, and that they may metabolize small organic acids, particularly formate, to provide CO2 and other soluble metabolites to the LCMS communities [58].

5.5.1 Serpentinomonas, here, there, and everywhere? As sequence based studies of terrestrial serpentinite systems accelerated, it was soon noted that Betaproteobacteria within the Comamonadaceae were prevalent in numerous geographically distinct systems. Initially, due to generally low biomass and technical difficulties in working with these systems, there was some concern that these sequences were the product of laboratory contamination, as they are related to common soil bacteria such as Ralstonia, which have also appeared as a frequent contaminants in laboratory reagents [101]. However, the continued appearance of this taxon and the close phylogenetic relationship of these sequences, across labs, methods, and sites, provided a degree of confidence in their occurrence. Work at the Tablelands Ophiolite related the abundance of Betaproteobacteria to the mixing regime and percentage of ultrabasic water [80]. These studies found that the betaproteobacterial taxon possessed genes related to hydrogenotrophy and carbon monoxide metabolism, consistent with the biogeochemistry of the site [57]. Further work by Suzuki et al. demonstrated the extraordinary adaptations that isolates of Betaproteobacteria strains from serpentinites, tentatively named Serpentinomonas, possessed to enable adaptation to the serpentinite environment [56]. This included optimal growth at pH 11 and up to 12.5, and the ability to use numerous carbon resources, including solid calcium carbonate through direct and indirect mechanisms. The authors also noted nearly identical gene sequences at other serpentinization influenced sites.

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Taken together, these results point towards the cosmopolitan occurrence of Serpentinomonas-like organisms in high pH continental serpentinites across the globe. These taxa appear to be versatile, metabolizing a range of carbon compounds (e.g. acetate, lactate, glucose, CO2 ). Furthermore, both laboratory and field based evidence point towards their occurrence and abundance at interfaces between high pH reducing fluids and surface waters [33, 47, 80]. What is most intriguing is that Serpentinomonaslike sequences have also been reported in high pH industrial waste sites, including the Superfund site in the Lake Calumet region of Illinois, USA [102]. The extent to which this reflects the global transport and distribution of Serpentinomonas between hyperalkaline surface environments versus site-specific adaptations remains to be studied.

5.5.2 What about the Clostridiales? Unlike the Betaproteobacteria, a consistent picture of the various taxa related to the Clostridiales has not yet emerged. In sulfur rich sites, wells such as CROMO and CVA, Dethiobacter-like species and SRB taxa appear to predominate [33, 82]. These sequences may be a consequence of local sulfur rich conditions, and potentially may represent a bridge between terrestrial and marine serpentinizing environments. Desulfotomaculum-like sequences have been reported in marine serpentinites at both LCHF and PHF [58, 65]. However, in freshwater systems, such as The Cedars, taxa related to Chloroflexi and the candidate phylum OD1 are more abundant. Studying the environmental controls upon the occurrence of Clostridiales and their interplay with lithoautotrophic taxa, and particularly with sulfur and carbon biogeochemistry, is a work in progress for microbial ecologists.

5.6 Evidence of activity and function Serpentinizing systems are often enriched in one or both of the volatiles H2 and CH4 . In some systems, it appears that abiotic processes may shunt H2 towards abiotic carbon reduction resulting in higher CH4 /H2 ratios. In other cases, more hydrogen is available. What limits biological methanogenesis? It is enigmatic that so much hydrogen exists in these highly reducing environments, yet there have been few archaea reported in terrestrial serpentinizing systems. Does it have to do with carbon dioxide availability, bioenergetics, or perhaps competition with acetogenic populations? The few terrestrial sites where methanogens have been reported are generally either near surface [52, 53] or highly dynamic [15]. The reality of the situation is that there are likely a spectrum of different scenarios impacting methanogenesis in serpentinite settings, from fluid dynamics, to burial and infiltration of preformed organic matter, to the competition between methanogens and other microflora, particularly acetogens, each of which must be considered on a site-specific basis. [16]

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In most terrestrial serpentinites, there has been little clear evidence to date that methane is being consumed and metabolized by the resident microbial communities. Anaerobic methane-oxidizing archaea (ANME) sequences have been observed at CVA and SEO [15, 82]. However, recent work suggests that ANME may not only consume but also produce methane [103]. Methane stimulated the growth of microcosms from CROMO, but no direct mechanism of carbon transfer was determined [34]. Furthermore, aerobic methanotrophic sequences were discovered in the outflow from the SEO. It is certainly possible that methane could be consumed in aerobic or microaerophilic mixing zones by typical aerobic methane oxidizing taxa. [33] Additionally, it should be noted that most serpentinite fluids harbor remarkably low biomass. Initial results from drilling expeditions, using the most sensitive cell quantification procedures available, revealed some of the lowest biomass yet detected in a subseafloor environment, on the order of 10s of cells per gram of material [68]. Even reports as far back as the 1920s note that serpentinites are extremely replete in biomass [90]. Cell counts in most fluids range from 105 to below 10 cells/ml. From a microbiological standpoint, these are extremely pure and clean fluids. What is limiting the biomass? Is it energy? Is it nutrients? It is likely a combination of multiple factors. For the most part bioenergetic processes appear to be limited by the availability of oxidants. Nutrients may also limit the system, both in terms of DIC availability and recharge rates. Situations that replenish and provide both these conditions are generally significantly higher in biomass. At the same time, there is often evidence of nutrients ‘left on the table’, such as elevated ammonium concentrations, which would typically be a limiting nutrient in other systems [34].

5.7 Relevance to early Earth, deep biosphere and mineral carbonation As serpentinization influenced habitats are common on the modern Earth, they are important to consider in terms of biosphere processes. In particular, because ultramafic rocks originate in the deep Earth, they represent vectors for the transport of carbon and energy from deep reservoirs to the near surface. The transformation of carbon, specifically to methane, and its subsequent assimilation or metabolism is an important part of that process. Because ultramafic rocks constitute the base of the lithosphere and extend into subduction zones, fluid circulation and serpentinization in these systems may ultimately be linked to the base of the Earth’s biosphere [22]. Particularly pertinent are the reactions that contribute to the copious abiogenic synthesis of organic compounds, and their roles in prebiotic chemistry and potential origins-of-life scenarios [7, 104]. Although the artificially synthesized molecules themselves are far from the simplest organisms, the physical-chemical context of alkaline hydrothermal systems have been considered for their role in providing the raw mate-

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rials and the chemical driving force, in the form of gradients, for early biosynthetic processes. Top-down approaches have also been informative in the study of prebiotic chemistry in alkaline hydrothermal systems. The highly cooperative, single-species biofilms of LCMS bear interesting similarities to genetic annealing theories and relationships to an early progenote that have been a foundation of many origins-of-life theories [13, 105]. Although the placement of common taxa from serpentinization influenced systems indicate that they are relatively evolved from the standpoint of SSU rRNA sequences, a recent study using comparative whole genome analyses provided fascinating insights. Comparison across the genomes of more than 2,000 isolates revealed the commonality of several gene sequences related to physiology and metabolisms of the last universal common ancestor (LUCA) [106]. The traits of the LUCA included carbon monoxide utilization, the Wood–Ljungdahl pathway of carbon dioxide assimilation, hydrogenases, and several other key traits that are also found in microorganisms in serpentinizing systems. As more isolates and genomes are obtained from serpentinites, it will be interesting to observe their phylogenetic placement relative to this new framework. The ecology of serpentinizing systems is also relevant in the search for life on other planets and moons. Serpentine minerals have been detected on the surface of Mars [107] and seasonal methane plumes have been reported in Mars’ atmosphere [108]. The geochemistry of Saturn’s moon Enceladus has been modeled, and is believed to have a subsurface ocean with alkaline chemistry influenced by serpentinization [109]. Plumes from Enceladus contain hydrogen, indicating that the raw fuels to support biosynthesis might exist in the subsurface ocean of this icy moon, and be evident in the plumes emitted by hydrothermal activity and cryovolcanism [110]. In both of these cases it is important to delineate the biosignatures of life processes and to carefully distinguish them from their abiotic counterparts, which has proven challenging even in modern Earth ecosystems.

5.7.1 Carbon sequestration in serpentinites Due to the copious volumes of calcium carbonate precipitates, ophiolites have been considered in geoengineering strategies aimed at carbon sequestration. Pilot studies have indicated that > 1 billion metric tons of CO2 can be buried per year in the Samail Ophiolite alone [23], although extensive field tests, engineering, and cost-benefit analyses remain to be conducted. Another study of carbon uptake in serpentine soils found enhanced uptake of carbon by plants in serpentine soils at Jasper Ridge, California under CO2 enriched conditions [91]. However, upon further examination, this carbon existed in mobile short-term carbon pools. Clearly the role of subsurface microbial communities in effecting the mobility and stability of carbon in these systems should be evaluated before any large-scale strategies are attempted. From studies of dynamic

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serpentinizing springs, such as the SEO, it appears that carbon dioxide may play a role in limiting methanogenesis. If CO2 is artificially added, it may stimulate CH4 production, and convert a net sink for greenhouse gases into a much more potent source.

5.8 Summary and future perspectives Within the past five years, the number of studies of serpentinite microbiology has exploded both in marine and terrestrial systems. The work has resulted in both a census of serpentinite hosted organisms and insight into their unique physiological adaptations. Importantly, the work has provided a view of the ecology of serpentinization influenced ecosystems and how they vary in different settings. Although there are some common characteristics, such as high pH, low Eh, and a preponderance of volatile compounds, there are significant variations in salinity, flow rates, and other factors that have imprints on the resident microbial communities. While sequence based studies of serpentinites have progressed, there remains a dearth of cultivars from these systems. An important focus area for microbiologists moving forward should be to learn from sequence based data to devise creative cultivation strategies, as these data are necessary to understand the mechanisms behind microbial adaptations to serpentinizing conditions, and to derive rates of various biogeochemical processes. A continuing challenge in these systems is the disentanglement of biogenic and abiogenic processes. These problems are exacerbated by the low biomass in many serpentinite ecosystems and relatively prolific abiogenic chemosynthetic processes. This puzzle is further complicated by substrate limitations driving overlapping biogenic and abiogenic signatures. A prime example of this challenge are the stable isotope signatures of methane from various serpentinizing systems and their common incongruence with microbiological data. The results of ongoing analyses of samples from drilling projects at the Atlantis Massif (IODP X357), Marianas forearc (IODP X366) and the Samail Ophiolite in Oman (ICDP) will provide important new information about the integration of geological and biological processes, and possibly expand far beyond what has been observed in springs. The focused interdisciplinary analyses of these samples will be critical in testing hypotheses developed from analyses of serpentinizing springs and other near surface samples, and for integrating serpentinizing systems into a comprehensive view of the Earth’s subsurface biosphere. Acknowledgment: Funding to MOS provided through the NASA Astrobiology Institute Rock-Powered Life grant (NAI CAN 7) through the University of Colorado, an Alfred P. Sloan Foundation Research Fellowship in Ocean Sciences, and a small grant from the NSF C-DEBI STC.

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Mingyang Niu, Qianyong Liang, Dong Feng, and Fengping Wang

6 Ecosystems of cold seeps in the South China Sea 6.1 Introduction Cold seeps were first discovered in the San Clemente fault zone at 1,800 m depth in the California Borderland in 1979 [1] and identified at the Florida Escarpment in 1983 [2]. Cold seeps are sites where cold, hydrocarbon rich fluids seep from subsurface reservoirs to the sea floor at specific sites on continental slopes [3]. Cold seeps are often called methane seeps because the seeping fluids are rich in methane. The temperature of the fluids is close to surrounding seawater, distinguishing them from hydrothermal vents, where extremely hot fluids are vented from the chimney in the deep sea floor. Methane gas accumulation is the main source and trigger of methane seeps. Methane seepage is driven by a variety of geological processes, such as plate subduction, salt diapirism, gravity compression and the dissociation of methane hydrates [3, 4]. When the gas pressure is high enough, gas can migrate through existing faults and sediment pores in the gas hydrate stability zone, forming seeps on the seabed [5], or initiate hydrofracturing above the sediment of the dissociation area, prompting upward gas migration with a considerable amount of overpressure produced by dissociation of methane hydrates [6]. Cold seeps are common deep water habitats on geologically active and passive continental margins, where they are often supported by subsurface hydrocarbon reservoirs [7, 8]. The upward hydrocarbon rich seep fluids can sustain some of the richest marine chemosynthetic ecosystems along modern continental margins, which are composed of a mosaic of habitats covering wide ranges of potential physicochemical constraints for organisms (e.g. in temperature, salinity, pH, and oxygen, CO2 , hydrogen sulfide, ammonia and other inorganic volatiles, and hydrocarbon and metal contents) [9]. Cold seeps can form a continuum of habitats that support species with affinities for the ecosystem [9–12]. Since the first discovery of a cold seep community off Louisiana [13], hydrocarbon seepage and its close association with chemosynthetic biological communities has been documented at a number of sea floor sites worldwide [14]. Following geological studies on the distribution of hydrocarbon rich fluids, an increasing number of cold seeps worldwide and in all tectonic contexts have been recognized, such as Monterey Canyon off the coast of California, in the Gulf of Mexico, the Black Sea and the Sea of Japan, and off the coast of Alaska. New seepage sites continue to be discovered. Cold seeps in the South China Sea (SCS) have received increasing scientific attention since their discovery in the early 2000s [15]. More than 30 seeps have been

https://doi.org/10.1515/9783110493672-006

140 | 6 Ecosystems of cold seeps in the South China Sea

Fig. 6.1: Image of the northern South China Sea showing the locations of all known seep sites. Two active seeps sites, Site F (also called Formosa ridge) and Haima are highlighted by red dots (Liang et al., 2017 [51]). Dark blue dots represent sampling sites of studied cores in cold seeps of the northern slope of the SCS

investigated and constant research efforts on these sites have produced an impressive body of information, the majority of which focuses on the geochemical characterization of seeps [16–19]. Ecosystems including the chemosynthetic microbial community, which is at the base of the ecosystem pyramid at seeps, are still at the beginning stage of investigation in the SCS [20, 21]. Here, we summarize available data collected on sediments from seeps located across the northern SCS continental slope (󳶳 Fig. 6.1). The geological structure of the seeps and megafauna ecosystem will only be briefly introduced (a detailed summary is described elsewhere [22–25]); the main part of this review focuses on the microbial ecosystem in the seeps of the SCS, introduces recent research progress and compares the microbial ecosystems at different seeps. The environmental factors that constrain or shape the microbial community, especially those involved in methane metabolism, will also be presented briefly in this review. Finally, the research frontier and focus for future study on microbial ecosystems in the SCS will be discussed.

6.2 Framework geology The SCS is located at the junction of three tectonic plates: the Eurasian, the Pacific and the Indian–Australian. The SCS is a classic representative of western Pacific marginal seas, surrounded by passive continental margins in the west and north and convergent margins in the south and east. It originally developed during the Cenozoic opening of the SCS via continental breakup and subsequent sea floor spreading. The margins of the northern slope of the SCS display a complex, tectonosedimentary framework.

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Submarine slide, diapir, accretionary wedge and polygonal faults spread widely on the northern slope. The northwest (NW) margin of the SCS is a passive tectonic margin forming the boundary with the broad Chinese continental shelf and the eastern margin is an active collisional margin of a prong of the circum-Pacific plate. Locations of seeps discovered in the SCS are Southwest (SW) of Taiwan, Dongsha, Shenhu and west of the northern slope of the SCS (󳶳 Fig. 6.1).

6.2.1 Southwest of Taiwan The island Taiwan is located at the junction of the Ryukyu and Luzon arcs along the western margin of the Philippine Sea (󳶳 Fig. 6.2a). The Taiwan orogeny is caused by an oblique arc-continent collision between the Philippine Sea Plate and the Eurasian Plate, and has been active since the Late Miocene (󳶳 Fig. 6.2a) [26, 27]. The Eurasian continent subducts eastward beneath the Philippine Sea Plate, along the N–S-trending Manila Trench, whereas the Philippine Sea Plate subducts northward beneath the Eurasian plate along the E–W-trending Ryukyu Trench. Off southern Taiwan, subduction of the South China Sea oceanic crust has resulted in the formation of an arc (North Luzon Arc), a forearc basin (North Luzon Trough), and an accretionary prism (Hengchun Ridge and Kaoping Slope; 󳶳 Fig. 6.2b). Formosa ridge (also called Site F) located SW of Taiwan deposits in a passive margin setting with rapid sedimentation rate and normal faults. It was formed by canyon erosion in the northern, western and eastern parts and is known as the most vigorous cold seep area on the SW of Taiwan Island (󳶳 Fig. 6.1). Large carbonates, which are comprised mainly of high-Mg calcite and aragonite, have been found in the Formosa cold seep area [28]. The carbon and oxygen isotopic detection of carbonates reflected dissolution of gas hydrates during their formation [16, 28, 29]. Alive and dead bivalve samples were obtained by the Chinese research vessel R/V Haiyang IV in 2004 and German research vessel F/S Sonne in 2005 [30, 31], and high concentrations of methane were detected in sediments of the seep area [30, 31]. In summer 2013, a large number of carbonate rocks were observed by the Jiaolong manned submersible [32] (󳶳 Fig. 6.3a). According to the results of in situ observations and isotopic analyses of the retrieved authigenic carbonate samples from the most vigorous cold seep system in the Formosa ridge area, Feng and Chen [16] found that methane seepage was initiated at least 10.6 ka ago, and there was enhanced fluid seepage around 6 ka BP with environmental conditions, followed by a decline from 2 kyr BP until today. In addition to rocks, densely packed mussels and galatheid crabs were also observed around the methane seepage area and we will go into details about these chemosynthetic communities in the following parts.

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󳶣 Fig. 6.2: Tectonic setting of Taiwan and its surroundings. The island of Taiwan is situated at the junction of the Ryukyu and Luzon arcs along the western margin of the Philippine Sea. Tectonically, the island of Taiwan consists mainly of the N–S trending mountain belt and its western foreland basin. (a) Bathymetric map showing the distribution of the East China Sea Shelf, Taiwan Strait Shelf and Kaoping Shelf around the island of Taiwan. Note how small the Kaoping Shelf off the southwestern coast of Taiwan is compared to those of the East China Sea and Taiwan Strait. (b) Schematic cross section showing the westwards advancing Taiwan orogen overriding the passive margin of South China Sea. Deep water turbidites derived from the Taiwan orogeny progressively overlap the passive margin succession, representing the underfilled stage of the early phase of foreland basin development (pictures modified from Hoshing Yu et al. [108])

Fig. 6.3: White crabs (Shinkaia crosnier) and mussels (Bathymodiolus plantifrons) at Site F (Formosa) (a) and Haima (b) cold seeps of the South China Sea (by courtesy of Dr. Dong Feng and Dr. Qianyong Liang). Images are ∼1–3 m across

6.2.2 Dongsha area The Dongsha area is situated on the northeastern continental slope of the SCS, close to the area SW of Taiwan (󳶳 Fig. 6.1). In recent years, many geological studies have been carried out in these regions and multidisciplinary approaches were applied, such as seismic methods combined with geochemical detection and drilling. The seismic profiles and geochemical data indicated that the western part of the Dongsha slope area with the deepest Cenozoic depression and minor slumping bodies might be favorable for gas hydrate formation. Bottom simulating reflectors (BSRs) and blanking zones indicating the existence of gas hydrates in the SCS were reported [33–35]. Based on the geochemical data at site 1144 of ODP leg 184, the chloride content in the pore water decreased from 26.9 ml/L to 17.8 ml/L with the depth [37]. The methane concentration reached a maximum in the depth interval 300–600 m below sea floor (bsf [37]). On the basis of these studies, it has been suggested that the Dongsha area may preserve methane hydrate in deep sediments.

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The Jiulong methane reef in the Dongsha area was discovered on buildups of carbonate reefs on several ridges during the joint Chinese-German R/V Sonne Cruise 177 (󳶳 Fig. 6.4) [33]. It was the first direct evidence of past methane seepage. The carbonates were strongly depleted in 13 C and abundant microbial rods and filaments were recognized within the carbonate matrix as well as the aragonite cements [34, 35]. Shells of bivalves including Calyptogena sp., Acharax sp., Conchocele sp., and Bathymodiolus sp. were found in this region, and irregular yellowish and whitish bacterial patches were found on the surface of the chemoherm blocks, lining fractures and on the sediment surface [34].

6.2.3 Shenhu area The Shenhu area is in the middle of the northern slope of the SCS, between Xisha Trough and Dongsha Islands, and located in the Pearl River Mouth Basin and Zhu II Depression (󳶳 Fig. 6.1). It covers an area of 25, 500 km2 , with water depths from 200 to 2,000 m. The thickness of sediments is ∼1,000–7,000 m with 0.46–1.9% organic matter, which contain an unusually high concentration of methane [37]. Based on seismic analysis, it has a shallow gas hydrate stable zone. High resolution seismic investigation showed that most of the bottom simulating reflectors were located 150–350 mbsf. In addition, from drilling samples, the geochemical data showed that the gas hydrate bearing sediments were not very deep (10 mbsf to 43 mbsf) and located above the base of the gas hydrate stability zone. The gas hydrate was mainly composed of methane, which derived from microbial production. This area was already known as a large source of oil and natural gas [38, 39]. Although no active seeps were found in the surficial sediment of the Shenhu area, a large amount of authigenic carbonates and two carbonate chimneys were collected [17, 35]. The geochemical analysis indicated that methane was the predominant carbon source for all carbonates (δ13 C values usually lower than −40‰) [19].

6.2.4 Qiongdongnan Basin The Qiongdongnan Basin located southeast of Hainan province and is one of the potential gas hydrate-bearing basins on the northwestern continental shelf. It is an oil bearing, fault-depression structural basin in which organic rich sediments of 5,000 m in thickness have accumulated since the Cenozoic [40]. Seismic data showed occurrences of high pressure diapirs and gas plumes, faults, and gas springs, all of which indicated that the Qiongdongnan Basin possessed suitable sources and tectonic conditions for gas hydrate formation [41]. The presence of gas hydrates in the Qiongdongnan Basin has been confirmed using geological and geophysical evidence [41–44]. In March 2015, a large active cold seep was discovered for the first time in the Qiongdong-

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nan Basin by the Haima remotely operated vehicle (ROV) (󳶳 Fig. 6.1). Since it was the first time the Haima ROV was used for a marine geological survey, the cold seep was named the ‘Haima Cold Seep’. In the cold seep area, Bathymodiolus platifrons and microbial mats were found in the surficial sediment (󳶳 Fig. 6.3b), and lots of authigenic carbonate samples were collected by the Haima ROV.

6.2.5 Xisha area The Xisha area is located in passive continental margin of the northwest SCS and is another potential gas hydrates area in the northwest SCS (󳶳 Fig. 6.1). Bottom simulating reflectors and geochemical anomalies that indicated the existence of gas hydrates have been recognized in sediments of the Xisha Trough [45–47]. The cold seeps have not been found in the Xisha area yet, but typical pockmarks have been discovered in the western Xisha Uplift [48, 49]. Some pockmarks are still sluggish in activity with methane-bearing fluids weakly seeping from subsurface sediments [48]. Luo et al. found that the pore water from the sediment of the pockmark anomalies of Cl− concentrations and δ18 O values were attributed to gas hydrate dissociation instead of clay mineral dehydration [50]. This pockmark field in the southwestern Xisha Uplift is likely to be a good prospective area for the occurrence of gas hydrates in shallow sediments.

6.3 Macroecology in cold seeps of the northern South China Sea The northern slope of the SCS is a fault depression and passive continental margin with thick sediments, in which a huge amount of organic carbon has been preserved since the Cenozoic [40]. These features of northern slope of the SCS make it favorable for hosting methane hydrates and cold seeps. So far, two active cold seeps, Formosa cold seep and Haima cold seep, have been found in the SCS. In summer 2013, the Chinese manned deep sea submersible Jiaolong observed symbiotic mussels (Bathymodiolus platifrons, Bathymodiolus aduloides), white galatheid crabs (Shinkaia crosnieri) and shrimps (Alvinocaris) (󳶳 Fig. 6.4) and two years later, a large active cold seep site was discovered for the first time in the west of the northern slope of the SCS by the Haima ROV. Tube worms (Paraescarpia echinospica), mussels (Bathymodiolus platifrons) and clams (Calyptogena sp. and Paraescarpia sp.) were also discovered in the seep area. In the Formosa cold seep, deep sea invertebrates were found in the surficial sediment of the area. Two groups of megafauna dominate: the mussel Bathymodiolus platifrons (󳶳 Fig. 6.4a) and the shrimp Shinkaia crosnieri [25] (󳶳 Fig. 6.4b), which comprised a Bathymodiolus platifrons-Shinkaia crosnieri community (󳶳 Fig. 6.3a). Alvinocaris longirostris [52] (󳶳 Fig. 6.4c, d), which was not previously recorded from

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Fig. 6.4: The dominant species of megafauna communities in the Formorsa cold seep area: Bathymodiolus platifrons (a), Shinkaia crosnieri (b), Alvinocaris longirostris (c,d), Lithodes longispina (e) (modified from Li et al. [25])

the South China Sea, was also dominant in this community. The Lithodes longispina (󳶳 Fig. 6.4e) crab species is the top predator in the community and feeds on the bivalve Bathymodiolus platifrons and possibly the anomuran Shinkaia crosnieri [25]. Other arthropod crustaceans in the Bathymodiolus platifrons-Shinkaia crosnieri community include: Munidopsis lauensis, Munidopsis tuberosa, Munidopsis verrilli, Uroptychus jiaolongae, Uroptychus spinulosus, Acanthephyra faxoni, and Globospongicola jiaolongi (󳶳 Fig. 6.5). Among these crustaceans, Uroptychus jiaolongae, Uroptychus spinulosus, and Globospongicola jiaolongi (󳶳 Fig. 6.5d, e, g, i) are newly recorded species in the South China Sea [25]. In the Bathymodiolus platifrons-Shinkaia crosnieri community of the Formosa cold seep, other invertebrates found include corals (Chrysogorgia sp. [25], 󳶳 Fig. 6.6a), sponges (Semperella jiaolongae [53], 󳶳 Fig. 6.6b), which often occurred in the muddy bottom, and worms (Branchipolynoe pettiboneae [54], 󳶳 Fig. 6.6c), which commonly inhabit in the shells of the mussel Bathymodiolus platifrons. In the Haima cold seep area, the composition of megafauna was similar to the species in the Formosa cold seep area. Clams (Calyplogena sp.), tube worms (Paraescarpia sp.) and mussels (Bathymodiolus platifrons) have been found using the Haima ROV [51]. The two major groups, Shinkaia crosnieri and Bathymodiolus platifrons, were not just dominant in cold seeps, but also in hydrothermal vents [55, 56]. A pattern of population differentiation for S. crosnieri and a homogeneity pattern for B. platifrons were found between the Formosa cold seep and other vents [57]. The main reason might

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Fig. 6.5: Species of arthropod crustaceans found in the Formosa cold seep area: Munidopsis lauensis (a), Munidopsis tuberosa (b), Munidopsis verrilli (c,f), Uroptychus jiaolongae (d,g), Uroptychus spinulosus (e), Acanthephyra faxoni (h), Globospongicola jiaolongi (i) (modified from Li et al. [25])

Fig. 6.6: Some invertebrates in the Bathymodiolus platifrons-Shinkaia crosnieri community of the Formosa cold seep: corals (Chrysogorgia sp. (a)), Sponges (Semperella jiaolongae (b)), worms (Branchipolynoe pettiboneae (c)) (modified from Li et al. [25])

be that B. platifrons has a higher degree of tolerance to environmental heterogeneity than S. crosnieri [57]. Deep sea cold seep or vent invertebrates such as Shinkaia crosnieri shrimp and Bathymodiolus mussels are sustained nutritionally by intracellular symbiotic bacteria (endosymbionts) as primary producers [58–60]. The shrimp Shinkaia crosnieri, galatheid crabs and mussels dwell in sulfide rich environments with host epibiotic microbial communities that utilize sulfide and methane as an energy source, whereas

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different species of shrimp and mussel harbor either methanotrophic, or sulfide oxidizers, or both types of symbionts, which provide the nutrition for S. crosnieri. In the Formosa Ridge, aggregation of B. platifrons, which survived in the sediment with methane hydrate, harbored endosymbiotic methanotrophs in their gills [61]. Similarly, S.crosnieri from the Okinawa Trough harbored numerous epibiont sulfur-oxidizing bacteria on their carapace [61]. Bathymodiolid mussels actively acquire thiotrophic and/or methanotrophic Gammaproteobacteria right from their juvenile stage [62]. These bacterial symbionts were maintained in bacteriocytes, a type of hemocytes with specialized cellular compartments for the storage of symbiotic bacteria [62]. Bacteriocytes may occasionally absorb nutrients via phagocytosis [62].

6.4 Microbial community Methane seeps are island-like habitats, harboring distinct microbial communities that occur worldwide but are locally selected by environmental factors [63]. In contrast with extensive works on geochemical characterization of seeps, relatively few studies on microbial ecosystems have been carried out in the SCS.

6.4.1 Formosa Ridge Bacteria in the sediment core from the seep area of the Formosa Ridge, Southwestern of Taiwan, have been investigated. The piston core DSH (555 cm in length; water depth of 3,009 m) was collected from the Formosa cold seep area in summer 2006 (󳶳 Fig. 6.1) [20]. Lithologically, core DSH was composed of silty clay interbedded with turbidite layers. Headspace methane concentration was about 2.1 mM at the top of the core, and increased with depth, reaching above 20.4 mM at the bottom of the core. Sulfate concentration showed a decreasing trend with depth, from 29.1 mM at the top to 12.6 mM at the bottom of the core. Correspondingly, hydrogen sulfide concentration increased from nearly 0 mM above 250 cm to 1.6 mM at 550 cm depth. The chloride concentration was approximately 550 mM above 450 cm and slightly decreased below this depth. TOC content ranged from 0.5% to 0.7% (dry weight). Based on the profiles of geochemical data in pore water, the sulfate-methane transition zone (SMTZ) ranged from 500 to 600 cmbsf in core DSH [20]. Total cell abundance measured with AODC (acridine orange direct count) decreased from 2.3 × 106 cells/g (wet weight) in the top layer to 5 × 105 cells/g (wet weight) at 200 cmbsf, then increased gradually with depth and reached the maximum value of 9.6 × 106 cells/g (wet weight) in the bottom layer. The profile of cell abundance was similar to that of the methane concentration in the core. The total bacterial abundance decreased with depth from 1.9 × 109 copies/g (wet weight) in top layer to 3.2 × 107 copies/g (wet weight) in the bottom layer [20].

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Some sediment subsamples were selected from the whole core DSH to investigate microbial diversity with the 16S rRNA clone library. Proteobacteria (mainly Deltaproteobacteria) was the dominant bacterial group in the surficial layer of the sediment core, and accounted for 17% of the whole bacterial 16S rRNA gene library [20]. Chloroflexi and candidate division JS1 became dominant in the middle (400–405 cmbsf) and the bottom (550–555 cmbsf) layers respectively. Most of the sequences from Chloroflexi were classified into Anaerolineae [64]. The abundance of JS1 and Chloroflexi related bacteria decreased with depth, varying from 7.5×104 copies/g at 300 cmbsf to 3.6×106 copies/g at 50 cmbsf, and accounted for only minor amounts of total bacterial abundance. The overall bacteria structure revealed at the Formosa seep was basically similar to what has been found in other cold seeps, such as at Hydrate Ridge, Guaymas Basin and the Gulf of Mexico [63, 65–67]. The major bacterial groups, Proteobacteria, Bacteroidetes and Chloroflexi, were found in all studied seep sites (󳶳 Fig. 6.7a). Phycisphaerae and Chloroflexi were found in high abundance in the Formosa seep. Moreover, candidate division JS1 was found in high abundance in the Formosa seep, similar to the methane seeps of Hydrate Ridge and the Gulf of Mexico (󳶳 Fig. 6.7a). In addition, most of those sequences that belong to candidate division JS1 and Chloroflexi from cold seeps and gas hydrate regions, and have high sequence identity among each other, seem to be cold seep environmental specific. Thus, these specific groups within the candidate division JS1 and Chloroflexi may have a role in methane metabolizing either directly or indirectly [64, 68]. Besides the clones affiliated with the three predominant phylogenetic groups mentioned above, other clones were affiliated with several minor groups, such as Acidobacteria, Actinobacteria, Chlorobi, Firmicutes, Planctomycetes, Spirochaetes, Verrucomicrobia, OP3, OP8, OP11, TM6 and WS3, and unclassified groups, each group accounting for a low percentage of the total sequences.

6.4.2 Jiulong Reef and Haima Seep In addition to Formosa samples, archaeal composition in the sediment cores from the Jiulong methane reef of the Dongsha area and the newly discovered Haima cold seep in the western part of northern slope of the SCS has been reported. The gravity core CL11 (767 cm in length, water depth of 1,607 m) was collected from the southern slope of the Jiulong methane reef during a cruise of the Chinese research vessel Haiyang IV for geological surveying in the summer of 2012 (󳶳 Fig. 6.1). Two piston cores QDN-14B (∼830 cm in length) and QDN-31B (∼750 cm in length) were retrieved from the active seep area named ROV1 (in the Haima cold seep area) and a nonseep control site, respectively, in spring of 2015 (󳶳 Fig. 6.1). The core CL11 is mainly composed of silt (68.49–78.53%), sand (0–2.06%) and clay (20.85–31.51%). A relatively higher percentage of sand and lower percentage of clay occurred in the upper interval (90–440 cmbsf) of the core. Accordingly, the sedimentary porosities of upper layer (above 550 cmbsf) are higher (> 65%) than those of lower

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󳶣 Fig. 6.7: Relative abundance of archaeal and bacterial classes in the cold seep sediments. The graphs show the relative sequence abundance of the most abundant bacterial classes (a), archaeal phyla (except ANME and methanogens) (b) and subgroups of ANME and methanogens (c), in different cold seep ecosystems (GB, Guaymas Basin hot seeps; GoM, Gulf of Mexico seeps; JS, Japanese Nankai Trough seep; NNS, Northern North Sea; NZ, New Zealand seep; HR, Hydrate Ridge seeps; DS, Dongsha seep; HM, Haima seep)

parts of the core. As for geochemical data of the core, TOC increased with depth, ranging from 0.8% to 1.2%, and methane increased below 600 cmbsf, reaching a peak at 744 cmbsf. According to these geochemical data, a clear SMTZ was observed between 600 and 800 cmbsf in the sediment core [69]. In the piston core QDN-14B retrieved from the active seep site of the Haima cold seep area, numerous gas bubbles were observed in charcoal grey sediment. A strong smell of hydrogen sulfide was recognized during core processing. In contrast, no visible gas bubble or sulfide smell was noticed in the typical yellow-brownish sediment from core QDN-31B collected from the nonactive control site [70]. For QDN-14B, the concentration of sulfate decreased sharply with depth, from 28 mM at 10 cmbsf to under detection limit (below 0.1 mM) at 400 cmbsf. The DIC concentration was generally constant in the upper 240 cmbsf (∼6 mM) and then increased to 27 mM at 480 cmbsf. The δ13 C value of DIC became more negative from −6‰ in the surface sediment to −53‰ at 300 cmbsf and then positively increased to −36‰ below 400 cmbsf. Based on the sulfate, δ13 CDIC and DIC profiles, the SMTZ of QDN-14B was estimated to be ∼300–400 cmbsf. In contrast, the sulfate concentrations varied slightly from 28 mM at the surface to 25 mM at the bottom of the QDN-31B core [70]. Microbial abundance was measured in all three cores. In core CL11, cell abundance based on AODC ranged from 2.13 × 107 to 1.05 × 108 cells/g in the whole core. From the top to bottom of core CL11, the pattern of cell abundance resembled S-shaped profiles. A relatively higher abundance of cells was observed at the depths of 17 cmbsf, 450 cmbsf and 740 cmbsf of core CL11. The archaeal abundance varied from 1.63×107 to 3.75×107 copies/g and had a similar pattern to the bacterial cell abundance [69]. In the QDN-14B core, the archaeal 16S rRNA gene copy number varied between 4.9 × 104 and 6.1 × 106 copies/g (wet weight sediment), while the bacterial 16S rRNA gene copy number varied from 1.9 × 106 to 3 × 107 copies/g (wet weight sediment). The abundance of the bacterial 16S rRNA gene was around two times higher than that of archaea. The ratio of archaeal vs. bacterial cells ranged from 7% to 68%. In the surficial layer, the ratio was more than 60% and then decreased with depth, except at 650 cmbsf, where it was 65%. In the QDN-31B core, the bacterial abundance varied between 2.1×106 and 1.4×107 copies/g (wet weight sediment) and decreased gradually with depth. The archaeal abundance varied between 1.4 × 104 and 1.2 × 106 copies/g (wet weight sediment), and decreased sharply with depth. The ratio of archaeal vs.

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bacterial cells varied between 1% and 25% and decreased with depth. The ratio was much lower than that in the QDN-14B core ([70]). For archaeal diversity analysis, seven sediment subsamples were taken from the top to bottom of core CL11, sixteen subsamples from core QDN-14B and thirteen from control core QDN-31B for investigatation with archaeal 16S-tagged Illumina sequencing methods [69, 70]. These studies revealed that Thermoplasmata of Euryarchaeota (mainly Marine Benthic Group D, MBG-D), Bathyarchaeota, Lokiarchaeota, Thaumarchaeota and Woesearchaeota were the dominant archaeal groups in the sediments. These archaeal groups have varied ecological functions and distributions. The MBG-D group of archaea were cosmopolitan in various marine sediments [71–78], and were suggested as protein degraders since genes encoding for extracellular protein-degrading enzymes were found in the cell genomes [79]. Meanwhile, MBG-D members were found in abundance and showed co-occurrence with methane-oxidizing archaea in cold seep sediments [74] and in anaerobic methane-oxidizing enrichment [80], suggesting a direct and indirect role for the MBG-D group in methane metabolization. Thaumarchaeota were commonly found in the water column and surficial sediments, where they could oxidize ammonia to nitrite aerobically. Recently, lots of research studies have reported that Thaumarchaeota can be found in anoxic sediments and may have the ability to oxidize ammonium in anoxic environments [81]. Bathyarchaeota were widespread in organic rich sediments [71–74, 77, 78, 82–87]. They were initially considered to be heterotrophs by single-cell genomic and metagenomic analyses [88, 89] as well as stable isotope probing experiments [90]. Recently, some lineages of Bathyarchaeota were suggested to be capable of methane metabolization and homoacetogenesis based on genomes obtained from metagenomic analysis [91, 92]. Lokiarchaeota were identified as a novel, deeply rooted clade of the archaeal TACK superphylum. Active Lokiarchaeota were found to dominate the sediment of the studied cores (DSH, QDN-14B, CL11) with high methane concentration, suggesting they may also directly or indirectly be involved in methane metabolization in marine sediments [83, 93]. Woesearchaeota, which were previously named Deep-sea hydrothermal vent Euryarchaeota Group-6, are distributed widely in marine sediments [94–96]. Woesearchaeota were primarily considered to have a symbiotic or parasitic lifestyle due to the fact that most of the core biosynthetic pathways were partial or absent [97]. Methane production and oxidation are generally dominated by archaea within the Euryarchaeota phylum. Methanogens can produce methane in anoxic sediments with simple compounds, such as CO2 and H2 , C1-compounds and acetate. They are initially assigned to six orders: Methanococcales, Methanopyrales, Methanobacteriales, Methanosarcinales, Methanomicrobiales, and Methanocellales [98]. Recently, a seventh order of methanogens, the Methanomassiliicoccales (previously referred to as Methanoplasmatales), was discovered in the termite gut and human faeces [99, 100]. Most members of Methanococcales, Methanopyrales, Methanocellales, Methanobacteriales and Methanomicrobiales are hydrogenotrophs, and produce methane with

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H2 /CO2 [101]. The Methanosarcinales comprise the most versatile species, which have the ability to carry out methanogenesis with various substrates described above, with the genera Methanosarcina and Methanosaeta known to have the capability to convert acetate into methane [102]. The members of Methanomassiliicoccales so far consists exclusively of obligate hydrogen dependent methylotrophs [103]. Anaerobic oxidation of methane (AOM) is considered to be conducted by ANME (anaerobic methan-oxidizing archaea) including ANME-1, ANME-2 and ANME-3 subgroups. Since ANME-1 and ANME-2 consist of phylogenetically and physiologically diverse subgroups, ANME-1 and ANME-2 clades have been further assigned into different subgroups. ANME-2 clade has been refined into two subgroups, ANME-2a/b (previously considered to be two separate subgroups) and ANME-2c; in addition, the ANME-1 clade has been assigned into ANME-1a and ANME-1b subgroups [104]. ANME-2 clade is phylogenetically related to cultivated members of the Methanosarcinales, ANME-1 clade is phylogenetically similar to Methanomicrobiales and Methanosarcinales [105] and ANME-3 clade is most related to Methanococcoides spp. [74]. In the sediment cores (CL11 and DSH) from the Jiulong methane reef and the Formosa cold seep area, only a few sequences affiliated with ANME-1, ANME-2a/b, Methanosarcinaceae and Methanomicrobiaceae were found in the methanogenic zone and surficial layers of sediment cores, respectively [69]. In the Haima cold seep area, methanogens and ANMEs were enriched in the studied sediment core (QDN14B), which include Methanosarcinales, Methanomicrobiales, ANME-2a/b, ANME-2c, ANME-1, and GoM-Arc1 groups (Niu et al., 2017 [70]). ANME-2a/b and ANME-1b were the predominant groups within methane metabolizing microbes in the Haima cold seep, and they showed niche separation in the sediment core. In the sediments of the Haima cold seep, ANME-2a/b were dominant in the upper layer of the SMTZ, and ANME-1b were more prevalent in the sulfated depleted deeper sediments. Fine phylogenetic analysis further divided the ANME-1b group into three subgroups with different distribution patterns: ANME-1b I, ANME-1b II and ANME-1b III. ANME-1b II and ANME-1b III appeared to have different trends: the relative abundance of ANME-1b II decreased gradually with depth, while ANME-1b III increased (Niu et al., 2017 [70]). Niche separation of ANME-2 and ANME-1 in the cold seep sediments has also been noticed in other marine sediments, such as the Eastern Japan Sea offshore area of Joetsu and other cold seeps [89, 106, 107]. Ruff et al. investigated the global distribution of ANME clades in cold seeps and found sediment depth, water depth, sediment temperature, methane and sulfate concentration all had a strong influence on the distribution of ANME clades [63]. A strong correlation between ANME-1 and DIC, and ANME-2a/b and sulfate was further observed in the sediments of the Haima cold seep [70]. More indepth observations and analyses are still needed to understand the niche separation of ANME clades. The archaeal communities of various global cold seeps have been compared (󳶳 Fig. 6.7b). Thermoplasmata were found in all nine cold seeps investigated, and Thaumarchaeota and Lokiarchaeota occured in most of the sediments of cold seeps.

154 | 6 Ecosystems of cold seeps in the South China Sea

Bathyarchaeota were found enriched in the cold seeps of the SCS. The composition of ANME and methanogens showed various patterns in the cold seeps, however, the role ANME-1 and ANME-2 as major methane oxidizers was widespread (󳶳 Fig. 6.7c).

6.5 Outlook Nearly two decades of surveys on the northern continental slope of the SCS have provided a large amount of geological and geochemical data, which have confirmed the presence of various hydrocarbon seeps in the area. However, the ecosystems, including megafauna and microbial community, of the cold seeps in the SCS are only at the beginning stage of investigation. Limited studies on the microbial ecosystems of the SCS have found some endemic ecotypes that may play important roles in methane metabolism. A close coupling of geochemical and microbiological investigation will lay the foundation for uncovering the geochemical and ecological roles of the microbes in the cold seeps. More intensive and in-depth multidisciplinary approaches, including methane transforming rate measurements, cell enrichment and cultivation, stable isotope probing, nanoscale secondary ion mass spectrometry, and metagenomics, metatranscriptomics, metaproteomics, and metametabolics, will shed light on the functional and syntrophic relationship of the organisms within the ecosystem. The cold seeps in the SCS can be considered fascinating natural laboratories and key candidates for ecosystem monitoring. Acknowledgment: This work has been financially supported by the National Natural Science Foundation of China (Grant: 91228201, 91428308 and 41525011) and partially supported by National Special Project on Gas Hydrate of China (Grant: GZH201100311, DD20160217).

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Jian Ding and Yu Zhang

7 Life at the hydrothermal vent field of the Southwest Indian Ridge 7.1 Overview of the hydrothermal vents The ∼ 50,000 km global ridge system, composed of mid-ocean ridges, fracture zones and back-arc basins, is an integral component of the Earth’s plate tectonic system. The system defines the diverging margins of tectonic plates according to the various separating rates: very fast (> 20 cm/yr), fast (8 cm/yr or more), slow (4 cm/yr or less) and ultraslow (< 2 cm/year). The concept of the ultraslow spreading class of ocean ridges was proposed after investigations of the Southwest Indian and Arctic ridges about two decades ago. Since then, ultraslow spreading ridge systems have been discovered in the Gakkel Ridge (0.3–0.7 cm/year) [1], the Southwest Indian Ridge (SWIR) (1.4–1.5 cm/year) [2], the Knipovich Ridge (1.5–1.7 cm/year) [3], the Terceira Rift Ridge (0.23–0.38 cm/year) [4], part of the North Fiji Basin (1.5–2.0 cm/year) [5]and so on. These ultraslow spreading rates are largely due to the influence of ridge geometry. Although the ultraslow spreading ridges were previously considered as part of the slow spreading ridges, it has been realized that the differences between ultraslow and slow spreading ridges are as great as those between slow and fast spreading ridges. The ultraslow spreading class of ocean ridge is characterized by intermittent volcanism and a lack of transform faults, where the mantle beneath is emplaced continuously on the seafloor over large regions [6]. Normally, the ultraslow spreading ridges consist of linked magmatic and amagmatic accretionary ridge segments. These amagmatic segments sometimes coexist with magmatic ridge segments for millions of years to form stable plate boundaries, or they may displace or be displaced by transforms and magmatic ridge segments as spreading rate, mantle thermal structure and ridge geometry change [7, 8]. As in the SWIR, ongoing hydrothermal venting activities have been observed [9–11]. Submarine hydrothermal vents are usually discovered at the ocean ridges, where the crust is thin enough to allow hot fluid to be emitted out of the magma chamber. This fluid from active hydrothermal vents can easily reach temperatures up to 300 °C, until it reaches the seawater and the temperature quickly drops to ambient levels. The chemical composition of a venting fluid is different from the seawater. Depending on the origin of the fluid, it is normally highly reduced and contains relatively high concentrations of minerals, sulfides, hydrogen and even organic carbon. During the cooling process, chimneys built up of carbonate minerals and sulfide minerals form along the venting path and a sharp physical-chemical gradient forms between the venting fluid and seawater. Various biochemical reactions occur from which a spehttps://doi.org/10.1515/9783110493672-007

162 | 7 Life at the hydrothermal vent field of the Southwest Indian Ridge

Fig. 7.1: Distribution of hydrothermal vents. This map was created using the InterRidge ver. 2.1 database

cial biological habitat is established [10]. Deep sea hydrothermal vents were first discovered in 1977 along the Galápagos Rift [12]. In the last four decades, numerous investigations have been carried out along most major rift systems, including the East Pacific Rise [13], the Mid-Atlantic Ridge [14] and the Central Indian Ridge [15], and were later extended to back-arc basins [16]. Extremely slow spreading ridges like the Southwest Indian Ridge [10, 11] and the Gakkel Ridge [17] were also investigated immediately after they were discovered. Intensive venting activities, significantly greater than predicted [8, 9], have been identified along the slow and ultraslow spreading ridges (󳶳 Fig. 7.1). Hydrothermal vents have been considered windows to explore the origin of life. The hypothesis that life comes from an underwater hydrothermal environment has been supported by considerable evidence since it was proposed in the late 1980s [18]. Hyperthermophiles have a slow rate of evolution, and they branch off at the root of the phylogenetic tree. Studies have confirmed that the abiotic formation of short-chain fatty acids, such as formic and acetic acids, does occur in hydrothermal systems [19]. Furthermore, coprecipitated (Ni, Fe)S can act as a catalyst and reduce CO2 and CO to amino acids [20], which can then be converted into peptides [21]. The discovery of the reversed tricarboxylic acid cycle, which is induced by FeS without enzymes, also provides us with a model of chemoautotrophy [22]. Hydrothermal vents have also been considered as hotspots for deep sea mining. Polymetallic sulfides formed by hydrothermal venting on the sea floor, together with other types of polymetallic nodules and ferromanganese crusts, have long attracted attention as alternative sources of metals to terrestrial deposits. The occurrence of many of these deposits in international waters has necessitated their regulation under the UN Convention on the Law of the Sea through the establishment of the International

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Seabed Authority (ISA). The ISA was established in 1994 with its headquarters in Jamaica for the purpose of regulating activities in international waters beyond the national jurisdiction of any country. By 2015, the ISA had 25 Registered Pioneer Investors subsequently called ‘Contractors’ for exploration (http://www.isa.org.jm). The decision to commence mining of any deep sea mineral will depend on the availability of metals from terrestrial sources and their price in the world market, as well as technoeconomic analysis based on capital and operating costs of the deep sea mining system. As a rough estimate, approximately 30% of the benefit could be generated from polymetallic nodule mining if the techniques allow, which is not currently the case. In spite of such potential, most of the deep sea mineral deposits can only be termed as ‘resources’ but not ‘reserves’ as they cannot be economically recovered under prevailing economic conditions, but may be exploitable in the foreseeable future. Techniques for the evaluation of the distribution and potential of deep sea mining as well as techniques for resource mining have been largely developed during the past decades simply due to the economic interest. In 2016, SMD (Soil Machine Dynamics Ltd) released the first commercial deep sea mining system, which is still being tested. Meanwhile, the ecological concerns are also a burning question. The mining activity will certainly disturb the local ecosystem, especially the hydrothermal ecosystems where a large number of fauna live. According to the ISA, any contractor has to provide solid evidence that deep sea mining will be ecologically sustainable. Some researchers argue that mining underwater could be more ecologically friendly than mining on land. For example, the acid water discharged from mining is causing much pollution in terrestrial mining areas but would not be problem in the ocean because it will be neutralized by the alkaline seawater.

7.2 Hydrothermal vents on the Southwest Indian Ridge The Indian Ocean has three spreading ridge branches, the Central Indian Ridge (CIR), the Southwest Indian Ridge (SWIR), and the Southeast Indian Ridge (SEIR). Their spreading rates represent rapid paradigm changes, varying from the ultraslow spreading of the SWIR (∼12 mm/year) to the intermediate spreading of the SEIR and CIR (50–60 mm/year) [23]. The SWIR separated the African and Antarctic plates before 100 Ma, and it is currently a major plate boundary within the global ocean [24]. The SWIR spans a distance of approximately 8 × 103 km and is located fully within the southern hemisphere. The SWIR extends 7,700 km from the Bouvet Triple Junction (BTJ) (55°S, 00°40󸀠 W, intersects the MAR and the Americas-Antarctica Ocean Ridges [AAR]) to the Rodriguez Triple Junction (RTJ) (25°30󸀠 S, 70°E, intersects the CIR and the SEIR) [25]. The oceanic crustal thickness adjacent to the SWIR is between 3 and 6 km. There are nonmagmatic and magmatic spreading ridges alternating along the length of the SWIR. The tectonic environment of the axial rift valley is diverse, with large variability in terrain; the deepest water depth is as much as 5,000 m. According

164 | 7 Life at the hydrothermal vent field of the Southwest Indian Ridge

to existing geological and geophysical data, dynamic volcanism and tectonic activity in some segments of the SWIR make it possible for hydrothermal systems to have heat and fluid flow channels, which provide favorable conditions for hydrothermal activity and the formation of massive sulfide deposits [24]. The first indirect evidence for the presence of hydrothermal venting, from six hydrothermal plume anomalies overlying the eastern SWIR, was obtained on the Fuji voyage in 1997 [10]. The first submersible investigations to locate hydrothermal vents or sulfidic mineralization along the ultaslow spreading SWIR were carried out in 1998 during the INDOYO cruise with RV Yokosuka and submersible Shinkai 6500 [26]. Since then, the SWIR has become a hotspot for hydrothermal polymetallic sulfide resource exploration [27–30]. Hydrothermal activity on the ultraslow spreading SWIR was first visually confirmed near 49°E in 2007 by expedition DY115-20 of D/V Dayang Yihao. Since then, eight new hydrothermal areas have been identified by Chinese research groups through four research cruises in the SWIR [24]. Sulfides, basalt, and biological samples were obtained in this region, which was subsequently named as the Longqi vent field (37°47󸀠 S, 49°39󸀠 E) [31]. The Longqi field, which holds a large polymetallic resource, is more than 2,300 km away from Kairei, the closest neighboring surveyed vent field, and more than 2,500 km away from Solitaire. Based on previous research survey results, China submitted an application to explore a 10,000 km2 region of the Southwest Indian Ridge containing polymetallic sulfides, and this application was subsequently approved by the International Seabed Authority at their 17th meeting on 19 July 2011. As regulated by the contract, this potential mining area needs to be systematically investigated and only 25% of the contract area will be retained in 2021 for a further mining operation [33]. Together with the exploitation proposal, a comprehensive proposal on the ecological preservation is required. During the past years, more and more data and samples have been generated regarding the biological communities and their ecological impacts. Now we are getting a better understanding of this unique hydrothermal system in the SWIR and the life that inhabits it.

7.3 The biologic communities distributed on the SWIR 7.3.1 Faunal communities The SWIR is particularly intriguing because it is the corridor between the Pacific Ocean and the Indian Ocean, thus the fauna herein are supposed to have signatures of both oceans, and because its ultraslow spreading characteristic makes the faunal communities unique [10]. Recently, a few studies have focused on elucidating how ventendemic species disperse and maintain connectivity between different hydrothermal vent fields, which is important for understanding their biogeography and speciation. For example, based on the morphological description, the ‘scaly-foot gastropod’ C. squamiferum (Mollusca: Peltospiridae) at the Longqi field has a brown shell and dark

7.3 The biologic communities distributed on the SWIR

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165

sclerites, while the C. squamiferum from Kairei has a dark shell and dark sclerites, and the C. squamiferum from Solitaire has a yellowish shell and white sclerites [34]. Each morphotype is consistent and specific to the corresponding location. These C. squamiferum from hydrothermal vents are gastropods possessing numerous dermal sclerites, which may be mineralized with iron sulfide, making them the only known living metazoan type to incorporate iron in their skeletons [35, 36]. Moreover, according to the gene flow results using a cytochrome oxidase c subunit I (COI) gene identified from C. squamiferum collected in Longqi, Kairei and Solitaire fields, there is low connectivity between SWIR and CIR vent populations within C. squamiferum [37]. However, as they are limited by sampling size, the conclusions above are subject to more evidence. With a series of China Ocean Mineral Resources Research and Development Association (COMRA) routine cruises for exploration of polymetallic sulfides found at the SWIR, the subsequent investigation of faunal and genetic diversity may have implications for management of the Longqi site and possibly other sites on the SWIR [33, 34].

7.3.2 Microbial communities The microbial ecosystem is fueled by biochemical reactions between the reduced venting fluid and oxidized sea water. Within the microbial groups, the chemolithoautotrophic microorganisms are the primary producers of hydrothermal microbial communities, and they are able to fix inorganic carbon using chemical energy obtained through the oxidation of reduced compounds such as hydrogen sulfide, hydrogen, and methane. Heterotrophic microorganisms that depend on the energy flow through these chemoautotrophs form another component of the microbial community. These are the metabolically versatile chemolithoautotrophic microorganisms, producing organic matter that forms the base of the food chains in these deep sea highly productive ecosystems [38]. As there are numerous detailed reports on other vent fields including the EPR and the MAR, suggesting that the microorganisms are playing important ecological roles in such ecosystems. However, the distribution pattern of microbial communities is still obscure at the SWIR, since only limited samples have been taken along SWIR (󳶳 Tab. 7.1). There are several studies focused on the microbial community on hydrothermal sulfides collected from SWIR. The first article detected the lipid biomarker for the investigation of related microbial groups. The biomarkers associated with methanogenic archaea that belong to Euryarchaeota and sulfate reducing bacteria (SRB) have been observed. The lipid composition of hydrothermal oxide also showed that archaeal activity was limited in hydrothermal oxide, and insead sulfur-oxidizing bacteria formed the main microbial community [39]. Then, in 2013, the molecular evidence for microorganisms participating in Fe, Mn and S biogeochemical cycling in two low temperature hydrothermal fields at the SWIR was published. Analyses of 16S rRNA gene sequences showed that Zetaproteobacteria, Pseudoalteromonas,

166 | 7 Life at the hydrothermal vent field of the Southwest Indian Ridge

Tab. 7.1: Studies of microbial diversity related to hydrothermal vent samples at the SWIR Sample style Samling site

Time

Depth

Method(s)

Reference

Sulfide

49.6°E

2007

–

Lipid biomarker

Lei et al. 2012 [39]

Plume

49°39󸀠 E,

2010

2,800 m 16S rRNA/aprA Clone library

Ren, et al. 2012 [45]

Sulfide

49.6474°E, 37.7805°S; 50.4678°E, 37.6587°S

2008

2,786 m 16S rRNA/aprA 1,745 m Clone library

Li, et al. 2013 [40]

Sediment

27°50.98󸀠 S, 63°56.17󸀠 E 2007

2,951 m 16S rRNA/NifH Clone library

Wu et al. 2014 [43]

Sulfide

37°66󸀠 S, 50°46󸀠 E; 37°78󸀠 S, 50°65󸀠 E

2008

1,744 m Metagenomic 2,783 m sequencing

Cao et al. 2014 [41]

Carbonate sediment

51.0091°E, 37.6081°S

2008

2,034 m 16S rRNA Clone library

Li et al. 2014 [42]

Plume

49°37󸀠 E

2010

2,380– 16S rRNA 2,935 m Clone library

Li et al. 2016 [44]

37°47󸀠 S

Leptothrix, and Pseudomonas were potential Fe and Mn oxidizers in the low temperature hydrothermal environments, but they were not present in equal abundance among the subniches. Furthermore, the aprA gene clone library indicated that members of Gammaproteobacteria and Alphaproteobacteria were involved in the S oxidation process, while members of Deltaproteobacteria, Nitrospirae, Firmicutes, and archaea might participate in the S reduction process. The Fe, Mn, and S oxidizers and reducers might actively participate in hydrothermal biogeochemical processes, which could influence the transfer of chemical species and the formation of biogenic minerals at the SWIR [40]. In 2014, an analysis of metagenomes of low temperature hydrothermal chimneys on the SWIR revealed an influential microbial sulfur cycle. Phylogenetic diversity of 16S rRNA selected from metagenome data indicated that Delta- and Gammaproteobacteria represented the most abundant classes for both sulfides, followed by Epsilon-, Alpha- and Zetaproteobacteria. The taxonomic results also revealed that a few dominant bacteria participated in the microbial sulfur cycle, particularly sulfate-reducing Deltaproteobacteria. Additionally, several carbon metabolic pathways, in particular the Calvin–Benson-Bassham (CBB) pathway and the reductive tricarboxylic acid (rTCA) cycles for CO2 fixation, were identified in sulfuroxidizing autotrophic bacteria [41]. Recently, papers describing the microbial communities and their potential metabolic pathways within the semiconsolidated carbonate sediments of the SWIR were published. The results revealed that Gammaproteobacteria, Acidobacteria, and Thaumarchaeota members dominated the bacterial and archaeal clone libraries respectively. Additionally, it was revealed that the Thaumarchaeota and Alphaproteobacteria were the potential players that participated in N and S cycles in this marine car-

7.3 The biologic communities distributed on the SWIR

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167

bonate sedimentary environment [42]. Wu et al (2014) also focused on the microbial community distribution pattern in a hydrothermal vent sediment sample that was obtained from a TVG3 site, southwest of the Kairei and Edmond hydrothermal field on the SWIR during the Chinese cruise on R/V Dayangyihao. Within the archaea, the marine benthic group E (MBGE) and marine group I (MGI) belonging to the phyla Euryarchaeota and Thaumarchaeota respectively, dominated the sediment sample. The dominant bacterial clones belong to the Gammaproteobacteria. Additionally, phylogenetic analysis based on the NifH gene indicated that bacteria play an important role in nitrogen fixation in the SWIR environment [43]. Quite recently, in 2016, the microbial communities of hydrothermal plumes at an active field 49°37󸀠 E collected during the DY115-21 cruise aboard the R/V Dayang Yihao (2010) was analyzed using molecular approaches. Phylogenetic analysis showed that the dominant groups were members of Alphaproteobacteria and Gammaproteobacteria and members of MG-I within the Crenarchaeota. Furthermore, no significant difference in microbial composition between plume sample and ambient seawater was detected, suggesting that the SWIR hydrothermal plumes were sourced from ambient seawater rather than from seafloor vent derived niches [44]. The results of bacterial diversity investigations in this article are quite similar to the distributional pattern in a plume reported in 2012 [45].

7.3.3 A case study on the microbial communities at the Longqi field In 2015, we obtained several chimney samples at the Longqi field during the Dayang 35cruise using the manned deep sea submersible vehicle Jiaolong The microbial communities in the chimney samples were then analyzed (namely JL90, JL94D, JL94H, and JL95) (󳶳 Fig. 7.2). These samples represent distinct characteristics in terms of temperature, pH and metal compositions (󳶳 Tab. 7.2).

Fig. 7.2: Photographs of hydrothermal vent samples collected from the Longqi field on the SWIR

168 | 7 Life at the hydrothermal vent field of the Southwest Indian Ridge

Tab. 7.2: Concentrations of elements in chimney deposits and physicochemical characteristics of their hydrothermal fluids at the Longqi field on the SWIR. Chimney Element(mg/kg) Fe S Zn Mn Mg Ca Pb Cu Al Depth (m) Location Fluid Max. temp. (°C) pH Salinity DO (mg/L)

JL90

JL94D

JL94H

JL95

261,600 300,200 211,200 3,786 20,960 9,268 1,417 806.8 661.6 2,746 49.6501525°E, 37.7832506°S

305,900 394,400 233,100 33 24 112 703 541 5,680 2,768 49.6487677°E, 37.7832151°S

286,500 302,400 2,938 14 324 46,000 57 67,330 141 2,778 49.6487677°E, 37.7832151°S

68,780 1,682 154 343,800 13,720 13,140 41 42 429 2,775 49.6477092°E, 37.7799720°S

145 4.85 4.0 24.7

13.3 – – –

362 – – –

379 3.42 4.0 14.3

Range

13.3–379 3.21–4.85 3.7–4.5 0.5–24.7

Bacterial and archaeal populations correlating with the nature of the vents were evaluated using tag pyrosequencing of the 16S small subunit (SSU) ribosomal RNA (rRNA) genes. In this study, high-throughput DNA based analysis on environmental samples was applied to investigate the microbial communities of chimney samples collected from the Longqi hydrothermal field, and a total of 110,473 and 200,630 amplicons on bacterial and archaeal 16S SSU rRNA genes, respectively, have been obtained, suggesting that the method was efficient. According to the Venn diagrams(󳶳 Fig. 7.3), JL94D and JL94H shared the same sampling location and also the highest percent of bacterial OTUs (operational taxonomic unit), 9.6%, followed by the overlap between JL90 and JL94D, 7.7%, JL90 and JL94H, 7.6%. It might because JL90 was quite near the location of JL94D/H. But the bacterial OTU overlap between JL95 and JL94 or JL90 was about or below 6.2%. Based on the whole profiles of phyla, 28 of 29 bacterial phyla and 11 of 12 archaeal phyla were shared in the four chimney samples (󳶳 Fig. 7.4, 󳶳 Fig. 7.5). Only the bacterial phylum of Fusobacteria was absent in JL95 and the archaeal candidate phylum SM1K20 was not detected in JL90. The species composition overlap reflected the microbial resemblance among the four chimneys at the Longqi field on the SWIR. Meanwhile, according to the relative abundance of categories, there are variations in the microbial communities among chimneys in the SWIR region. Microbial community structures were clearly correlated with the environmental parameters, and among all the considered parameters the in situ temperature was the most influential. Previous studies demonstrated that the biogeographical patterns of

7.3 The biologic communities distributed on the SWIR

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169

Fig. 7.3: Venn diagrams showing the estimated OTU (97 identity threshold) richness shared among bacterial (A) and archaeal (B) communities from hydrothermal vent chimneys at the Longqi field on the SWIR. Shared OTU richness estimates were calculated using the program Qiime (version 1.9.0). Venn diagrams were plotted using the Venn diagram package of R. Numbers in the Venn diagrams indicate the number of OTUs. The table below the Venn diagrams shows the percent of sharing OTUs for each pair within all four chimney samples

microbial communities were shaped in part by local fluid geochemistry in active hydrothermal vent chimneys [46], mineralogy on inactive seafloor sulfide deposits [47] and geological processes, such as eruption, on diffuse-flow vents [48]. In this study, the temperature of fluid ranges from 13.3 °C at low temperature vents to 145 °C at high temperature vents, where temperatures of 362 °C and 379 °C were reached, which demonstrates quite a difference among chimney samples (󳶳 Tab. 7.2). Potential energy sources for deep sea vent chemoautotrophy include reduced sulfur compounds, molecular hydrogen, reduced metals and ammonium. The microorganisms potentially capable of using the above energy sources have been detected in our samples (󳶳 Tab. 7.3). Classic sulfur-oxidizing bacteria have been detected as the dominant type, suggesting a strong sulfur metabolizing potential in all tested chimney samples. In this study, Gammaproteobacteria were the most abundant groups, mainly consisting of sulfide oxidizing bacteria within the families of Thiotrichaceae Ectothiorhodospiraceae, Thiohalophilus and Piscirickettsiaceae in the bacterial library of all detected chimney samples (󳶳 Tab. 7.3). Epsilonproteobacteria were known to play a significant role in carbon, nitrogen and sulfur cycling and were consistently shown to be the most numerically abundant bacteria in different hydrothermal and subsurface environments, including sediment [49], hydrothermal fluids [50], hydrothermal plumes [51], and vent chimneys [52]. Based on our V4 tag sequence data, bacterial am-

Nitrospira; Nitrospirales; Nitrospiraceae; Thermodesulfovibrio Deltaproteobacteria; Desulfarculales; Desulfarculaceae; Desulfatiglans Deltaproteobacteria; Desulfobacterales; Desulfobacteraceae Deltaproteobacteria; Desulfobacterales; Desulfobulbaceae Epsilonproteobacteria; Campylobacterales; Campylobacteraceae; Sulfurospirillum Thermodesulfobacteria; Thermodesulfobacteriales; Thermodesulfobacteriaceae; Thermosulfurimonas

Sulfate reduction

49.653

0.429 0.006 0.034 3.028 0.360 0.022

0.016 36.326 0.193 5.360 2.908 0.971

JL90

40.742

0.657 0.557 1.786 5.398 0.627 0.060

0.105 24.724 0.040 5.955 0.642 0.191

JL94D

46.277

22.053 0.154 0.604 3.790 0.194 1.044

1.780 8.633 0.855 4.160 0.282 2.728

JL94H

Relative abundance (%)

7.636

0.246 0.008 0.016 0.131 0.008 0.008

0.000 1.246 1.065 1.950 0.295 2.663

JL95

The relative abundance in sequencing library is for each sample’s Miseq dataset. Taxa are designated by class (phylum for Crenarchaeota and Thaumarchaeota), order, family, and genus.

a

Sum of S oxidation and sulfate reduction

Aquificae; Aquificales; Aquificaceae; Hydrogenivirga Epsilonproteobacteria; Campylobacterales; Helicobacteraceae Gammaproteobacteria; unknown; unknown; Thiohalophilus Gammaproteobacteria; Thiotrichales; Thiotrichaceae Gammaproteobacteria; Thiotrichales; Piscirickettsiaceae Gammaproteobacteria; Chromatiales; Ectothiorhodospiraceae

Bacterial/Archaeal Taxonomy

S oxidation

Bacteria

Function

Tab. 7.3: Potential ecological function of tag sequences for which obvious metabolisms can be inferreda .

170 | 7 Life at the hydrothermal vent field of the Southwest Indian Ridge

7.428 0.028 0.072 0.290

Epsilonproteobacteria; Nautiliales; Nautiliaceae; Nitratifractor

Betaproteobacteria; Nitrosomonadales; Nitrosomonadaceae; Nitrosomonas

Alphaproteobacteria; Rhizobiales

Nitrate reduction

Nitrification

N fixation

Sum of ammonia, nitrite oxidation, nitrification and N fixation

H oxidation

Zetaproteobacteria; Mariprofundales; Mariprofundaceae; Mariprofundus

Deltaproteobacteria; Desulfuromonadales; Desulfuromonadaceae; Desulfuromusa

Fe(II) oxidation

Fe(III) reduction

Crenarchaeota; Thermoprotei; Desulfurococcales Crenarchaeota; Thermoprotei; Thermoproteales

Thaumarchaeota

Sulfate reduction

Ammonia oxidation

Total archaea

Alphaproteobacteria; Rhodobacterales; Rhodobacteraceae; Roseobacter

Mn oxidation Total bacteria Archaea

Sum of Fe(II) oxidation and Fe(III) reduction

1.122

Gammaproteobacteria; Methylococcales; Methylococcaceae; Methylothermus

CH4 oxidation

46.531

45.667

0.538 0.326

0.068 60.194

2.655

0.238

1.295

0.39

0.002

0.489

0.002

Sum of H oxidation

Aquificae; Aquificales; Aquificaceae; Hydrogenobacter Aquificae; Aquificales; Hydrogenothermaceae; Persephonell Epsilonproteobacteria; Campylobacterales; Hydrogenimonaceae; Hydrogenimonas

5.739

Nitrospira; Nitrospirales; Nitrospiraceae; Nitrospira

Nitrite oxidation

1.196

Gammaproteobacteria; Chromatiales; Chromatiaceae; Nitrosococcus

Ammonia oxidation

1.068

1.062

0.003 0.003

1.234 47.569

4.43

1.801

2.468

0.161

0.366

0.080 0.261 0.025

0.797

0.562

0.010

0.105

0.005

0.115

33.207

24.636

7.460 1.111

0.353 58.604

3.993

0.591

1.084

2.318

7.064

1.908 4.958 0.198

0.917

0.450

0.004

0.062

0.026

0.375

Relative abundance (%)

Bacterial/Archaeal Taxonomy

Function

Tab. 7.3: (continued)

42.763

42.745

0.014 0.004

0.049 15.502

0.803

0.041

0.705

0.057

0.057

0.016 0.033 0.008

6.957

2.622

1.418

0.057

0.992

1.868

7.3 The biologic communities distributed on the SWIR |

171

172 | 7 Life at the hydrothermal vent field of the Southwest Indian Ridge

Fig. 7.4: Clustering analysis tree of the 16S rRNA bacterial community structure of chimneys (JL90B\JL94DB\JL94HB\JL95B) from the SWIR, the EPR and the MAR. LS7=Lucky Strike. Lucky Strike vent field located at the MAR (38); CH7=chimney sample from the EPR9°N M vent

plicons consisted of Epsilonproteobacteria, ranging from 1.3% at high-temperature vent chimney to 37.6% for diffusive vent chimney samples. Certain Epsilonproteobacteria sequences are dominant, and closely related to the known chemosynthetic, sulfur-oxidizing genera Sulfurovum and Sulfurimonas. Similar Epsilonproteobacteria communities were also found in cool, diffusive flows at the Axial Seamount on the Juan de Fuca Ridge [53] and the biofilms growing on the chimney walls at the Loki’s Castle vent field [54]. Interestingly, the most abundant Epsilonproteobacteria OTU belongs within the genera Sulfurovum and Sulfurimonas, which were recovered from all active chimneys in our study, and also recovered from inactive sulfides in the EPR [55]. It is possible that these groups represent widely distributed species at active sulfides and survivable relict populations at inactive chimneys by oxidizing sulfide minerals. Thaumarchaeota is an abundant and ubiquitous phylum of archaea that plays a major role in the global nitrogen cycle [56], which might be the major ammonia oxidizing archaea (AOA) among the recovered microorganisms at high-temperature vent chimneys JL94H and JL95 with relative abundances of over 40%, and also for JL90 at ∼24% (󳶳 Tab. 7.3). Moreover, the ammonia oxidizing bacteria (AOB) within Nitrosococcus were recovered and represented the abundant genus (over 1%) in JL90 and JL95. Related genera of nitrifier, the Nitrospira and Nitrosomonas, were found in all samples and with highest abundance in JL95. Recently, a completely nitrifying bacterium from the genus Nitrospira was reported [57], indicating that the globally distributed nitrite oxidizers fundamentally changed the picture of nitrification and might act as key microbial communities for nitrogen cycling on the high-temperature chimney JL95 and other samples at the Longqi hydrothermal field. To evaluate the effects of geological and geochemical characteristics on microbial communities on the surface of active chimneys in the Longqi field at the ultra-

7.3 The biologic communities distributed on the SWIR

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173

slow spreading ridge of the SWIR, we compared the bacterial and archaeal distribution pattern with habitats of active chimney both from the slow spreading ridge of the MAR and the fast spreading ridge of the EPR (󳶳 Fig. 7.4, 󳶳 Fig. 7.5). The results showed that all the archaeal communities of chimneys in the SWIR and the EPR were clustered into different branches than the high-temperature vent chimney in the MAR. That both CH7 and JL94D were clustered in the same branch might be with the reason for the absolute abundance (over 79%) of the Woesearchaeota. The high-temperature vent chimney LS7 from the MAR was highly dominated by Epsilonproteobacteria, which is quite different from the bacterial composition of other active chimneys from the SWIR and the EPR, and might lead clustering into a separate branch. The bacterial communities of JL95 and CH7 from the EPR were surprisingly clustered in to a group with limited Epsilonproteobacteria, but with dominant Gammaproteobacterial sulfur oxidizers. Besides the Epsilonproteobacterial and Gammaproteobacterial sulfur oxidizers mentioned above, Fe- and ammonia-oxidizing chemoautotrophs were also identified from all chimneys in our study, just as reported by previous studies on other hydrothermal vent fields [58]. The major order of Gammaproteobacteria was Xanthomonadales in JL95, which was also observed in the microbial community of deep sea sediments [59]. Additionally, the genus Sulfurovum and Sulfurimonas within Epsilonproteobacteria was recovered from all active chimneys at the SWIR, and also from inactive sulfides in the EPR [55]. These groups represent the most widely distributed species at active sulfides and the survivable relict populations at inactive

Fig. 7.5: Clustering analysis tree of the 16S rRNA archaeal community structure of chimneys (JL90A\JL94DA\JL94HA\JL95A) from the SWIR, the EPR and the MAR. LS7=Lucky Strike. Lucky Strike vent field located at the MAR (Flores et al., 2011); CH7=chimney sample from the EPR 9°N M vent

174 | 7 Life at the hydrothermal vent field of the Southwest Indian Ridge

chimneys by oxidizing sulfide minerals. Within archaeal communities, the culturable genera of Pyrococcus and Thermococcus were recovered frequently in molecular environmental surveys at hydrothermal vents [60], but were absent in our archaeal tag sequences, which might be because of the distinct geographic locations and related geochemical conditions. To summarize, this case study reported the distribution and diversity of the prokaryotic communities on the surface of chimneys collected from the Longqi field on the SWIR. The 16S SSU rRNA gene analysis suggested that bacterial communities were highly diversified among all the detected samples. Compared to bacteria, the lower diversity of archaeal phylotypes agreed with other molecular surveys indicating that marine hydrothermal vent archaeal diversity is relatively limited [61]. Combined with the functional analysis of bacteria and archaeal communities, sulfur oxidation and reduction may be important energy metabolism pathways in low and high temperature vent chimneys with high abundance of SOB within Gammaproteobacteria and Epsilonproteobacteria. Meanwhile, ammonia oxidation may be another major pathway to providing energy for microbial ecology systems on high temperature active chimneys. Our results provided more details for the characterization of the microbial roles in ecologic and minerogenic processes at the SWIR, especially in S and N cycling. Acknowledgment: This work was supported by China Ocean Mineral Resources R&D Association (Grant No. DY125-22-04). We would like to thank the pilots, officers, and crew members of DaYangYiHao/Jiaolong (Dayang 35 cruise) for their dedication and expertise, which made this study possible. We also thank the Chief Scientist and crew members of the cruise of R/V Atlantis and ROV Jason II AT26-10 (December 27, 2013– January 26, 2014) for accommodating us.

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Scott D. Wankel, Annie Bourbonnais, and Chawalit Charoenpong

8 Microbial nitrogen cycling processes at submarine hydrothermal vents 8.1 Introduction In comparison to the abundance of other redox-active elements commonly found in hydrothermal vent systems (e.g. sulfur, iron, carbon), the abundance of reactive nitrogen (N) is generally low, and perhaps this is why it has not been as widely studied. Nevertheless, as a required nutrient for all life as we know it, its distribution places implicit constraints on the magnitude, extent and forms of life that can persist on Earth. The study of the abundances and forms of nitrogen therefore sheds important light on the nature of biological communities and mechanisms at work in these extreme environments, which in turn directly reflect how life has evolved and continually adapts to acquire, metabolize and recycle this key element for life. In fact, given the relatively nonreducing nature of the Earth’s early atmosphere (e.g. [1, 2]), the availability of reduced nitrogen in the form of ammonium (NH+4 ) is considered a prerequisite for the origins of life from prebiotic precursor compounds. Hydrothermally hosted transformations of nitrogen-bearing compounds have been implicated as playing a central role in the origin of life on Earth [3–5] and potentially elsewhere in our solar system [6–8]. From an ecosystem perspective, the abundance and distribution of nitrogen-bearing compounds plays a critical role in regulating the viability of a wide variety of metabolic pathways that are tightly coupled with the cycling of other important elements including carbon, oxygen and sulfur; this is especially apparent in hydrothermal vent ecosystems. As a nutritional element, organisms generally require N in its most reduced form, thus NH+4 generally represents the most bioavailable form for assimilation. This being said, highly specialized diazotrophic organisms have evolved the ability to directly fix gaseous dinitrogen (N2 ) and may play important roles in vent ecosystems. In addition to serving as a nutrient, oxidized forms of nitrogen can also serve as important electron acceptors for energy conservation by microbial communities. Biological reduction of oxidized forms such as nitrate (NO−3 ) and nitrite (NO−2 ) can be coupled with the oxidation of organic carbon or through chemolithotrophic pathways involving other chemical species such as reduced species of sulfur, iron or hydrogen (e.g. [9–11]). Finally, some forms of N can serve as energy sources (NH+4 and NO−2 ) for microbial communities catalyzing chemoautotrophic metabolisms including NH3 oxidation (to NO−2 using O2 ), NO−2 oxidation (to NO−3 using O2 ) and anaerobic NH+4 oxidation (to N2 coupling oxidation of NH+4 with reduction of NO−2 ). Before discussing specific microbially catalyzed processes under hydrothermal conditions in detail, we first provide a brief overview of hydrothermal venting and https://doi.org/10.1515/9783110493672-008

180 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

fluid mixing, including a general description of the distribution of the most common N species. We also briefly introduce the major microbial N transformations and how fluid mixing operates to favorably poise energetic yields for catalysis of these reactions. In addition we introduce the use of commonly applied tools for investigations of nitrogen cycling processes – specifically the application of natural abundance stable isotopes as a tool for detecting and quantifying active transformations as well as the interrogation of genetic information contained in DNA and RNA at both the organismal and environmental levels (‘omics’).

8.1.1 Hydrothermal fluid venting and mixing The chemical composition of submarine fluids venting at oceanic spreading centers and back-arc volcanic systems integrates the influence of a wide range of physical, chemical, and biological processes during the convective circulation of seawater through the oceanic lithosphere (idealized in 󳶳 Fig. 8.1). The composition of ‘endmember’ fluids that have not been influenced by mixing with seawater reflect the net result of deep subsurface reactions and may emerge under focused flow at the seafloor at temperatures as high as 407 °C through chimneys comprised of sulfide minerals. Nevertheless, young oceanic crust is highly permeable, which facilitates extensive subsurface mixing of hot endmember fluids with cold (2 °C) seawater and results in large areas of lower temperature venting of diffuse fluids. In conjunction with the unique composition of endmember fluids relative to seawater, the short residence times of fluids in subsurface mixing zones and the kinetic barriers for many chemical reactions (especially those involving transfer of electrons), the chemical composition of these mixed low temperature fluids is typically far from thermodynamic equilibrium, and therefore they contain a substantial amount of potential chemical energy. Thus, when temperatures of these mixed fluids cool below the proposed temperature limit for life (∼122 °C; [12]), the resulting wealth of chemical energy supports complex ecosystems in near vent environments including the seafloor, the water column and the subsurface biosphere.

8.1.2 Nitrogen species found in hydrothermal vent fluids Given the combination of chemical disequilibrium and the diverse and highly adapted biological communities subsisting on energetic gain from metabolic transformations, hydrothermal fluids can exhibit broad ranges in the concentration of aqueous N species. In general, bottom seawater represents a cold source of oxidized components, including O2 as well as ∼20–50 µM NO−3 (󳶳 Fig. 8.1). In concert with the supply of energy (whether in the form of organic matter or as reduced chemical species such

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as H2 , H2 S, Fe(II), etc.), the biological demand for electron acceptors generally varies as a function of energetic potential. Under most conditions, the reductive consumption of bottom water O2 as cold bottom seawater is circulated into the warm subsurface is rapidly followed by reductive consumption of NO−3 as biological and chemical reactions ensue. Tellingly, although aqueous NO−3 concentrations in high temperature endmember vent fluids have not been widely reported, the few that do exist suggest near-zero abundances [13, 14]. 󳶳 Fig. 8.2 illustrates how variable mixing of cold, NO−3 rich seawater with hot, NO−3 poor vent fluid might affect the relative concentrations of NO−3 (plotted against Mg2+ as a conservative tracer of the proportion of seawater in a particular sample). Under temperatures permitting biological NO−3 consumption, deviations from the conservative mixing line might be expected as NO−3 is removed from solution. In contrast, much wider variations exist in reported levels of NH+4 (here using NH+4 to refer to ∑ NH+4 = NH+4 + NH3 ), with levels at sedimented spreading centers reaching as high as 15.3 mmol/kg in endmember fluids [15]. These extremely high levels are the result of low temperature remineralization and thermal decomposition (or pyrolysis) of organic matter (󳶳 Fig. 8.1) accumulating in the overlying sediments in system such as Guaymas Basin, Escanaba Trough and Okinawa Trough [15–19]. In stark contrast, given the absence of any substantially organic rich overburden, NH+4 levels are much lower in the context of unsedimented hydrothermal systems, with typical values ranging from 1–50 µmol/kg in endmember fluids [13, 20, 21]. The exact origin of this NH+4 and the processes underlying observed variations is unclear, however it could involve abiotic reduction of seawater NO−3 or N2 to NH+4 via high temperature reactions, microbial reduction of seawater NO−3 to NH+4 in shallow, low temperature mixing zones, and/or thermal degradation of microbial biomass associated with the subsurface biosphere. Whatever the level of NH+4 in the high temperature endmember fluids, variable mixing of cold, oxic, NH+4 depleted seawater and hot, NH+4 -containing vent fluids will also result in a conservative distribution of NH+4 with respect to Mg2+ , deviating only in response to processes that are either producing or consuming NH+4 (󳶳 Fig. 8.2). Dissolved N2 gas, although rarely quantified explicitly, is generally abundant in hydrothermal fluids (both high temperature and diffuse fluids). While air-saturated seawater in the deep ocean contains approximately 0.59 mmol/kg, reported values of N2 content in hydrothermal fluids have generally ranged from seawater equilibrated values up to 7 mmol/kg [22–24]. Higher values are largely interpreted as reflecting differential contributions of magmatic N2 , having an isotopic composition depleted in 15 N [25–28]. In addition to these more abundant species, the cycling of nitrogen also involves a number of intermediate species, the study of which has shed light on reaction mechanisms, as well as thermodynamic and kinetic considerations of energy gain by microbial communities. These N species include NO−2 , hydroxylamine (NH2 OH), nitric oxide (NO), nitrous oxide (N2 O), and dissolved organic nitrogen (DON). In practice, however,

182 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

most of these compounds are in low abundance and/or highly reactive, making their study in hydrothermal systems challenging. As such, very little effort has been made to explicitly quantify these intermediates in the context of hydrothermal vent systems. Notably, a few recent studies have begun reporting measured concentrations of NO−2 and N2 O (for example [13, 29, 30]). Both of these compounds can be products of either reductive or oxidative biological transformations, and their presence in these studies was interpreted as reflecting the influence of microbially catalyzed redox reactions. Study of the distribution and isotopic composition of these intermediate species represents a promising (though challenging) path forward for gaining more insight into the nature of specific processes.

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󳶣 Fig. 8.1: Cross sectional conceptual schematic of a hydrothermal vent system illustrating fluid pathways and prominent reactions involving transformations of major nitrogen bearing species. (a) Off-axis recharge of cold bottom seawater (containing NO−3 and N2 , but not NH+4 ) is convectively drawn into permeable ocean crust where elevated temperatures induce water-rock reactions leading to development of significantly reducing redox conditions, which are predicted to catalyze NO−3 reduction to NH+4 and/or N2 . Concentration and δ 15 N of NH+4 in high temperature vent fluids will reflect these high temperature reactions, while low temperature vent fluid compositions will also be influenced by subsurface microbially catalyzed processes. (b) Diagram illustrating major nitrogen transformations and species. Abiotic reduction of NO−3 (red arrows) may proceed as parallel, unidirectional reactions producing both NH+4 and N2 as a function of temperature and redox. In low temperature venting fluids, NH+4 originates from mixing of both endmember NH+4 as well as biologically derived NH+4 (either from pyrolyzed organic matter or DNRA supported by shallow inmixing of bottom seawater NO−3 ). The warm subsurface and exterior biosphere may acquire nutritional N by either N2 fixation or assimilation of NO−3 or NH+4 . Shallowly circulated NO−3 and/or NO−2 may be microbially reduced in the warm biosphere by organisms catalyzing denitrification, anammox (using NO−2 ; not depicted) or DNRA. Where NH+4 encounters oxic conditions ‘new’ NO−3 may be produced by nitrification (blue arrow), through a NO−2 intermediate (not shown). See text for details

Fig. 8.2: Illustration of the influence of fluid mixing on the predicted composition of samples collected from hydrothermal vent environments. Conservative mixing lines are indicated for mixing of vent fluids rich in NH+4 and free of NO−3 with bottom seawater having high NO−3 and virtually no NH+4 . Deviations from these mixing lines may offer insights into key reactions occurring involving NH+4 and/or NO−3 . For example, samples falling below the mixing line would suggest nonconservative consumption of a compound (e.g. NH+4 being removed by biological assimilation or nitrification)

8.2 Energetic considerations A majority of microbially catalyzed nitrogen transformations are redox reactions involving the transfer of electrons from one reactant species to another. The energy obtained from these reactions is used to fuel microbial growth and sustain various metabolic functions essential for microbial life. Within most hydrothermal systems there is vast range of reaction conditions represented by the sharp geochemical gradients, which offer a diverse (and often dynamic) system of chemical disequilibria suitable for a variety of possible redox reactions. Kinetic barriers act to maintain these

184 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

chemical disequilibria long enough for microbes to take advantage of the potential energy, with increasing temperature often beginning to overcome these barriers and limiting energy availability. Fundamentally, an understanding of the thermodynamic potential of these reactions offers a first order perspective on which reactions may be more energetically favorable (e.g. how much energy can be gained through its catalysis) under different environmental conditions (e.g. concentrations, temperature, pH, pressure, etc.), and therefore, which metabolic reactions might be expected to occur. However, while an impressive diversity of microbially catalyzed redox reactions have been demonstrated, the existence of energy-gaining potential does not by itself mean that the reaction will be occurring, only that the potential exists for energy to be gained by its catalysis. Mixing of hot, reducing hydrothermal fluids with cold, oxic seawater combines large arrays of electron acceptors and donors in chemical disequilibria; thus, a diversity of redox species may be available for microbially mediated transformations of nitrogenous species. Depending on the temperature, pressure, and concentrations of the species involved, ∆Gr for these reactions may vary considerably [31]. The amount of energy available from any chemical reaction is termed the overall Gibbs free energy of the reaction (∆Gr ), where negative values indicate a spontaneous reaction (exergonic) and positive values indicate a nonspontaneous reaction (endergonic). This can be calculated by: Kr ∆Gr = RT ln Qr where R is the ideal gas constant, T is the temperature in Kelvin, and Kr and Qr are the equilibrium constant and the activity product of the reaction respectively. Alternatively, ∆Gr can be calculated using the standard partial molar Gibbs free energy of the reaction (∆G0r ) over a wide range of temperatures for different reactions using the following equation [31]: ∆Gr = ∆G0r + RT ln Qr Thus, factors governing the thermodynamic favorability of a reaction may vary sharply across the mixing gradients commonly encountered in hydrothermal environments. Determining values of Qr under real-world hydrothermal conditions can be challenging, especially since many nitrogenous reactants are reactive intermediates (e.g. NO−2 , NO, N2 O), which can be notoriously difficult to quantify and/or predict. Factors including temperature, pH and product/reactant activities will often shift sharply, simultaneously modifying values of both ∆Gr and Qr and therefore the energetic favorability of the reaction. Specifically, these gradient driven changes in reaction conditions can dramatically influence the solubilities of gaseous N species (e.g. N2 , N2 O, NO, NH3 ), the speciation of some forms of N (NH3 /NH+4 ) and of course the activities of reactants, including those to which nitrogen redox reactions may be directly or indirectly coupled. As an example, the energetic yield of the reduction of NO−3 to N2 coupled to oxidation of H2 (a form of autotrophic denitrification) is illustrated in 󳶳 Fig. 8.3

8.2 Energetic considerations

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| 185

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Fig. 8.3: Illustration of the energetic yield of an autotrophic denitrification reaction. Contours are ∆Gr values (in kJ/mol N2 ) for the reaction 2NO−3 + 5H2(aq) + H+ → N2(aq) + 6 H2 O at 350 bar and 4, 40, 80, and 120 °C as a function of pH and log aH2 . The activity of H2 O(l) is taken to be unity, pressure is set to 350 bar, and values of aNO−3 and aN2 are set as 3 × 10−5 and 6.63 × 10−4 respectively, representing common conditions encountered in the deep ocean.

as a function of H2 activity and pH over four different temperatures from 4 °C to 120 °C (holding pressure and NO−3 concentration constant at 350 bar and 35 µM respectively). As shown, the thermodynamic favorability of the reaction increases (more negative ∆Gr values) as H2 activity increases, while favorability decreases at higher pH values and temperatures. Such redox transformations of nitrogen are generally coupled with any of a number of other redox-sensitive elements under hydrothermal conditions (H, O, C, S, Fe and Mn), the reactivity of which are of course also subject to the sharp gradients of

186 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

hydrothermal fluid mixing. For example, reactivity may be affected by mineral precipitation (Fe, S, Mn, other metals), which may reduce reactivity by removing elements from solution or increase reactivity by fostering mineral surface catalyzed reactions. The degree to which microbes influence and are influenced by these mineralization reactions is also an exciting avenue of research. Complete discussion of important considerations for calculation of energetic yields of the range of redox couples involving nitrogen, including those involving biomolecules, is beyond the scope of this chapter (for more complete treatment readers are referred to [31–34]).

8.3 Microbially catalyzed nitrogen cycling processes As described above, the abundance of aqueous inorganic N species in low temperature mixed fluids is highly variable and influenced by disequilibrium arising from mixing of hot NH+4 bearing and NO−3 free endmember fluids with cold NH+4 poor and NO−3 bearing seawater. Superimposed on this mixing of contrasting fluids is the distribution and abundance of N species, which is also regulated by a suite of microbially catalyzed processes (󳶳 Fig. 8.4). Given temperatures hospitable to life and the absence of O2 (after its consumption by microbial respiration and/or chemical reaction with reduced species such as Fe(II) or sulfide), the reduction of NO−3 , the most oxidized N species, is energetically favorable and rapidly catalyzed, producing NO−2 . This respiratory process of NO−3 reduction, catalyzed by the enzyme NO−3 reductase (nar), represents the first step before a key branching point in the N cycle. As shown in 󳶳 Fig. 8.4, the fate of the product NO−2 can be reduction to either gas phase or aqueous phase species, representing an important shunt for either N loss from the ecosystem (as gaseous NO, N2 O, or N2 ) or N recycling and retention (as NH+4 ). The continued respiratory and stepwise reduction of NO−2 to NO (by NO−2 reductase [nir]), to N2 O (by nitric oxide reductase [nor]) and ultimately to N2 (by nitrous oxide reductase [nos]) is known as denitrification. Denitrification can be coupled with either heterotrophic oxidation of organic carbon or with chemoautotrophic oxidation of species including H2 , Fe(II) or reduced sulfur species [11]. In contrast, the reduction of NO−2 to NH+4 , representing catalysis of a single six-electron transfer reaction by a different form of NO−2 reductase (nrf ), functionally acts to produce NH+4 , which is readily utilized by the microbial community. This process is known as ‘dissimilatory nitrate reduction to ammonium’ – or DNRA – and has also been shown to be catalyzed by both heterotrophic and chemoautotrophic microbes. A third reductive fate for NO−2 is catalyzed by the autotrophic process known as anammox, whereby the reduction of NO−2 is coupled to the anaerobic oxidation of NH+4 to produce N2 gas, thus representing an additional loss of fixed N from the ecosystem. Reduced N species, including NH+4 and NO−2 , can also be used as sources of energy by a group of chemoautotrophic organisms catalyzing the two-step process of nitrification. The first step in the oxidation of NH+4 to NO−2 , facilitated by NH3 -oxidizing bac-

8.3 Microbially catalyzed nitrogen cycling processes |

187

NO3-

V

Oxidation state

IV NO2-

III

NO2-

II I 0

-III

N2O N2

NH4+

Org N

Fig. 8.4: Schematic of microbially mediated nitrogen cycling processes discussed and the genes that encode for the enzymes that catalyze them

teria and/or archaea, is catalyzed by the enzyme ammonia monooxygenase (amo), therefore requiring the presence of molecular O2 . Similarly, NO−2 can be oxidized to NO−3 by chemoautotrophic NO−2 -oxidizing bacteria using the enzyme nitrite oxidoreductase (nxr), also in the requisite presence of O2 . Finally, acquisition of N for building biomass also comprises important avenues for N flow through these ecosystems (󳶳 Fig. 8.4). Microbial communities have evolved a number of pathways for assimilation of nitrogenous species for biosynthesis, all of which ultimately involve incorporation of N in its most reduced form. Heterotrophic assimilation of N in the form of organic matter (particulate or dissolved) generally involves the use of proteinases for breaking down larger molecules, followed by direct assimilation of N in the form of amino acids. Deamination of peptides can also release NH+4 directly during microbially catalyzed organic matter decomposition or ‘remineralization’. A wide range of microbes is able to assimilate NH+4 directly, first by using active NH+4 transport (amt) across the cell membrane [35] followed by funneling into either the glutamate dehydrogenase pathway (gdh) or glutamine synthetase (gln)–glutamate synthase (glt) pathway (GS-GOGAT) [36–38]. Nitrate can also be directly assimilated after being transported into the cell, reduced to NO−2 by an assimilatory NO−3 reductase (nas) and then fully reduced to NH+4 by assimilatory NO−2 reductase (nirB) [39–41]. In this same way, environmental NO−2 , although typically low in abundance, can also be directly assimilated. Finally, microorganisms known as diazotrophs have evolved to assimilate N2 using the highly specialized nitrogenase enzyme (nif ) that enables breaking of the triple bond of N2 and ultimately production of NH+4 for biomass synthesis [42, 43]. Thus, microbial communities have evolved a number of widely adaptable assimilatory pathways that ultimately aid in supporting high

188 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

levels of productivity in the energy rich contexts of hydrothermal vent systems. In contrast to the large expanse of the sunlit surface ocean, where uptake by photosynthetic organisms rapidly depletes dissolved inorganic forms of N, access to nitrogen is not generally considered a limiting nutrient in the deep sea. Nevertheless, an ecosystem level understanding of energy flow as well as underlying constraints on life in these environments remain fundamentally linked to the distribution, transfer and transformation of these nitrogen species.

8.4 Stable isotopes as indicators of microbial processes Among the available tools for investigating the occurrence and extent of environmental processes, stable isotopic composition of nitrogen (and oxygen) can provide uniquely integrated perspectives on the relative activity of various nitrogen cycling processes. In the environment, operative cycling pathways are difficult to discern directly through the use of concentration measurements alone. However, stable isotopes act uniquely to integrate processes over space and time, providing more information than can be gained from more conventional geochemical techniques. Owing to the fact that rates of chemical and biological transformations occur slightly faster for isotopically lighter species (e.g. molecules with 14 N and 16 O react slightly faster than those containing 15 N or 18 O respectively), changes in the ratio of heavy to light isotopologues reflect the activity of processes distinct from physical mixing. Standard stable isotope delta notation (δ) converts these isotopic ratios (‘15 R’, e.g. 15 N/14 N) of a sample relative to a standard into a metric in parts per thousand, or ‰, via the definition (for N): δ15 N = 1, 000 ⋅ [(15 Rsample /15 Rstandard ) − 1], where the standard for nitrogen isotopes is the 15 R of N2 in air. The same general definition is applied to oxygen, whereas ratios are reported relative to the 18 R of Vienna Standard Mean Ocean Water (VSMOW). From an ecosystem perspective, the bulk δ15 N of living biomass ultimately reflects the δ15 N of the flux-weighted source of N being incorporated at the base of the food web, with distinctive enrichments occurring between trophic levels [44, 45]. In the case of primary production by diazotrophic organisms, because N2 fixation by the molybdenum-containing nitrogenase imparts very little isotopic fractionation, microbial biomass generated and/or supported by diazotrophic organisms is generally very close to that of the source N2 (e.g. ∼ −2 to 0‰; [46, 47]). In contrast, because the δ15 N of deep sea NO−3 is generally closer to 5‰, microbial biomass ecosystems subsisting largely on a NO−3 source may have distinctive δ15 N values. This compositional distinction has indeed been used to draw conclusions about the possible importance of N2 fixation in a variety of marine ecosystems, including hydrothermal vents (e.g. [48]). As the result of the isotopic fractionation during most enzymatically catalyzed redox reactions, transformation processes often impart unique and distinguishable

8.4 Stable isotopes as indicators of microbial processes

| 189

patterns that can be used to identify the relative roles of certain processes and/or contributions of distinct sources. For example, the isotopic difference between two pools (e.g. NO−3 and NO−2 ) may reflect transformation of one reactant pool into another product pool. Under steady state conditions, the difference between the δ15 N of the reactant and product pools is approximately equal to the magnitude of the kinetic isotope effect (15 ε), where 15 ε = (15 α − 1) ⋅ 1, 000 and 15 α is equal to the ratio of the rate constants for each isotopologue (15 α =14 k/15 k). Thus, in general, as a reactant pool is consumed its δ15 N or δ18 O value will increase in direct relation to the strength of the isotope effect (15 ε) and the relative fraction of reaction progress. This is expressed as: δ15 Nreactant = δ15 Ninitial +15 ε ∗ fcon where fcon is the fraction of reactant consumed. Under closed system dynamics (where the reactant is not continuously replenished) the evolution of the reactant pool follows Rayleigh distillation dynamics, which can be approximated as: δ15 Nreactant = δ15 Ninitial −15 ε ∗ ln(frem ) and where frem is the fraction of reactant remaining. Thus, given some knowledge or assumptions about concentrations and expected isotope effects for a given process, the relative extent of a reaction can be estimated by these approaches. Alternatively, given that isotope effects may be stronger for one process than another (e.g. consumption of NO−3 by assimilation is generally thought to have a lower 15 ε than its consumption by denitrification), apparent values of 15 ε can be calculated and used to approximate the relative contribution of more than one process. For example, under diffuse flow venting, where seawater is shallowly entrained, NO−3 consumption may be incomplete before it is returned to the seafloor, meaning the expression of isotopic fractionation by consumption processes will be apparent. Although direct observations are few, depletions of NO−3 have been ‘routinely observed’ in diffuse flow environments [14]. This incomplete consumption often leaves some fraction of NO−3 remaining within the low temperature fluids, the typically elevated isotopic composition of which reflects the nature of active subsurface processes (e.g. [13, 30]). Finally, in examining shifts in stable isotopic composition under these settings, mixing processes also need to be carefully considered. Mixing of two fluids containing different concentrations and stable isotopic compositions results in hyperbolic mixing curves (curved lines in 󳶳 Fig. 8.5). Thus, a diffuse fluid sample (yellow circle in 󳶳 Fig. 8.5) having a lower NO−3 concentration and higher δ15 N value may result from a singular consumption process (having a 15 ε of 10‰ in this case). Alternatively, it could also result from mixing of background seawater with a more highly enriched pool (stemming from a consumption process having a 15 ε of 25‰ in this example: curved lines in 󳶳 Fig. 8.5). Thus, great care must be taken to account for dilution/mixing processes (using conservative elements such as Mg) when quantitatively examining the stable isotopic distribution of N species.

190 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

Fig. 8.5: Illustration of processes influencing stable isotopic composition (of NO−3 in this case), including biological consumption and mixing. Background seawater (blue square) contains 30–40 µM NO−3 , while high temperature endmember vent fluids contain zero NO−3 . Thus, mixing of these two fluids will result in dilution of NO−3 concentrations, yet no change to its N isotopic composition (red triangle). In contrast, mixtures of vent fluid and seawater under low temperature conditions (yellow circle) may reflect the influence of microbial processes (including those consuming and/or producing NO−3 ). Under open system conditions (see text), the isotopic fractionation during NO−3 consumption will lead to a linear increase in the δ15 N values as NO−3 consumption proceeds, the slope of which will correspond to the strength of the isotope effect (15 ε). Thus, the sample represented by the yellow circle may be interpreted as consumption of NO−3 having an isotope effect of ∼10‰. Alternatively, the sample may also represent the integrated influence of isotopic mixing processes. Because mixing of two fluids containing NO−3 with different concentrations and δ 15 N values will be hyperbolic (mixing curves), the sample could also represent mixing between background seawater and a more highly fractionated and consumed NO−3 pool (green diamond)

Examination of coupled isotope dynamics (e.g. δ15 N and δ18 O in NO−3 , NO−2 or N2 O) can provide an even more powerful approach for constraining the relative roles of simultaneously occurring cycling processes. Specifically, while for some processes the dynamics of N and O isotopes are tightly coupled – e.g. both 15 N/14 N and 18 O/16 O are influenced by a transformation catalyzed by a single enzyme – other N cycling processes involve separate N and O pools and reaction pathways, effectively decoupling these isotope systems and shedding light on the relative roles of different N cycling processes. For example, the enzymatic reduction of NO−3 by NO−3 reductase (whether dissimilatory or assimilatory) across a wide variety of microbes has been demonstrated to result in a closely coupled evolution of δ15 N and δ18 O of the remain-

8.4 Stable isotopes as indicators of microbial processes | 191

ing NO−3 along a 1 : 1 slope, e.g. 18 ε : 15 ε ∼ 1 [49, 50]. Notably, however, the reduction of NO−3 by an auxiliary periplasmic NO−3 reductase (nap) appears to exhibit a lower 18 ε : 15 ε relationship near 0.6 [49, 51]. Similarly, the reduction of NO− by NO− reduc2 2 tase has also recently been demonstrated to exhibit a tight coupling of the 18 ε and 15 ε, albeit with distinctly different coupling ratios depending on the type of NO−2 reductase (18 ε : 15 ε ∼ 0.75 for copper-containing nirK and 18 ε : 15 ε ∼ 0.1 for iron-containing nirS; [52]). Thus, the reductive consumption of oxyanions by enzymatic catalysis generally imparts distinct coupled isotope fractionation, which can be used to constrain the contribution of these processes in the environment. In contrast to the coupled nature of N and O isotopes during oxyanion reduction, nitrification, i.e. the production of NO−2 and NO−3 , represents a distinctive decoupling of N and O isotope systems with differential impact on the respective δ15 N and δ18 O of these pools. Ammonia oxidation, the strongly fractionating first step in the oxidation of NH+4 to NO−2 , initially produces NO−2 with a δ15 N much lower than that of the initial NH+4 substrate, with N isotope effects being similar for both bacteria and archaea [53– 55]. The second step of nitrification, NO−2 oxidation, has an unusual inverse isotope effect, resulting in NO−3 having a δ15 N greater than its substrate [56], however, this effect is only expressed in the NO−3 pool if NO−2 is allowed to accumulate. Hence, given incomplete consumption of the NH+4 substrate and complete consumption of the NO−2 intermediate, the δ15 N of the produced NO−3 will likely be much lower than ambient deep ocean NO−3 . In contrast, if oxidation of NH+4 and NO−2 is complete, then the new NO−3 will have a δ15 N value equal to the substrate NH+4 [57]. Ammonium sources in hydrothermal vent systems are not well constrained, but can originate from a variety of sources including pyrolysis of organic matter in buried sediments [58] or in the subsurface biosphere, abiotic interactions between NO−3 and metal oxides and/or sulfides [59], and high T abiotic reduction of N2 to NH+4 [4, 60, 61]. By comparison, the O atoms appended to the product NO−2 pool during NH+4 oxidation derive from both molecular O2 (50%) and water (50%); the final O atom appended to the NO−3 pool during NO−2 oxidation also derives from water [62, 63]. Recent work has also shown that the incorporation of each O atom is accompanied by distinctive kinetic isotope effects [64–66]. Additionally, intermediate NO−2 undergoes oxygen isotopic equilibration with ambient water, a process that occurs especially rapidly at pH < 6 [67, 68]. Thus, given conditions promoting isotopic equilibration (e.g. low pH, higher temperature), the contribution of O atoms during NH+4 oxidation to NO−2 may be effectively ‘reset’ by this process. At any rate, the δ18 O values of newly produced NO−2 and NO−3 are clearly independent of the N source and therefore provide a second dimension for resolving multiple processes.

192 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

8.5 Genetic evidence for nitrogen cycling processes at vents In addition to geochemical and isotopic lines of evidence, genetic data have emerged as incredibly valuable tools for demonstrating the potential and activity of several functional genes central to microbial N cycling, including studies of hydrothermal vent ecosystems. As shown in 󳶳 Fig. 8.4, transformations of the nitrogen cycle are largely catalyzed by distinctive and well-characterized enzymes. When an organism requires one of these proteins, the genetic code contained in its DNA for directing enzyme synthesis is upregulated, resulting in transcription into messenger RNA (mRNA), ultimately directing synthesis through ribosomal translation. The functional genes encoded in organismal DNA represent the genetic potential of a specific metabolic process. In light of this, quantification of functional gene abundances in the environment (through quantitative polymerase chain reaction or qPCR) has been used to infer the potential importance of various processes. Notably, however, drawing relationships between gene abundance and catalyzed rates of particular transformations is not straightforward, in particular since the presence of a specific gene reveals nothing about whether it is being actively upregulated in response to environmental cues and/or physiological triggers. More recent work has focused on quantifying the active transcription of these genes as reflected in the abundance of mRNA (via reverse transcriptase or RT-qPCR). As genetic sequencing techniques have become more affordable and bioinformatics approaches more user friendly, the emerging use of ‘omics’ tools (metagenomics, metatranscriptomics, proteomics) has also begun shedding light on the relative abundance of genes and gene expression for nitrogen based transformations in hydrothermal vent ecosystems. These approaches result in exceptionally large datasets of the genetic potential or the transcribed activity of microbial communities from the environment, from which specific investigation of the occurrence and activity of N cycling genes can be interrogated. Such approaches will undoubtedly mature in the near future through combination with metabolic modeling, incorporating the function and properties of genes, enzymes and pathways together with improved understanding of microbial physiological controls. With its multitude of transformations, a better understanding of the nature of nitrogen cycling in hydrothermal vents will undoubtedly follow. In the sections below, we discuss in more detail the different N transformations in the subsurface biosphere of hydrothermal vents. Here we aim to summarize the current state of understanding of N cycling processes occurring in the context of submarine hydrothermal vent ecosystems, highlighting novel features and characteristics as well as gaps that should be addressed by future work. A variety of approaches have been employed in the study of nitrogen cycling; below we have organized our synthesis according to processes, drawing from the range of techniques aimed at elucidating their presence, rate and/or role in nitrogen speciation.

8.6 Nitrogen fixation

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8.6 Nitrogen fixation Biological nitrogen fixation involves the enzymatic reduction of dissolved N2 to NH+4 for assimilation and biomass synthesis, a reaction catalyzed by the nitrogenase enzyme complex (󳶳 Fig. 8.6) [43]. The ability to fix N2 is limited to a highly specialized group of microbes and is conventionally considered to be well suited for environments where the abundance and availability of other forms of fixed N (i.e. NO−3 , NO−2 , NH+4 , DON and particulate organic nitrogen [PON]) are limiting since catalysis by nitrogenase is energetically intensive. Classic environments in the ocean include surface layers of oligotrophic gyres and other oligotrophic waters [69–71]. However, growing evidence also suggests that other environments that are less often considered may also harbor active N2 fixation, including benthic sediments, hydrocarbon seeps and hydrothermal vents [72–75]. One of the earliest lines of evidence supporting active N2 fixation in hydrothermal systems came from nitrogen isotopic composition (δ15 N) of vent animals from the Galapagos and 21°N vents in the eastern Pacific [48]. The biomass of these animals (notably vestimentiferan worms and clams) displayed low 15 N/14 N ratios, which was interpreted as evidence that they were primarily supported by a food web heavily supplied by N sourced from N2 fixation. Similar lines of evidence from the study of carbonate samples collected from active chimneys at the serpentinite hosted Lost City hydrothermal field [76] also reported similarly low δ15 N measurements from the organic nitrogen fraction within the carbonate samples (as low as 0.1‰), again interpreted as reflecting an apparent significance of N2 fixation in this system.

Nitrogen Fixation Nitrogenase complex Fe

16 ATP protein 8 Fdxreduced 8 Fdxoxidized 16 ADP + 16 Pi

Fe-Mo protein

N2 + 8H+

NifK NifD NifH NifD NifK 2 NH3 + 2 H2

Fig. 8.6: Enzymes comprising the nitrogenase complex for microbial nitrogen fixation. The nitrogenase complex is comprised of multiple proteins in which 16 ATPs are used to shuttle electrons from a reducing agent (e.g. ferredoxin (Fdx)) and to drive reduction potential through the Fe protein (a homodimer encoded by the nifH gene) and through the Fe-Mo protein (a heterotetramer encoded by nifD and nifK) for the ultimate reduction of N2 to NH3 . The process also catalyzes the production of H2 , which is generally utilized by diazotrophic organisms for energy through an uptake hydrogenase (not shown)

194 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

The detection of gene sequences encoding for the nitrogenase enzyme (nifH) has also revealed the potential for active N2 fixation in hydrothermal vents. The presence of nifH genes in subseafloor diffuse hydrothermal vent fluids and deep seawater was reported from Marker 33 on Axial Volcano and near Puffer on the Endeavour Segment of the Juan de Fuca Ridge [77]. The nifH genes found in the vent fluids and from the surrounding bottom seawater exhibited distinctive phylogenies, with the latter group exclusively related to those of methanogenic archaea and the former mainly related to those of anaerobic sulfate reducers and Clostridia. It was also shown that the seawater from the Axial Valley of the Endeavour Segment on the Juan de Fuca Ridge contained archaeal nifH genes [78]. Indeed, quantitative detection of nifH genes on both mature and newly forming sulfide chimneys suggests that organisms catalyzing N2 fixation may be important ecological players over a broad spectrum of ecosystem conditions [79]. Physiological studies have also been conducted on N2 fixing organisms isolated from various deep sea hydrothermal vent sites. Early work demonstrated that representative heterotrophic and chemolithotrophic microorganisms capable of N2 fixation could be cultured from diffuse hydrothermal fluids sampled at the Galapagos Spreading Center [80]. A hyperthermophilic N2 fixing archaeal strain (FS406-22) isolated from Marker 33 (Juan de Fuca Ridge) was shown to have an optimal growth temperature of 90 °C, with active N2 fixation confirmed by both incorporation of 15 N2 and expression of nifH genes [75]. Two other hyperthermophilic methanogenic archaeal isolates, Methanocaldococcus (Mc 1–85N) and Methanothermococcus (Mt 5–55N), were also shown to be able to assimilate both N2 and NH+4 , but not NO−3 , when grown at temperatures of 85 °C and 55 °C respectively [81]. Strikingly, these N2 fixing methanogenic isolates produced biomass that was more depleted in 15 N with larger isotopic fractionation (∼ −4‰) than previously reported for N2 fixing photosynthetic microbes (e.g. near 0‰; [46, 82]). This apparently larger N isotope fractionation in N2 fixing methanogenic strains has implications for the interpretation of natural abundance of N isotopes in food web and N cycling in hydrothermal vents. Since the δ15 N derived from N2 fixation is generally considered to be close to 0‰, in systems where these methanogenic N2 fixers are active, the fixed N for this food web could be skewed toward more negative δ15 N values. Thus, if a primary source of N derives from N2 fixation by these methanogens, the δ15 N value of the microbial biomass N produced by this input could be as low as −4‰. Similarly, the discovery that so-called alternative nitrogenases (e.g. V and Fe forms) may have larger isotope effects than the Mo form raises the possibility that estimation of N2 fixation by N isotopes may be more complex than previously appreciated in environments where these are active [83]. The role of alternative nitrogenases in the redox dynamic context of hydrothermal vents, however, remains entirely unexplored. Without consideration of these potential sources of lower δ15 N, the relative contribution of N2 fixation could be overestimated.

8.7 Nitrogen assimilation

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8.7 Nitrogen assimilation All organisms need nitrogen for biosynthesis and growth, making assimilation of an exogenous source of nitrogen central to their ecological success. Aside from the highly specialized microorganisms capable of N2 fixation, microbial assimilation of nitrogen for biomass synthesis generally occurs either through heterotrophic decomposition of large organic molecules and assimilation of smaller N-bearing organic molecules such as amino acids, or by direct assimilation of inorganic forms of nitrogen such as NH+4 or NO−3 (󳶳 Fig. 8.7). Here we summarize investigations into the microbial ability to assimilate these inorganic forms. Early work on N assimilation processes in deep sea vent organisms was conducted on the mouthless, gutless vestimentiferan worm Riftia pachyptila, which is solely reliant on the food supplied by its endosymbiotic bacteria [84]. Isotope tracer experiments (15 N) targeting NH+4 and NO−3 assimilation by these endosymbionts of Riftia

Fig. 8.7: Enzymes involved in the assimilation or inorganic forms of nitrogen. Extracellular forms of nitrogen such as NO−3 , NO−2 and NH+4 can be actively transported across cellular membranes for assimilation. Nitrate reduction to nitrite (1) can be catalyzed by nitrate reductase (encoded by either nasA for assimilatory nitrate reductase or narG for the respiratory nitrate reductase, both producing NO−2 as the product). Intracellular nitrite reduction (2) to ammonium is catalyzed by nitrite reductase (encoded by nirB). Intracellular ammonium assimilation can occur through two mechanisms, largely depending on substrate availability, both producing glutamate as a substrate for major biosynthesis pathways. The glutamate dehydrogenase (3) pathway (Gdh encoded by gdh) is the low affinity pathway, wherein the α-ketoglutarate and NH+4 are used to form glutamate directly. Under lower NH+4 availability, the combined action of the glutamine synthetase (GS encoded by glnA) and glutamate synthase (GOGAT encoded by gltB) pathways (4) act to produce glutamate for biosynthetic pathways

196 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

pachyptila from 13°N on the East Pacific Rise revealed that N assimilation was primarily supported by uptake of NO−3 rather than NH+4 [84]. A technically sophisticated follow-up study on tube worms collected from the same sites and employing flowthrough, pressurized aquaria containing the entire organism also demonstrated significant preferential NO−3 uptake relative to NH+4 , confirming NO−3 as the main N uptake source in these endosymbionts [85]. These results contrast with the conventional paradigm for inorganic N uptake where NH+4 is thought to be favored over the uptake of NO−3 . However, these findings were interpreted as reflecting expectations based on ambient concentrations of NH+4 (low) and NO−3 (high) found in the habitat of these animals. N assimilation has also been investigated by quantification of the expression of genes associated with N assimilation (󳶳 Fig. 8.7). For example, a study of N acquisition by symbionts of the tubeworm Ridgeia piscesae also reported expression of genes encoding for enzymes required for both assimilatory NO−3 reduction (nas) and NH+4 assimilation (gln, glt) from sites at both the Axial Volcano and Main Endeavour vent fields of the Juan de Fuca Ridge [86]. Similar dynamics were also observed for N assimilation by symbionts of hydrothermal vent snails, Alviniconcha, at the Eastern Lau Spreading Center [87]. In both studies, expression of genes responsible for assimilatory NO−3 reduction as well as NH+4 assimilation suggested that symbionts were metabolically poised to take up NH+4 either from the surrounding environment or from the reduction of NO−3 to NH+4 , thought to be mediated by a periplasmic NO−3 reductase (nap). Thus, in contrast to previous studies of Riftia symbionts from the East Pacific Rise, where NH+4 levels are quite low [84, 85], N assimilation in other hydrothermal vent symbioses from environments hosting higher NH+4 in fluids, or perhaps otherwise modulating NH+4 acquisition (e.g. pH speciation considerations), appear to reflect a more versatile suite of N assimilation strategies. Unfortunately, far fewer studies have been conducted on microbial N assimilation by nonsymbiont communities at hydrothermal vents. Within the neutrally buoyant plume of the Juan de Fuca Ridge, which emanates from high temperature endmember fluids with very high NH+4 concentrations > 1 mM [58], NH+4 assimilation rates were substantially enhanced (up to 26.4 nM d−1 ) in comparison to background seawater (1.5−5.0 nM d−1 ) [88]. These elevated NH+4 assimilation rates were estimated to account for at least 47% of total net NH+4 removal rates. Although not necessarily definitive evidence, stable isotopic measurements of both NO−3 and NH+4 have also suggested active biological removal of these compounds, possibly from microbial assimilation, in studies of both the Juan de Fuca Ridge region [13] as well as the Loihi Seamount in Hawaii [30].

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8.8 Denitrification and anammox Bioavailable (or fixed) N is removed from ecosystems through conversion to gaseous forms of N by either denitrification or anammox, processes representing important sink terms for nutrient budgets in all known ecosystems. Denitrification refers to the stepwise reduction of NO−3 to NO−2 , NO, N2 O and finally N2 (󳶳 Fig. 8.8). In most regions of the ocean, denitrification is thought to be performed by anaerobic heterotrophic bacteria coupling the respiratory reduction of NO−3 to the oxidation of organic matter [89, 90]. In hydrothermal vent habitats however, several studies have indicated that denitrification may be substantially catalyzed by autotrophic organisms, wherein NO−3 reduction is coupled to oxidation of H2 , reduced sulfur species (S0, S2 O2− 3 , H2 S) or reduced iron (Fe(II)) [30, 91, 92]. We note that the co-occurrence of abiotic NO−3 (and more likely NO−2 ) reduction to N2 in low temperature vents is also possible, for example coupled to Fe(II) oxidation (e.g. ‘chemodenitrification’) [93], although little is known about the possible importance of this process in hydrothermal vent systems.

Denitrification Periplasm H2O + N2

2H++ NO31 NapA H2O + NO2 H

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NorC NorB

Q H2O + NO2 NAD+

NADH + H+

2H++ NO3- NarH 1 NarG

Cytoplasm

Fig. 8.8: Enzymes involved in denitrification. Reduction of nitrate to nitrite (1) is catalyzed by a membrane associated respiratory nitrate reductase (Nar encoded by narGHI). Alternatively, reduction may also be catalyzed by a periplasmic nitrate reductase (Nap encoded by napA), though this reaction may not be directly involved in cellular respiration. In the context of denitrification, both processes produce nitrite that may be further utilized in the respiratory chain. Nitrite reduction to nitric oxide (2) is catalyzed by either a Fe-containing nitrite reductase (encoded by nirS) or a Cu-containing nitrite reductase (Nir encoded by nirK). Nitric oxide reduction (3) to nitrous oxide is catalyzed by nitric oxide reductase (a cNOR encoded by norCB), while the product nitrous oxide may be reduced (4) to dinitrogen by nitrous oxide reductase (Nos encoded by nosZ). Denitrification can be considered as any case in which the reductive production and release of any N-bearing gas (e.g. nitric oxide, nitrous oxide or dinitrogen) occurs. Supply of electrons is maintained by shuttling via the quinone pool (QH). DH refers to the NADH dehydrogenase complex

198 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

Anammox Cytoplasm

NO35

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NAD+

Nxr H2O

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2 +

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+3 e-

H

ADP + Pi 4e-

HZS

+

H+

NirS NirK

1

H+

H2O

3

2

4

N2 + 4 H +

N2H4

Anammoxosome Fig. 8.9: Enzymes involved in the anaerobic oxidation of ammonium (anammox). Nitrite reduction (1) is catalyzed by either the hemeprotein nitrite reductase nirS or the Cu-containing nirK protein, the nitric oxide product of which is combined with ammonium (2) by way of a hydrazine synthesis enzyme (HZS). Following hydrazine production inside the anammoxosome (3), hydrazine is converted to dinitrogen (4) by hydrazine hydrolase (HZH). Note that energy metabolism also involves the production of nitrate from nitrite (5) by action of a nitrate:nitrite oxidoreductase (Nxr). DH refers to the NADH dehydrogenase complex

By comparison, anammox is the enzymatically catalyzed oxidation of NH+4 coupled to the reduction of NO−2 , producing N2 gas (󳶳 Fig. 8.9), and is carried out by anaerobic autotrophic bacteria [94–97]. To date, anammox activity has been identified in diverse hydrothermal vent habitats (active smoker, animal holobionts and microbial mats) on the Mid-Atlantic Ridge [98] as well as within diffuse fluids of the Juan de Fuca Ridge [13]. Several lines of evidence confirm bioavailable N loss in the microbially amenable context of diffuse hydrothermal vent fluids. First, NO−3 concentrations in diffuse fluids often fall below conservative mixing lines between NO−3 free, high temperature (low Mg2+ ) vent fluids and NO−3 rich, low temperature (high Mg2+ ) seawater (󳶳 Fig. 8.2). Such nonconservative behavior has been recorded at hydrothermal vents of the Juan de Fuca Ridge [14, 29] as well as the Loihi Seamount, Hawaii [30]. Notably, in contrast to the sites on the Juan de Fuca Ridge, fluids from Loihi Seamount were low in sulfide and high in dissolved iron, representing a unique model of low temperature venting, biogeochemical reactions and microbial community. Second, the consumption of NO−3 was also reflected in the co-occurring increase of the δ15 N and δ18 O of NO−3 in diffuse fluids (with respect to background seawater NO−3 ), concomitant with decreased NO−3 concentrations [13, 30] (󳶳 Fig. 8.5). Since simple mixing of seawater with high temperature NO−3 free fluids, which would lead to decreases in concentrations by dilution, would not lead to changes in the isotopic composition, these elevated isotopic com-

8.8 Denitrification and anammox | 199

positions unequivocally reflect isotopic fractionation by NO−3 -consuming processes in low temperature fluids. Third, in addition to observed changes in concentrations and isotopes, direct measurement of denitrification rates in ex situ incubations has also been used to provide insight into the metabolic potential of these environments. Denitrification rates of up to ∼1 µmol N l−1 day−1 were measured in 15 N-labeled incubations in diffuse vent fluids of the Juan de Fuca Ridge, corroborating the inferred activity of denitrification based on NO−3 stable isotope data [29]. In contrast, anammox rates measured in the same study at the Juan de Fuca sites were markedly lower, never reaching higher than ∼5 nmol N l−1 day−1 . In comparison, anammox rates of up to 30 nmol N l−1 day−1 were documented in a hydrothermal chimney along the Mid-Atlantic Ridge [98], suggesting that while rates of anammox may be low relative to denitrification in hydrothermal systems, microbial communities capable of catalyzing anammox may be widespread, in particular where anaerobic conditions persist. In contrast to ex situ experiments, in which fluid samples are brought up from the seafloor in specialized titanium samplers and incubations are conducted in shipboard laboratories, more sophisticated approaches enabling in situ incubation experiments can provide additional insights [91, 99, 100]. Recent in situ experiments conducted at ambient pressure and temperature permitted the selective amendment of different electron donors (H2 and H2 S) and acceptors (O2 and NO−3 ) to be applied to fluids collected directly on the seafloor [91]. Rates of denitrification measured using these in situ incubation systems at a diffuse flow site at the East Pacific Rise (‘Crab Spa’) ranged from 27 to 398 µmol N l−1 day−1 [91], substantially higher than those reported for the Juan de Fuca site [29]. Variations in native denitrifying microbial communities and their metabolic capacity notwithstanding, a number of possible experimental aspects may also have contributed to observed differences in rates between these two studies. These include: 1) the ten-fold higher NO−3 amendments used during the in situ incubations that might have stimulated denitrifying bacterial activity; 2) decompression and/or experimental manipulation of fluids, which may impact microbial activity; and/or 3) the helium purging or preincubation of fluid samples, which may have depleted electron donors (e.g. H2 , H2 S) and consequently decreased autotrophic bacterial activity. Reported denitrification rates, though sparse, vary widely depending on several environmental factors. In general, denitrification rates in diffuse hydrothermal vent fluids are substantially higher than rates reported for the open ocean (up to 9 nmol N l−1 day−1 ) or for coastal oxygen-deficient zones (up to 33 nmol N l−1 day−1 ) [89]. The denitrification rates measured at the diffuse flow site at the Juan de Fuca Ridge [29] were more similar to reported chemolithoautotrophic denitrification rates in the Namibian shelf waters (600 nmol N l−1 day−1 ) [101]. On the other hand, the higher rates reported from the in situ incubations at the East Pacific Rise [91] are more similar to denitrification rates of up to ∼30 µmol N l−1 day−1 measured in cold seep sediments of the Gulf of Mexico [102] and Guaymas Basin hydrothermally altered sed-

200 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

iments [103]. Notably, rates reported for hydrothermal sediment incubations at Guaymas Basin were as high as the highest reported rates from estuarine systems [103]. These observed ranges in denitrification rates reflect potentially wide variations in the types of environments sampled, the absolute numbers of cells being incubated, and the concentrations and nature of various available substrates (organic carbon, H2 S, H2 , etc.). Indeed, the presence of sulfide, which is commonly high in fluids from ridges and sedimented hydrothermal systems, is known to greatly influence the nature of nitrogen cycling pathways [104–106]. Sulfide can inhibit denitrification [104] or promote N2 O production by denitrification [104, 107], while also serving as an important energy source for autotrophic denitrification by nitrate-reducing sulfuroxidizing bacteria, especially Epsilonproteobacteria (e.g. [11], and reference therein). Nevertheless, these studies are beginning to provide a foundation for understanding the nature of reductive processes in deep sea hydrothermal settings. For example, the apparently low anammox rates in comparison to denitrification may stem from the strictly anaerobic nature of anammox bacteria, relative to denitrifying bacteria, which are often only facultative anaerobes. Perhaps the often dynamic nature of hydrothermal fluid venting, which is subject to variations imposed by tidal cycles, mineralization and/or macrofaunal intervention, gives rise to a more variable zonation in redox conditions, for which denitrifying bacteria are better adapted. It could also be that slowly growing anammox bacteria with a doubling time of approx. two weeks [108] are simply outcompeted by faster growing autotrophic denitrifiers with a doubling time as short as 1.5 hours [109]. Regardless of the differences in absolute rates, these studies suggest that denitrification likely comprises the major fixed N-loss process in the subsurface biosphere of hydrothermal vents. Nevertheless, it is also clear that a high proportion of nitrate reduction may also result in the conservation of N through processes such as DNRA and NO−3 assimilation. Surveys of the microbial community through both culturing and culture independent molecular methods have also revealed the presence and activity of microbes mediating fixed N loss in hydrothermal vent habitats. Over 20 different chemolithoautotrophic NO−3 reducing organisms (using a variety of inorganic electron donors including H2 and reduced S compounds) have been isolated and cultured from hydrothermal vents ([11] and references therein), with some reducing NO−3 to N2 , [92, 110] and others reducing NO−3 to NH+4 [111–113]. This represents an important difference for ecosystem N budgets (see DNRA in Section 8.9 below). Nitrate accumulation of concentrations up to 4,000 higher than ambient levels have been documented in vacuoles of autotrophic sulfide-oxidizing bacteria (Beggiatoa sp.) from bacterial mats at Monterey Canyon cold seeps and Guaymas Basin hydrothermal systems [114]. It was suggested that Beggiatoa sp. could use this NO−3 as a terminal electron acceptor in respiration, producing either NH+4 or N2 as a waste product. Nitrate respiration coupled to sulfide oxidation was also found to occur in intracellular chemoautotrophic symbionts of the hydrothermal vent tubeworm Riftia pachyptila, with the end product being NO−2 or N2 (denitrification) [115].

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| 201

Epsilonproteobacteria are routinely identified as major components of hydrothermal vent chemolithoautotrophic bacterial communities [116–120]. Culture studies of Epsilonproteobacteria isolates from vent environments, including Sulfurimonas autotrophica, Sulfurovum lithotrophicum, Sulfurimonas gotlandica, and Sulfurimonas paralvinellae, have revealed the metabolic coupling of sulfur or hydrogen reduction with denitrification in the context of a multitude of hydrothermal vent habitats (artificial colonization devices, sediments, animals, diffuse fluids, plume) [10, 11, 120–122]. In addition, SUP05, a Gammaproteobacterium mediating sulfur and hydrogen oxidization coupled to denitrification in the ocean [123], has also been identified in studies of hydrothermal fluids and neutrally buoyant plumes [29, 124, 125]. Metagenomic and metatranscriptomic studies have also confirmed the presence and activity of key genes related to NO−3 reduction and denitrification (󳶳 Fig. 8.8) at vents. This includes dissimilatory NO−3 reductases (narG and napA), dissimilatory NO−2 reductases (nirS and nirK), nitric oxide reductase (norB) and nitrous oxide reductase (nosZ) (e.g. [79, 103, 126, 127]) (󳶳 Fig. 8.4). For instance, a high abundance of genes involved in the entire denitrification pathway were observed in the metagenome of a hydrothermal vent chimney from the Mothra field on the Juan de Fuca Ridge, which were mostly related to Beta- and Alphaproteobacteria [127]. Sequencing of a fosmid clone containing a narG fragment exhibited the highest identity with Thiobacillus denitrificans, a Betaproteobacteria mediating sulfur driven autotrophic denitrification. In contrast, no denitrification genes were identified in a sample collected from a carbonate chimney from Lost City on the Mid-Atlantic Ridge, implying that denitrifiers might not be prevalent in all vent communities [127]. Of the two types of NO−3 reductase linked to respiratory electron transport (the membrane bound NarG and the periplasmic NapA), the membrane bound NO−3 reductase NarG was found in most groups of vent Proteobacteria (Beta, Alpha, Gamma) [127, 128]. However, the periplasmic NO−3 reductase NapA was found to be highly conserved and widespread in the Epsilonproteobacteria, which generally represent the most abundant bacteria at vents [129]. Nap has been shown to be a high affinity NO−3 reductase and may therefore be better adapted for NO−3 depleted hydrothermal vent habitats [130]. Indeed, different temperature regimes (and perhaps more importantly the geochemical gradients they represent) may play an important role in the partitioning of prominent pathways controlling the biological fate of NO−3 . Most of the mesophilic (25 °C >T< 40 °C) Epsilonproteobacteria isolated from hydrothermal vents were shown to mediate NapA-catalyzed respiratory NO−3 reduction to N2 , while thermophilic (T> 40 °C) Epsilonproteobacteria appear to more typically produce NH+4 as a product of NO−3 respiration [129]. Indeed, higher temperature fluids also often contain higher concentrations of sulfide, which may be playing a key role in dictating the reductive fate of nitrate. A number of 16S rRNA sequences related to the anammox bacterial genera Candidatus Kuenenia and Scalindua have been amplified from a variety of hydrothermal vent environments including mussel gills and vent chimneys at sites on the MidAtlantic Ridge [98, 131]. The diversity and abundance of anammox bacteria has also

202 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents

been investigated using the anammox functional gene (󳶳 Fig. 8.9), hydrazine synthase (hzsA), in both cold seeps and hydrothermal sediments of the Guaymas Basin [131]. In general, most of the sequences were closely related to the genus Candidatus Scalindua, with a high interspecies diversity. Quantitative PCR amplification of a fragment of the hszA gene revealed a lower abundance of anammox bacteria in hydrothermal sediments compared to nearby cold seeps [131]. Anammox bacterial sequences of the genus Candidatus Scalindua were also retrieved in low temperature hydrothermal FeSi rich precipitates from the Lau Integrated Study Site in the Southwestern Pacific Ocean [132]. q-PCR analysis however showed that aerobic ammonia-oxidizing archaea were two to three orders of magnitude more abundant than anammox bacteria in these precipitates. Ladderane lipids, a molecular biomarker of anammox bacteria, have also been used to confirm the presence of anammox bacteria in chimney samples of the Mid-Atlantic Ridge [98] and sediments of the Guaymas Basin [131]. However, to the best of our knowledge, no 16S rRNA gene sequences related to anammox bacteria have been recovered from diffuse fluids, likely owing to their low abundance [29].

8.9 Dissimilatory nitrate reduction to ammonium (DNRA) In addition to denitrification, DNRA represents a competing and/or parallel NO−3 reduction process that operates to conserve fixed N within the ecosystem (e.g. N is not converted and lost as a gaseous phase N product). Several bacteria performing DNRA have been identified in vent environments with a few examples already briefly discussed above, i.e. Beggiatoa sp. [114] and many thermophilic Epsilonproteobacteria [129]. Thermophilic and chemolithoautotrophic bacteria capable of DNRA, Thermovibrio ammonificans sp. [129], Caminibacter mediatlanticus sp. [133], and Nautilia nitratireducens sp. [112], have been isolated from active hydrothermal vent chimneys on the East Pacific Rise and the Mid-Atlantic Ridge. Metatranscriptomic analyses have revealed some genes associated with DNRA within the chemoautotrophic symbionts of the snail Alviniconcha from hydrothermal vents of the Eastern Lau Spreading Center in the southwestern Pacific [87]. DNRA functional genes (󳶳 Fig. 8.10) have also been identified in the symbionts of the hydrothermal vent tubeworm Ridgeia piscesae [86]. In this study, all symbionts contained the periplasmic NO−3 reductase (nap) but the not the periplasmic NO−2 reductase (nrf ), the next step in the DNRA sequence [130]. An alternative pathway for DNRA was proposed in which nap-catalyzed NO−3 reduction in the periplasm was followed by NO−2 assimilation by the narK NO−2 /NO−3 transmembrane transporter. NO−2 was assumed to then be reduced to NH+4 and assimilated within the cytoplasm [86]. A number of environmental factors likely play key roles in determining the relative magnitude of DNRA relative to denitrification. Although generally lower than denitrification, rates of DNRA of up to 150 nmol N l−1 day−1 were measured using ex situ 15 N incubations in hydrothermal vent fluids of the Juan de Fuca Ridge [29]. Sulfide was

8.9 Dissimilatory nitrate reduction to ammonium (DNRA) |

203

Dissimilatory Nitrate Reduction to Ammonium (DNRA) Periplasm 2H++ NO3-

NO2- + 8H+ NrfA

1 NapA H2O + NO2 -

NapB

2 NH4+ + 2H2O

CymA MQ2 MQ Cytoplasm

Fig. 8.10: Enzymes involved in dissimilatory nitrate reduction to ammonium (DNRA). Reduction of nitrate to ammonium is initiated by the two-electron reduction of nitrate to nitrite (1), primarily thought to be conducted via the periplasmic nitrate reductase (Nap encoded by napA). The final six-electron reduction of nitrite to ammonium (2) is catalyzed by a nitrite reductase (Nrf encoded by nrfA). Electron supply is facilitated by menaquinones or similar intracellular redox shuttles and membrane bound cytochromes, such as the tetraheme cytochrome CymA

shown to play a central role in the relative rates of denitrification and DNRA in sediment incubations of hydrothermally altered sediments from the Guaymas Basin [103]. Production of NH+4 was also observed in fluids from the East Pacific Rise during in situ incubations of vent fluids at ambient temperature and pressure, especially when both NO−3 and hydrogen were added [91]. In fact, dissolved hydrogen availability was determined to be an important factor for DNRA, consistent with the view that DNRA is mostly supported by the oxidation of hydrogen. This is in contrast to chemoautotrophic denitrification in which either hydrogen or reduced sulfur may be used as an electron donor [11]. Furthermore, both the largest levels of NH+4 production from DNRA and the largest increase in cell numbers occurred in incubations performed at higher temperature incubations (50 °C as opposed to 24 °C) [91], also suggesting that DNRA may be a more important respiratory pathway at higher temperatures [11]. Indeed, it was shown that DNRA can occur at temperatures of up to 70 °C [112, 113, 133]. Notably, however, DNRA never represented more than 55% of the respired NO−3 in the in situ incubation experiments [91], reflecting the multiplicity of fates for NO−3 under reducing hydrothermal conditions. Although not the main focus of the present discussion, it is relevant to point out that oxidized forms of N are also abiotically reduced under high temperature conditions often encountered along hydrothermal flow paths. Indeed the presence of NH+4 in high temperature endmember fluids is often generally assumed to simply reflect abiotic reduction of bottom seawater NO−3 during convective transport through subsurface reaction zones. Studies have indeed suggested that NO−3 can be reduced abi-

204 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents otically to NH+4 with Fe(II) as a catalyst at temperatures from 22 to 200 °C (and clearly higher) [61, 134, 135]. Closer examination, however, suggests that subsurface reactions may be more complex than previously acknowledged. For example, NH+4 can often be depleted relative to conservative mixing with crustal seawater (󳶳 Fig. 8.2) [13], suggesting the operation of other loss processes. In the absence of biological consumption, the N isotopic composition of NH+4 in high temperature fluids, if derived from bottom seawater NO−3 , is expected to be equal to the isotopic composition of the original deep seawater NO−3 source (∼5–6‰). Indeed, δ15 N values of NH+4 in high temperature fluids from Axial Volcano were not significantly different from the δ15 N of background seawater NO−3 , consistent with the quantitative conversion of seawater NO−3 into NH+4 [13]. However, NH+4 concentrations in high temperature fluids were approximately only half those of the source NO−3 concentrations (∼40–50 µM), suggesting perhaps a nonfractionating NH+4 removal process, such as high temperature NH+4 ion substitution into secondary minerals during hydrothermal circulation [136, 137] or abiotic NH+4 oxidation to N2 . In contrast, high temperature fluids of the Endeavour and Cobb Segments on the Juan de Fuca Ridge exhibited markedly higher NH+4 concentrations than background seawater and lower δ15 N values, likely indicating contributions from an additional subsurface source of buried sedimentary organic matter [13, 58].

8.10 Nitrification Nitrification (󳶳 Fig. 8.11) describes the autotrophic two-step process in which NH+4 is oxidized first to NO−2 and then to NO−3 (note: we use ‘NH+4 ’ to refer to the environmental pool of reduced N, while the autotrophic oxidation process is believed to act on the NH3 phase). During these reactions both NH+4 and NO−2 are used as sources of chemical energy in support of CO2 fixation. Here we will limit our discussion to the scope of biological nitrification that is conventionally regarded to be carried out in two steps by two separate groups of microbes, namely (1) ammonia oxidizers, including archaea (AOA) and bacteria (AOB), and (2) NO−2 oxidizers, which to date have been limited to NO−2 oxidizing bacteria (NOB). Although complete nitrification (NH+4 oxidation to NO−3 or ‘comammox’) by a single microorganism was recently reported for a culture of Nitrospira [138, 139], to date nothing is known about the potential role of this combined pathway under hydrothermal conditions. As can be seen below, the study of this twostep process has almost entirely focused on the first step of ammonia oxidation, leaving much to be learned about the composition and activity of the organisms catalyzing NO−2 oxidation (Concluding Remarks in Section 8.11). As discussed above, fluids emanating from hydrothermal vents are generally anoxic, reducing and warm. Typically these fluids also contain elevated concentrations of NH+4 in comparison to the surrounding cold, oxic seawater, which usually contains NH+4 concentrations near detection level (< 20 nM). As with the chemolithoautotrophic oxidation of other reduced species found in vent fluids (e.g. H2 S, Fe(II),

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Nitrification (NH4+ and NO2- oxidation) Periplasm H2O + NH2OH

NO2- + 5H+ +4 e2

Hao

H2O + NH2OH Cytoplasm

AmoA

Nxr

+2 e1

3 NH3 + O2 + 2H+

NO2+ H2O

ATPase

ADP + Pi NO3+ 2H+

H+

ATP

Fig. 8.11: Enzymes involved in nitrification. The first step in microbial nitrification (both bacterial and archaeal) is the oxidation of ammonium to hydroxylamine (1) catalyzed by ammonia monooxygenase (AMO encoded by amoA). In ammonia-oxidizing bacteria, the second step involves oxidation of hydroxylamine to produce nitrite (2), catalyzed by a hydroxylamine oxidoreductase (Hao). In ammonia oxidation by archaea, this nitrite production step appears to be biochemically distinct from the bacterial pathway and is not yet well understood. Nitrite oxidation to nitrate (3), conducted by nitrite-oxidizing bacteria including the newly discovered comammox bacteria, is facilitated by a nitrite oxidoreductase (Nxr)

CH4 , H2 ), the presence of a hydrothermal supply of NH+4 offers a unique supply of energy for nitrifying organisms. As such, nitrification has been documented in the context of NH+4 supply from hydrothermal vent fluids over a range of studies using a variety of methods including endmember mixing, stable isotopes, rate measurement incubations and quantification of genes and transcripts of known nitrifying organisms. Although very few studies have quantified the dynamics of low concentration variations in NH+4 concentrations of diffuse fluids, mixing diagrams between vent fluids containing NH+4 and bottom seawater containing virtually no NH+4 have indicated nonconservative behavior of NH+4 [13, 30]. Where NH+4 concentrations fall below mixing lines, the consumption of NH+4 is suggested and may be attributed to nitrification or possibly assimilative consumption (󳶳 Fig. 8.2). Indeed, NH+4 dynamics under low temperature diffuse conditions often exhibit variable indications of production or consumption, and therefore unequivocal detection of NH+4 oxidation by mixing analysis is not always straightforward [13, 30]. In addition, variations in subsurface hydrological flow paths, which may variably interact with buried organic matter (󳶳 Fig. 8.1), may also influence endmember (and diffuse flow) concentrations [58]. In general, however, in low temperature diffuse fluids the observed correlations between NH+4 and NO−3 sug-

206 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents gest a strong relationship between transformation processes between NH+4 and NO−3 including nitrification [30]. Additional direct lines of evidence for nitrification have also been made through ex situ incubation measurements. In general these data are limited to studies of NH+4 rich hydrothermal plumes such as those found along parts of the Juan de Fuca Ridge as well as within the Guaymas Basin hydrothermal system. Ammonia oxidation rates of up to 91 nM/day were measured using 15 N-labeled incubations of fluids collected from within the neutrally buoyant hydrothermal plumes at the Main Endeavor Field, compared with much lower rates measured in background seawater just above the plumes (up to 4.7 nM/day) [88, 140]. Nonetheless, even these rates measured just outside the plume were remarkably higher than those typically reported from other deep nonvent environments, such as those measured below oxygen minimum zones in the Eastern Tropical North Pacific (≤0.5 nM/day: [141]), the Arabian Sea (< 0.1 nM/day; [142]) and the Sargasso Sea (< 0.02 nM/day; [143]). The high NH+4 oxidation rates within and around this NH+4 rich plume were comparable to rates reported from surface water ([144] and references therein), reflecting the broad importance of hydrothermal NH+4 supply and autotrophic production supported by ammonia oxidation in this region of the deep ocean. Natural abundance isotopic data has also been used to infer the influence of nitrification supported by hydrothermally derived NH3 . Specifically, the measurement of N isotopes in NH+4 (the reactant) as well as the combined isotopes of N and O in NO−3 (the product of nitrification) in diffuse fluids has been used to suggest an important role for nitrification. The strong N isotope effect during ammonia oxidation leads to pronounced increases in δ15 N values of the residual NH+4 pool. Reports of elevated δ15 N values for NH+4 in diffuse fluids from both the Juan de Fuca Ridge [13] and at Loihi Seamount [30] suggest the influence of a strongly discriminating processes such as nitrification. From the analysis of NO−3 isotopes, which offer a ‘two-dimensional’ perspective from both δ15 N and δ18 O, the influence of nitrification is apparent in the deviation of the NO−3 dual isotopic composition from a trajectory of 1 : 1 (e.g. changes in δ18 O vs. changes in δ15 N) during its consumption. If the NO−3 dual isotopic composition only reflects reductive consumption (whether by assimilation or denitrification), changes to the δ15 N of the NO−3 pool are expected to evolve in parallel with its δ18 O values [49]. Where the dual isotope evolution of the NO−3 pool diverges from a 1 : 1 trajectory (falling above a line originating in the composition of the deep sea NO−3 : δ15 N ∼5.5‰ and δ18 O ∼1.8‰), the simultaneous influence of NO−3 production can be inferred [13, 30, 145, 146]. Indeed, the information contained in the combined concentrations and δ15 N and δ18 O of NH+4 and NO−3 can be used together to leverage a more quantitative assessment of the relative rates of a multiple processes. Using a simple box model, this multicompound, multi-isotope approach was used to demonstrate the combined influence of NO−3 reduction, NH+4 oxidation and NO−2 oxidation in three diffuse fluids of the Juan de Fuca Ridge (Axial Volcano, the Cobb Segment and

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the Endeavour Segment; [13]). The results suggest tight metabolic connections within microbial N cycling communities hosted in these environments. The presence of nitrification has also been demonstrated by a variety of studies targeting specific genetic and molecular markers specific to nitrifying organisms. Environmental genetic studies of nitrification have involved both phylogenetic identification – matching environmental 16S rRNA sequences to those of known nitrifying microorganisms – as well as the detection and quantification of functional genes, mostly focused on the gene encoding the alpha subunit of the ammonia monooxygenase enzyme (amoA), responsible for the first transformation step during the production of NO−2 . Initial studies aiming to quantify the abundance of nitrifying bacteria in the NH+4 rich plume of the Main Endeavor Field on the Juan de Fuca Ridge used fluorescence in situ hybridization (FISH) with 16S rRNA primers targeting sequences of known ammonia-oxidizing bacteria (AOB) from the Betaproteobacteria (e.g. Nitrosomonas sp.). This method suggested that up to 10.8% of the total bacterial population was comprised of AOB [88, 140]. Contemporaneously, the discovery of ammonia-oxidizing archaea (AOA; falling within the Thaumarchaeota) and their ubiquity in marine environments [147–150] raised questions about the relative role of AOB and AOA in catalyzing NH+4 oxidation across a range of ecosystems, including deep sea hydrothermal vents. Nevertheless, these first plume studies concluded that the total archaea population was substantially lower than that of AOB and that nitrification was more likely to be catalyzed by AOB than by AOA the in the neutrally buoyant NH+4 enriched plume rising above the Main Endeavor Field [88]. This dominance of AOB over AOA observed in the NH+4 rich vent plume, however, does not appear to translate to other types of environments in hydrothermally influenced deep sea ecosystems. At the Lau Integrated Study Site in the Lau Basin of the western Pacific, an investigation of ammonia-oxidizing organisms in low temperature Fe-Si rich hydrothermal precipitates determined that archaeal amoA gene abundance was at least two orders of magnitude higher than that of Betaproteobacteria AOB. This suggests that AOA comprises the majority of the ammonia-oxidizing community and may constitute as much as half of the microbial biomass found in these mineral precipitates [132]. More recent studies, leveraging high-throughput sequencing approaches for interrogating the genetic entirety of the microbiome through metagenomics and metatranscriptomics, are beginning to shed more light on the extent of processes occurring in hydrothermal vent systems. A study using GeoChip (a microarray based, high throughput technique) compared a range of functional gene abundances (N, S, and C cycling) in the contexts of both mature hydrothermal vent chimney rock and new (∼15 day old) mineral deposits from sulfide chimneys at the Mothra Field along the Juan de Fuca Ridge [79]. Functional genes related to both bacterial and archaeal nitrification (amoA) were detected in both environments. In the mature chimney, amoA

208 | 8 Microbial nitrogen cycling processes at submarine hydrothermal vents genes accounted for ∼3% of all quantified functional genes, and while most of the amoA sequences belonged to the bacterial domain (i.e. Nitrosomonas and uncultivated Betaproteobacteria), archaeal amoA genes were also notably present. In samples from the newly formed sulfide chimney, amoA genes accounted for ∼1.5% of the total number of functional genes, indicating rapid colonization and utilization of NH+4 by the resident nitrifier population. In addition, metagenomic and metatranscriptomic studies conducted in the Guaymas Basin hydrothermal plumes indicated that ammonia oxidation is one of the main metabolic pathways that support the in situ microbial productivity. This was evidenced by the abundance of NH+4 transporter and amo transcripts as well as the proliferation of AOA within the archaeal community [151–153]. Moreover, some of the most abundant transcripts in the Guaymas plume community are those related to NO−2 oxidation (encoded by nxr genes) and associated energy metabolism. These nxr transcripts are believed to belong to Nitrospirae, a group of NOB whose presence and activity may have eluded detection by previous studies because of the much lower abundance of their genes compared to those of AOA [154]. Microbial communities in the diffuse vent fluids (24 °C) at the Marker 113 vent at Axial Seamount were also investigated using metagenomic and metatranscriptomic approaches [126]. Here, NH+4 concentrations in the diffuse fluid taken 2 cm above the seafloor were 3.5 µM in comparison to < 0.6 µM in the surrounding bottom seawater. Interestingly, the target gene for NH+4 oxidation, amoA, was identified in the metagenome but not in the metatranscriptome, suggesting the metabolic potential for nitrification but perhaps limited levels of its activity during the time of sampling [126]. Additionally, most of the identified amoA genes were identified as belonging to the Thaumarchaeota (the archaeal phylum containing AOA). Finally, the presence of specific intact polar lipids (IPLs), the basic structural lipids making up the lipid bilayer of microbial cell membranes, may also help reveal the presence and proliferation of some nitrifying organisms, as IPLs are thought to derive from recently living organisms and are not thought to persist long after cell death [155, 156]. Crenarchaeol, a unique class of IPL exclusively synthesized by the Thaumarchaeota [157–159], was found in samples extracted from Fe-Si rich mineral precipitates from the Lau Integrated Study Site (Lau ISS), again suggesting the presence and recent activity of AOA in this vent system [132]. The apparently higher relative abundance of AOA at a number of sites challenges earlier studies claiming that AOB are mainly responsible for nitrification. Indeed it appears that the niche partitioning between populations of AOB and AOA in the context of deep sea hydrothermal vent ecosystems may be complex and/or spatially variable. Nonfree-living nitrifiers may also play an important role in N cycling in some vent systems, as suggested by a study of the activity and diversity of nitrifiers living in close association with NO−3 respiring, sulfide-oxidizing Beggiatoa mats growing on hydrothermal sediments in the Guaymas Basin [160]. Guaymas Basin hydrothermal fluids are known to circulate through an exceptionally thick overburden of organic rich sediment, giving rise to fluids hosting extremely high concentrations of organic

8.11 Concluding remarks and future directions

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substrates derived from the thermal degradation of buried organic matter [15, 19]. Near the sediment-seawater interface where these fluids emerge, large microbial mats have been widely reported [103, 161, 162]. In these biogeochemically active microbial mat communities, exceptionally high nitrification rates (74–605 µmol N l−1 mat−1 day−1 ) were reported, greatly exceeding those reported from within a hydrothermal vent plume [88, 140] and in deep ocean margin sediments in general [163]. These elevated rates reflect the generally amplified nature of biogeochemical activity of these organic rich sediments, which support transformation rates more similar to those reported from organic rich environments such as coastal and estuarine sediments (e.g. [164– 166]). Thus far, only autotrophic nitrification has been discussed, wherein energy derived from the oxidation of NH+4 or NO−2 is used to fix inorganic carbon for biomass synthesis. However, heterotrophic nitrification has also been reported from deep sea hydrothermal vent ecosystems. Early studies from 13°N along the East Pacific Rise and the Guaymas Basin and from Snakepit (3,500 m) along the Mid-Atlantic Ridge [167, 168] detailed the physiology of microbial isolates capable of producing NH+4 (via ‘ammonification’ or ‘remineralization’) and in turn oxidizing this pool to NO−2 (but not NO−3 ). These heterotrophic nitrifying organisms were isolated from various hydrothermal vent habitats including fluids, sediments, chimney rocks and invertebrate tissues [167]. As these early studies were based on isolation, it is impossible to draw a conclusion as to how abundant these heterotrophic nitrifiers are in situ or how much they contribute to overall nitrification activity. Closer examination of the role of these organisms in vent ecosystems is clearly warranted and could have important consequences for understanding nitrogen flow.

8.11 Concluding remarks and future directions As utilization and transformation of nitrogen is central to sustaining all known forms of life, a detailed understanding of the chemical, physical and geological factors controlling its availability provides the necessary framework within which we can examine how life evolves, adapts and thrives in hydrothermal vent systems, perhaps even shedding light on factors that contributed to the origins of life on Earth. Every major biological N cycling process that has been studied in the surface ocean and terrestrial systems has also been shown to occur somewhere in the context of hydrothermal vent systems. Yet the factors controlling the relative proportions of transformation pathways and fluxes, while related, can be vastly different from those that are more completely understood in other systems. Furthermore, despite the apparent similarities in types of processes occurring, there remains a multitude of unexplored questions about the nature of nitrogen fluxes within these geologically important systems. At the planetary scale, the role that the global circulation of seawater through the ocean crust plays in nitrogen budgets remains poorly constrained. Indeed, the effects of crustal

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seawater circulation may represent important sinks for oxidized N and/or sources of reduced forms of N (whether locally, regionally or globally), with possible influences on N isotope budgets (albeit likely small relative to processes such as N loss in oxygen minimum zones and N2 fixation at ocean basin scales). Future studies will undoubtedly continue to examine the importance of various N cycling processes in an effort to better understand the controls on their occurrence, magnitude and relative impacts on N budgets at ecosystem scales. Importantly, these efforts will draw upon a diversity of analytical approaches, combining geochemical measurements of rates and isotopic composition with physical measurements of mass flux as well as with increasingly refined examination of microbial community composition and activity. Important progress will be made through the burgeoning use of high-throughput sequencing approaches (omics) permitting an increasingly more sophisticated understanding of the complex interaction of microbial groups and the regulation of their metabolic potential. For example, important questions remain concerning the factors governing the distribution and activity of the chemolithoautotrophic ammonia-oxidizing bacteria and archaea. Finally, intriguing questions arise at the ecosystem scale concerning the environmental regulation of N loss (by anammox and denitrification) versus N recycling (by assimilation and DNRA), including evidence that these processes coexist in a heterogenous fabric of inhibition and stimulation regulated by dynamic temperatures, pH, mineralogy and fluid chemistry. While the range of N cycling processes examined at hydrothermal vents thus far appears to have identified virtually every major player, the N cycle is notoriously complex and continues to offer new surprises. With respect to the N cycling at hydrothermal vents, several poorly examined facets remain that may provide important insights into life in these systems. For example, given the wide range of temperature, pH, reactive mineral surfaces and dissolved catalytic elements encountered, there is a multitude of abiotic reactions that may occur involving both oxidative and reductive fates for N species. Considering, for example, the high levels of dissolved and mineral-associated forms of iron in many of these vent systems, the possible roles of processes such as chemodenitrification (reduction of NO−3 and/or NO−2 by Fe(II)) or feammox (oxidation of NH+4 and/or NO−2 by Fe(III)) [93], which have been observed under soil and sediment incubations (e.g. [169]), may act to regulate the availability of N species in some systems. In addition, despite a number of reported isolates performing heterotrophic nitrification [167], its importance remains unexamined, especially in the context of vent systems. Furthermore, the process of nitrite oxidation (producing NO−3 ), generally thought to be catalyzed by nitrite-oxidizing bacteria, remains virtually unexamined across most vent studies, and yet provides an important redox link between NO−2 and NO−3 . In particular, a number of emerging studies are beginning to suggest the widespread occurrence of rapid redox cycling between NO−2 and NO−3 , perhaps even involving anoxic reoxidation processes [145, 170, 171] or even organisms catalyzing the complete nitrification pathway (comammox) [138, 139]. Additionally, the possible influence of alternative nitrogenases has yet to be closely examined in the context of hy-

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drothermal vents and could have especially important implications for N isotope mass balance calculations. It is also becoming increasingly clear that multicellular agents of N transformations may be more prevalent in many submarine hydrothermal ecosystems than previously believed [119, 172, 173]. For example, while fungi have been identified at a variety of hydrothermal vent sites and habitats, their ecological role in these systems is unclear [174, 175]. Furthermore, it remains undetermined whether they are capable of catalyzing denitrification as has been demonstrated in a large proportion of terrestrial fungi [176]. Closer examination of the metabolic potential of these multicellular organisms may provide important insights into understudied facets of nitrogen turnover in these ecosystems. Finally, almost nothing is known about the sources and fate of dissolved organic nitrogen (DON) in the context of hydrothermal vent systems. There has been an active debate about the role that hydrothermal circulation plays in modifying the DOC pool in the deep ocean [177, 178]. To our knowledge, no work has been done on DON in vent settings, yet much could also be learned by closer examination of its distribution and N isotopic composition. In the context of high temperature focused flows, the abundance and composition of any DON should shed light on high temperature reactions occurring in deeply seated reaction zones. Additionally, the nature of DON in low temperature diffuse fluids may be useful for determining the relative importance of subsurface processes and N sources, such as N2 fixation or NH+4 from remineralization of subsurface organic matter.

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multi-sampler with in situ preservation for microbial and biogeochemical studies. Deep Sea Research Part I: Oceanographic Research Papers 2014, 94:195–206. Taylor CD, Doherty KW, Molyneaux SJ, Morrison III AT, Billings JD, Engstrom IB, Pfitsch DW, Honjo S. Autonomos Microbial Sampler (AMS), a device for the uncontaminated collection of multiple microbial samples from submarine vents and other aquatic environments. Deep Sea Research Part I: Oceanographic Research Papers 2006, 53:894–916. Lavik G, Stührmann T, Brüchert V, Van der Plas A, Mohrholz V, Lam P, Mußmann M, Fuchs BM, Amann R, Lass U. Detoxification of sulphidic African shelf waters by blooming chemolithotrophs. Nature 2009, 457:581–584. Bowles MW, Joye SB. High rates of denitrification and nitrate removal in cold seep sediments. The ISME Journal 2011, 5:565–567. Bowles MW, Nigro LM, Teske AP, Joye SB. Denitrification and environmental factors influencing nitrate removal in Guaymas Basin hydrothermally altered sediments. Front Microbiol 2012, 3:377. Joye SB. Denitrification in the marine environment. In: Collins G (ed). Encyclopedia of Environmental Microbiology. New York, John Wiley and Sons, 2002, 1010–1019. Joye SB, Hollibaugh JT. Influence of sulfide inhibition of nitrification on nitrogen regeneration in sediments. Science 1995, 270:623–625. Marino R, Howarth RW, Chan F, Likens GE. Sulfide inhibition of molybdenum-dependent nitrogen fixation by planktonic cyanobacteria under seawater conditions: a non-reversible effect. Aquatic Biodiversity 2003, 171:277–293. Sorensen J, Tiedje JM, Firestone RB. Inhibition by sulfide of nitric and nitrous oxide reduction by denitrifying Pseudomonas fluorescens. Applied and Environmental Microbiology 1980, 39:105–108. Strous M, Kuenen JG, Jetten MSM. Key physiology of anaerobic ammonium oxidation. Applied and Environmental Microbiology 1999, 65:3248–3250 . Sievert SM, Scott KM, Klotz MG, Chain PS, Hauser LJ, Hemp J, Hügler M, Land M, Lapidus A, Larimer FW. Genome of the epsilonproteobacterial chemolithoautotroph Sulfurimonas denitrificans. Applied and environmental microbiology 2008, 74:1145–1156. Götz G, Banta A, Beveridge T, Rushdi A, Simoneit B, Reysenbach A-L. Persephonella marina gen. nov., sp. nov., and Persephonella guaymasensis sp. nov., two novel, thermophilic, hydrogen-oxidizng microaerophiles from deep-sea hydrothermal vents. International Journal of Systematic and Evolutionary Microbiology 2002, 52:1349–1359. L’Haridon S, Reysenbach A-L, Tindall B, Schonheit P, Banta A, Johnsen U, Schumann P, Gambacorta A, Stackebrandt E, Jeanthon C. Desulfurobacterium atlanticum sp. nov., Desulfurobacterium pacificum sp. no. and Thermovibrio guaymasensis sp. nov., three thermophilic members of the Desulfurobacteriaceae fam. nov., a deep branching lineage within the Bacteria. International Journal of Systematic and Evolutionary Microbiology 2006, 56:2843–2852. Pérez-Rodríguez I, Ricci J, Voordeckers JW, Starovoytov V, Vetriani C. Nautilia nitratireducens sp. nov., a thermophilic, anaerobic, chemosynthetic, nitrate-ammonifying bacterium isolated from a deep-sea hydrothermal vent. International Journal of Systematic and Evolutionary Microbiology 2010, 60:1182–1186. Vetriani C, Speck MD, Ellor SV, Lutz RA, Starovoytov V. Thermovibrio ammonificans sp. nov., a thermophilic, chemolithotrophic, nitrate-ammonifying bacterium from deep-sea hydrothermal vents. International Journal of Systematic and Evolutionary Microbiology 2004, 54:175– 181. McHatton SC, Barry JP, Jannasch HW, Nelson DC. High nitrate concentrations in vacuolate, autotrophic marine Beggiatoa spp. Applied and Environmental Microbiology 1996, 62:954– 958.

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Jeffrey Marlow and Roland Hatzenpichler

9 Assessing metabolic activity at methane seeps: a testing ground for slow growing environmental systems 9.1 Introduction Microbial communities mediate the planet’s foundational biogeochemical cycles, mobilize nutrients, support higher trophic levels, and perform essential ecosystem services. High-throughput genetic sequencing and advanced computational tools have revolutionized our understanding of biodiversity and habitability over the last decade, providing a window into microbial systems. 16S rRNA gene surveys have uncovered rich stories of microbial diversity across a wide range of habitats, from coastal lagoons [1] to intercontinental dust [2] and seafloor hydrothermal systems [3]. Single-cell genomics, as well as computational binning of metagenomic databases, have filled in thousands of genomes and pointed to fermenting bacterial phyla [4] and the hydrocarbon metabolizing potential of Bathyarchaeota [5] and Verstraetearchaeota [6]. Taken together, the genomics revolution has reconfigured the tree of life [7, 8] and described compelling new habitats, from the human gut [9] to the deep subsurface [10]. Despite these exciting developments, an essential aspect of the role of microbes in ecosystem dynamics is missing from the genomics pipeline: metabolic activity. DNA based studies demonstrate genetic potential, but offer limited insight into cell viability or realized biochemical processes; indeed, molecular products derived from many genes are not present at any given time [11, 12], and relic genetic material can persist in soil for years, complicating the interpretation of broad genetic surveys [13, 14]. However, metabolic activity – the ability of organisms to conserve energy and transform their surroundings – is precisely what makes microorganisms so powerful: focusing on this parameter should be a central priority for environmental microbiologists seeking to understand our planet. The challenge of measuring metabolic activity is particularly pronounced at marine methane seeps, where uncultured, slow growing, intricately interconnected microbial communities carry out globally relevant biogeochemical processes [15, 16]. The anaerobic oxidation of methane (AOM), catalyzed by primary producer consortia of anaerobic methanotrophic (ANME) archaea and sulfate-reducing bacteria (SRB), produces little energy [17, 18], yet these organisms support oases of secondary microbial mats, meiofauna, and clam beds [19, 20]. Geologic evidence of seep supported habitats extends back hundreds of millions of years [21], and methane associated pathways are believed to be among the most ancient of microbial metabolisms [22].

https://doi.org/10.1515/9783110493672-009

224 | 9 Assessing metabolic activity at methane seeps

Although anoxic methane oxidizing enrichments have been nurtured [23], efforts to isolate ANME have been unsuccessful to date, as precise metabolic and environmental needs have been difficult to disentangle from consortia and community based relationships. The dearth of other seep hosted microbial isolates is likely primarily due to a lack of attempted approaches, though downflow hanging sponge reactors have proven successful in cultivating methane linked microbes from similar habitats [24]. As a result of these challenges, assessments of metabolic activity at methane seeps have embraced and advanced culture-independent tools. This outlook began as a necessity, but has increasingly been viewed as a strength. After all, integrated microbial communities are the default mode of life [25, 26], and the vast majority of microorganisms remain uncultured [8, 27]. The challenges associated with linking identity to function within this community centric framework have led to an array of creative solutions, which we describe below. In this context, methane seeps have served as an important testing ground for activity based tools and techniques that have since seen application at an array of terrestrial, marine, and subsurface sites. Our assessment of activity oriented research at methane seeps encompasses multiple scales of investigation, from a broad view of community-wide changes to metabolism specific analyses seen through catabolic and anabolic lenses. The research community’s focus on the distinguishing metabolisms of methane seeps – methane oxidation coupled to sulfate reduction – is reflected in the methods highlighted herein, though the opportunity for a wider scope of study is briefly discussed. We begin with system-wide parameters that inform metabolic activity through observation, such as bulk geochemical transformations and microbial colonization (Sect. 9.2). Catabolic activity studies focus largely on the sulfate-dependent AOM by tracking carbon- (methane, bicarbonate) and sulfur- (sulfate, sulfide) bearing molecules in lab based incubations (Sect. 9.3). Anabolism oriented work tracks biomolecule production associated with seep simulating conditions, frequently examining information-bearing molecules in order to link function with phylogenetic identity (Sect. 9.4). We address pressing empirical challenges (Sect. 9.5) before assessing metabolic activity based tools of the future (Sect. 9.6). The culture-independent techniques that have been customized for investigations at methane seeps have exposed surprising new metabolisms and revealed the importance of seeps in biogeochemical processes. These tools have been adapted for a wide range of environmental microbiological investigations, helping a diverse community of researchers better understand how complex microbial systems impact ecosystems worldwide.

9.2 Observational approaches to quantifying seep hosted activity

|

225

9.2 Observational approaches to quantifying seep hosted activity In order to be most relevant from an observational, environment-wide perspective, activity measurements should be conducted in situ with minimal perturbation of the system. This is a challenging proposition in any real-world setting, but the difficulties of working in the deep sea make meaningful in situ work particularly daunting. Operationally, the high pressure, low temperature, time constrained, and robot mediated nature of experimental manipulation at seeps has limited the roster of reliable instrumentation. Hydrologic and microbiological complexity inherent to such habitats – e.g. variable fluid fluxes and flow paths [28], overlapping zones of methanogenesis and methanotrophy [29] – make it difficult to isolate variables. Nonetheless, modeling efforts and colonization studies offer insight into microbial activity at seeps.

9.2.1 Modeling rates of activity from geochemical profiles One of the earliest methods of observation based rate quantification involved the derivation of bulk sediment hosted activity rates from geochemical profiles. In this approach, concentrations of relevant species (potentially including CH4 , SO2− 4 , H2 S, − and HCO3 ) from in situ [30] or recovered core [31] pore waters are measured, and mathematical models are used to determine steady state rates of chemical transformation. In the simplest version of a reaction transport model, rates of methane consumption (R) are determined from one-dimensional concentration (C) profiles and calculated temperature dependent diffusion coefficients (DS,T ), according to equation 9.1 R = DS,T (

δ2 C ) δz2

(9.1)

where C is the dissolved methane concentration, z is depth, and DS,T is between 0.7–1× 10−5 cm2 s−1 for the most commonly applicable temperature range of 2–10 °C [32]. One challenge with such reductionist models is that they are underdetermined with regard to reactant transport: a high flow rate with rapid consumption could produce the same depth profile as low flow and slow uptake. For this reason, the incorporation of other physical parameters like advection rate and porosity, when measurable, are useful; fluid flow in particular signifies the dominant factor in determining the ultimate extent of AOM [33]. Several studies have developed such advection-diffusion-reaction models e.g. [34, 35], based on one-dimensional mass conservation [36, 37]. Treude et al., for example, use equation 9.2 when calculating rates of AOM (R) at Hydrate Ridge, OR [38], where ϕ is porosity, t is time, Θ is tortuosity, u is the advective flow rate, and other

226 | 9 Assessing metabolic activity at methane seeps

variables are represented as in equation 9.1.

ϕ

δC δ DS,T δC δC = (ϕ 2 )− (Θu) − ΘR δt δz δz Θ δz

(9.2)

Other factors, including biomass growth [33, 39] or kinetic [40] and thermodynamic limitations [41], can also be incorporated for higher fidelity treatment of well characterized sites (see Regnier et al., 2011, for a review on seep associated models [42]). Advection-diffusion-reaction models are useful in their quantification of net methane consumption, which is frequently the desired metric on an ecosystem-wide scale. However, reliable measurements of geochemical concentrations and advection rates (which can change substantially with the tidal cycle [43]) remain a challenge, and the elucidation of lineage-specific activity remains unaddressed.

9.2.2 Colonization Rates An alternative approach to measuring in situ activity entails multitimepoint observations derived from samples recovered during distinct sampling campaigns. In one study, a 14-month incubation of sterile glass slides in Eel River Basin seep sediment demonstrated growth of bacteria whose δ13 C signatures indicated minimal incorporation of isotopically light, methane-derived carbon; ANME representatives were not observed [44]. A more detailed study characterized microbial colonization and diversity on sterile carbonate and wood substrates after 13 months at actively seeping and seemingly inactive sites [45]; similar experiments at hydrothermal vents have demonstrated Epsilonproteobacterial colonization of titanium rings after 20 days [46]. This time resolution circumvents high frequency temporal bias in seepage activity – which can vary over days [28] or weeks [47] – and begins to constrain the timescales over which communities can spread across newly amenable substrate. Colonization by Thiotrichaceae and Helicobacteraceae representatives at active seeps suggests a prevalence of sulfur-oxidizing niches, while ANME archaea point to anoxic, methane rich endolithic habitats and seeding of microbial constituents from below. The observation that newly colonized rocks and ‘native’ carbonates contained distinct microbial assemblages suggests that 13 months is insufficient time for a mature, steady-state community to develop; more extended time series studies could further illuminate this process of successional activity. Repeated in situ sampling is useful for assessing activity and rate based features under natural conditions to measure parameters of interest, potentially including biomass growth or authigenic carbonate precipitation. As access to seep sites becomes more reliable – e.g. through cabled seafloor observatories – these modes of analysis are likely to become more ingrained in seep based science.

9.3 Catabolism based methods |

227

9.3 Catabolism based methods Transporting sediment and carbonate rocks from seep impacted environments to the lab enables a much greater range of activity based experiments, allowing researchers to track specific aspects of catabolic or anabolic metabolism. These experiments are typically performed in micro- or mesocosm incubations (e.g. tens to hundreds of milliliters of substrate), sacrificing features of environmental realism (advective transport of reactants and wastes, temporal variation) but providing more consistent control of chemical concentrations and sampling. The most common mode of activity measurement quantifies reactants and products of catabolism, the metabolic reactions that conserve energy in support of other cellular processes. These studies typically focus on methane and sulfate dynamics, so carbon, hydrogen, and sulfur are the best-established atomic tracking systems.

9.3.1 Tracking methane catabolism Methane metabolism in anoxic sediments involves a multifaceted set of reactions. Most broadly, ANME archaea partner with SRB to enact sulfate dependent AOM [18] (Reaction 1), while methanogens (comprising seven orders of Euryarchaeota and possibly members of the newly characterized Bathyarchaeota [5] and Verstraetearchaeota [6]) produce methane from a number of precursors, most prominently hydrogen and carbon dioxide (Reaction 2) or acetate (Reaction 3). − − CH4 + SO2− 4 → HCO3 + HS + H2 O

(rxn. 1)

4H2 + CO2 → CH4 + 2H2 O

(rxn. 2)

CH3 COOH → CH4 + CO2

(rxn. 3)

When quantifying rates of AOM, several reactants and products can be tracked. These procedures capture distinct aspects of a methane-processing pathway, and terminology is frequently used loosely, obscuring the true meaning of a study. We seek to clarify the situation by distinguishing between four commonly measured aspects of AOM as well as methanogenesis, which frequently occurs in colocated sediment horizons and whose activity can complicate interpretations (󳶳 Fig. 9.1). 9.3.1.1 Methane activation Methane activation signifies the cleavage of at least one C–H bond but does not rule out the back-reaction to reform methane [48]. A back-reaction at any point in the reverse methanogenesis pathway would maintain the initial methane derived carbon atom, making isotopic labels attached to carbon ineffective in quantifying methane activation. Hydrogen isotopes, however, can provide insight, and techniques track-

228 | 9 Assessing metabolic activity at methane seeps

(Net) Methane Consumption Methanotrophy

Biomass

Autotrophy

Catabolic Intermediates

CH4

HCO3-

(Complete) Methane Oxidation

Mcr

HCO3-

Methane Activation

Methanogenesis

CH4

e(Complete) Sulfate Reduction

HSMcr

Biomass

SO42-

Catabolic Intermediates

Biomass

HCO3Autotrophy

Autotrophy Sulfate Consumption

Catabolic Intermediates

H2

HCO3-

Fig. 9.1: A schematic diagram showing a network of catabolic and selected anabolic transformations amenable to activity oriented inquiry. Arrow thicknesses are intended to give a relative sense of metabolite fluxes under most AOM experimental conditions but are not to scale. See text for additional details

ing deuterium [49] and tritium [50] from methane into the aqueous phase have been developed. The radiotracer 3 H−CH4 was first synthesized in order to quantify aerobic methane oxidation rates in the Eel River Basin water column [50]. The developers viewed 3 H−CH as an improvement over 14 CH radiotracers because of its 2,000-fold higher 4 4 specific activity, meaning that added concentrations could be kept as low as 3 nM and artifacts of experimental conditions could be avoided. Subsequent experiments have measured aerobic methane activation rates of North Sea water [51], Arctic fjords [52], and the eastern North Pacific [53]. Discrepancies between 14 C and 3 H based rates were attributed to different concentrations of added methane (< 2 nM 3 H−CH4 , 450 nM 14 CH ) [52] or priming effects and method backgrounds [53]. 3 H−CH studies have not 4 4 fully engaged with the nuances involved with measuring hydrogen- versus carbonlinked atoms, typically equating 3 H presence with H14 CO3 − production [50–53], an approach that neglects biomass bound 3 H and 14 C, as well as back-reactions that would increase aqueous 3 H concentrations without full methane oxidation. To date, the 3 H−CH4 method has only been used in aerobic methane activation studies. An alternative method using monodeuterated methane (CH3 D) has been tested in both aerobic and anaerobic methane activating systems; following an incubation

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period, D/H ratios of filtered water samples are measured on a stable isotope water analyzer [49]. When directly compared with the 14 C radiotracer method in AOM incubations, calculated rates of tracer appearance in the aqueous phase were consistently offset by a factor of two, demonstrating both the applicability of the method and the relevance of anabolic pools. To constrain the prevalence of the methane-forming back-reaction, the changing proportion of CH3 D and CH4 was measured with nuclear magnetic resonance (NMR) spectroscopy, revealing that 4–16% of aqueous D was attributable to activated, back-reacted methane during AOM [49]. Using hydrogen as an isotopic probe of methane oxidation remains an immature field because the partitioning of methane derived hydrogen into catabolic and anabolic products is poorly understood. For this reason, a more thorough accounting of hydrogen atoms liberated through methane activation and oxidation would be useful; a combined study of D/H water ratios and anabolic incorporation of D (via nanoSIMS [54]) would begin to provide clarification. 9.3.1.2 Methanotrophy Methanotrophy, in the strictest sense, refers to the incorporation of methane derived carbon into cellular constituents. However, because much of their biomass carbon is sourced from dissolved inorganic carbon (DIC) [55], ANME can best be described as mixotrophs that combine autotrophy (DIC-carbon) and methanotrophy (methanecarbon). A more complete assessment of methanotrophic measurement approaches is provided in Sect. 9.4. 9.3.1.3 Complete methane oxidation Chemical oxidation strips electrons from an atom, and while the oxidation of methane’s carbon from CH4 to HCO−3 entails a loss of eight electrons, any stage beyond the initial reaction of methane with the CoM–S–S–CoB heterodisulfide constitutes methane oxidation. However, metabolites of intermediate oxidation states are rarely measured, and practitioners typically report DIC concentration as a proxy for methane oxidation (Reaction 1). We propose ‘complete methane oxidation’ as a more precise term associated with such measurements. Tracking changes in DIC as a metric of methane oxidation rate – even when done while measuring methane concentration – is insufficient in the context of mixed microbial communities. DIC is the end product of many heterotrophic metabolisms, and given the heightened concentrations of higher hydrocarbons in seep hosted sediments [56, 57], the origin of fully oxidized carbon is difficult to determine unless isotopic labels are used judiciously. Oxidized carbon can also enter autotrophic pathways and be recycled into reduced compounds. Tracking radiocarbon (14 C) from methane (14 CH4 ) into DIC is perhaps the best-established method for quantifying rates of complete methane oxidation. The approach was first codified by Reeburgh [58] based on similar techniques measuring methane

230 | 9 Assessing metabolic activity at methane seeps production from 14 C-acetate or lactate [59]; procedural enhancements from Iversen and Blackburn [60] and Treude et al. [61] have since streamlined the process and minimized potential errors associated with incomplete analyte recovery. After dissolved radioactive methane is injected, the sample is incubated at the desired conditions. Because the detection of 14 C atoms is so sensitive, experiments only require small volume amendments (typically a few tens of microliters) and brief incubation periods (hours to several days). To halt microbial activity, NaOH is added, and several parameters are measured. Total methane (CH4 ) is determined via gas chromatography from a control, time-zero sample. Postincubation residual 14 CH4 is measured by scintillation counting after headspace combustion at 850 °C and capture of the formed 14 CO2 with phenylethylamine. Biologically formed 14 CO2 is quantified by scintillation counting following acidification of the sample (driving any adsorbed or precipitated oxidized carbon into the gas phase) and capture of the resulting 14 CO2 gas in phenylethylamine. When the volume of the sample (v) and duration of the experiment (t) are taken into account, the rate of methane oxidation is calculated as shown in equation 9.3. Methane Oxidation =

14 CO ⋅ CH 2 4 + 14 CO2 ) ⋅

(14 CH4

v⋅t

(9.3)

Since its development, the 14 C method has arguably become the most pervasive approach used in rate quantifications; it has been employed in lab based incubations of marine sediments from the Gulf of Mexico [62], Saanich Inlet [63], Skan Bay [64], Cape Lookout Bight [65], Guaymas Basin [66], and the Chilean margin [61], as well as carbonate rocks from Hydrate Ridge [67]. A new approach allows researchers to introduce even less 14 CH4 reactant [68] by exploiting the 103 −109 -fold more sensitive 14 C measurements of accelerator mass spectrometery [69]. 13 C can also be used to measure rates of complete methane oxidation, an approach that was most substantively developed for the examination of ‘trace methane oxidation’ by cultures of known methanogens [70, 71]. 13 CH4 is added to the sample, and concentrations and isotopic ratios of headspace carbon dioxide and methane are measured by mass spectrometry (see Moran et al., 2007 for additional details [71]). Fractionation factors and back-reactions of labeled DIC to labeled methane should also be incorporated into rate equations depending on the amount of label and timescale of the experiment. In the context of environmental communities, detection limits could render this approach challenging, particularly when ∼ µM addition constitutes a major perturbation, as in the water column above methane seeps [68]. With incubations from more methane replete environments such as anoxic seep sediments, 13 CH4 has been used to probe the relative rates of AOM upon the addition of H2 and methyl sulfides [72]. This approach revealed a decoupling of sulfate reduction and complete methane oxidation at low sulfate levels [73], and confirmed methane oxidation coupled to anthraquinone2,6-disulfonate (AQDS) reduction [74]. Quantifying 13 C distributions among oxidized

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and reduced pools via nuclear magnetic resonance spectroscopy has been the method of choice for detailed in vitro biochemical studies on methyl coenzyme M reductase systematics [48, 75]. 9.3.1.4 Net methane consumption Methane consumption involves a decrease in methane concentration but is endproduct agnostic. Studies reporting this parameter frequently measure decreases in aqueous or headspace methane concentration and may or may not account for methyl coenzyme M reductase (Mcr) back-reaction or methanogenesis occurring in the same sample [29]. Thus, when isotopically unconstrained methane concentration is reported, “net methane consumption” is a more precise diagnosis. Several efforts have quantified methane concentration within experimental incubations to determine net methane consumption rates. These include simultaneous tracking of methane loss (via a gas chromatography-flame ionization detector) and sulfide production to demonstrate 1 : 1 stoichiometry at a range of pressures [76], and mass spectrometry based concentration assessments at the inlet and endpoint of a ‘simulated seep’ flow-through reactor [77, 78]. The technique was also used, alongside stable isotope probing, omics studies, and supplementary geochemical measurements, to demonstrate a novel ANME lineage linking AOM to nitrate reduction [79]. Net methane consumption, as measured through changes in headspace or aqueous methane concentrations, can be a relatively straightforward way to capture systemwide microbially mediated processes, but additional details on interacting metabolic pathways or phylogenetic constraints on activity require more nuanced tools. 9.3.1.5 Methanogenesis The anoxic sediments of methane seeps can host conditions that are energetically amenable to both methanogenesis and AOM [38, 80], particularly at sites of high organic load input [65] or serpentinization [81]. Both methanogens [82] and ANME [83] have also been demonstrated running their ‘primary’ reactions in reverse, complicating the interpretation of experiments with a unidirectional pathways focus. To better constrain simultaneous methane production and methane oxidation, catabolism oriented isotopic labels on oxidized carbon species can be incorporated into incubation conditions. To determine the role of both aceticlastic and bicarbonate based methanogenesis in Gulf of Mexico seep sediments, Orcutt et al. added 14 C-bicarbonate and 14 C-acetate to incubation experiments and tracked 14 CH4 production with headspace combustion and scintillation counting. The results were integrated into parallel incubations quantifying complete methane oxidation rates, demonstrating that AOM was ten times as rapid as 14 C-bicarbonate reduction [29, 84]. Such studies highlight the complexity of methane metabolism and indicate how multiple pathways can interact to regulate seep geochemistry.

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9.3.2 Tracking sulfur catabolism Methane oxidation is frequently the primary interest of researchers investigating methane seeps, but given its intimate link with sulfate reduction, measuring sulfur related processes can help clarify interorganism relationships and reveal hidden metabolic linkages. Canonical dissimilatory sulfate reduction is a three-step process converting dissolved sulfate to adenosine 5󸀠 -phosphosulfate (APS), sulfite, and ultimately sulfide; the entire process is an eight-electron reduction, and its pairing with methane oxidation to HCO−3 thereby results in a 1 : 1 stoichiometry. Because of this concordance, sulfur catabolism measurements can be used to both track AOM and place methane into a broader context as a carbon source in anoxic zones. Most straightforwardly, sulfate concentrations can be quantified using ion chromatography [38, 62], high-performance liquid chromatography [85], or reaction with barium chloride [86], while sulfide can be assessed through the Cline assay [87]. These approaches have supported studies of the zonation of AOM driven sulfide mineralization [88] and the effect of temperature or pressure on sulfate-linked AOM rates [76]. Stable and radioisotopes have also been used to probe dissimilatory sulfate reduction, offering the benefit of more specific molecular tracking and higher sensitivity [89]. Na2 35 SO4 is introduced to experimental incubations, and after the desired time period zinc acetate is added to stop microbial activity and fix H2 35 S as Zn35 S. Inorganic sulfides are recovered by chromous acid digestion [90], and sulfide and aqueous sulfate activities are determined by scintillation counting. Sulfate reduction rates (SRR) are then calculated according to equation 9.4 SRR =

35 [SO2− 4 ] ⋅ α SO4 ⋅ (aH2 S)

t ⋅ v ⋅ (a35 SO2− 4 )

(9.4)

where [SO2− 4 ] is the sulfate concentration, α SO4 is the isotope fractionation factor for sulfate reduction, t and v are the incubation time and volume, and a values designate chemical activities of the relevant 35 S-containing pools. Measures of sulfate reduction, when quantified alongside methane oxidation, help clarify the role of methane as a driver of overall heterotrophic activity and can expose intriguing aspects of the AOM-SR processes. Studies of sediments from geologically distinct seeps attribute varying levels of total sulfate reduction to methane oxidation, ranging from 7–11% in the Black Sea [34] to nearly 100% in high flux, organic poor sediments from the Bullseye seep [91]. Combining concentration based rate measurements with isotopic analysis has been particularly fruitful in developing interpretive frameworks for in situ methane and sulfate isotopic distributions. Such studies have shown that methane δ13 C values are linked to sulfate concentrations, δ34 S reflects methane concentrations [92], and sulfate δ18 O is suggestive of sulfide reoxidation [93, 94].

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9.3.3 Tracking catabolism of ‘nontraditional’ electron acceptors Historically, investigations of AOM metabolism have focused on sulfate as the associated electron acceptor [18, 95], a choice that reflects the dominant metabolic coupling in marine (high sulfate) methane seep settings. However, alternative electron acceptors have been implicated in other systems: nitrate in enrichment bioreactors inoculated with wastewater sludge [79, 96] and nitrite in terrestrial sediments [97]. AOM linked to metal reduction, meanwhile, was first demonstrated in seep sediment incubations [98]. Since then, strategic amendments with methane, sulfate, iron, manganese, as well as selective inhibitors like molybdate and bromoethanesulfonate, have pointed to cryptic cycling [99, 100] and suggested a role for metals in seep [94], marine [101, 102], and freshwater [103, 104] AOM. Studies pursuing metal reduction add quantification of metal species to the assessments of methane oxidation and sulfate reduction discussed above. Ferrous and total iron (following reduction with hydroxylamine hydrochloride) can be quantified with the ferrozine assay [105], while reduced manganese can be assessed spectrophotometrically after mixing with formaldoxime [106]. Nonetheless, the roles of biotic and abiotic processes in metal associated AOM remain difficult to interpret; reactive intermediates intercepted from putative metabolic reactions [107] or metal isotope distributions could provide useful insight in the coming years.

9.4 Anabolism based methods Anabolism is the process through which molecules or elements are incorporated into biomass. This side of metabolic activity touches more tangentially on the ecosystemwide, biogeochemical cycling implications of microbial communities; instead, it exposes components such as cell growth and proliferation, and, through the assessment of information-carrying molecules, provides a more granular look at phylogenetic identity and metabolic mechanisms. Here, we briefly review anabolism based studies of methane seep systems, whose uncultured core metabolism (AOM) has prompted the development and enhancement of several analytical tools. This assessment will cover a wide continuum of informational biomolecules, from cell counts to biomoleculespecific stable isotope probing and the hijacking of translational machinery to install chemically useful residues in incipient proteins.

9.4.1 Cell quantification Counting the number of cells in a sample is among the most fundamental measures of the microbiological response to empirical conditions, demonstrating whether the microbial community is expanding or contracting, while also providing a basis for per-

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cell rate calculations. In the context of seep sediments, the efficient, unbiased isolation of cells is an initial obstacle, since biomass can adsorb to charged clay particles or be shielded from in situ lysis reagents [108]. The use of a Percoll density gradient, in concert with thermal treatment and sonication, effectively concentrates microbial biomass from seep sediments [109, 110] and carbonate rocks [67] (see Dawson et al. for a more detailed protocol [111]). Alternatively, chemical treatment followed by sonication and Nycodenz and/or polytungstate density centrifugation resulted in high cell extraction efficiencies in Arctic Ocean subseafloor sediment [112] and Nankai Trough subduction zone sediments [113]. These processes do not appear to disrupt spatially associated consortia [15], but the retention of active cells following the procedures has not yet been demonstrated. Isolated biomass can be used for direct cell counts, lineage-specific characterization through fluorescence in situ hybridization (FISH), and quantitative PCR (qPCR). Cell counts with general DNA dyes (e.g. acridine orange, DAPI) have demonstrated the proliferation of ANME-SRB consortia in Hydrate Ridge seep sediment after two years of lab based incubations. Supplementing the analysis with FISH probes to confirm phylogenetic identity, a consortia doubling time of approximately seven months was calculated [114]. qPCR was used as a proxy for cell counts to demonstrate lineage-specific growth rate enhancements among ANME [78]. Loss of DNA during extraction could limit the interpretive power of qPCR results. While water column studies have noted a concordance between qPCR and other cell quantification methods [115], a study from nonseep associated anoxic marine sediments, which compared results from general DNA stain, FISH, CARD-FISH (CAtalyzed Reporter Deposition-FISH), and qPCR approaches, came to a different conclusion. Cell abundances determined by general stains and the sum of domain-level CARD-FISH numbers were similar, but qPCR derived gene copy numbers demonstrated a highly variable underestimate of overall abundances (i.e. cell counts from general DNA staining) [116]. These types of paired analyses suggest that physical substrates such as marine sediment could play a complicating role in the recovery of DNA and associated assessments of microbial activity and proliferation.

9.4.2 Stable isotope probing Stable isotopes of hydrogen, carbon, oxygen, and sulfur have proven effective in characterizing catabolism in microbial communities from methane seeps (see Sect. 9.3.1– 9.3.2 above). In the context of anabolic processes – many of which involve informational molecules that help link identity with potential or realized function – protocols based on carbon, nitrogen, and, to a lesser extent, hydrogen incorporation are well established. Here, we briefly demonstrate the types of analyses enabled by stable isotope probing (SIP) along the continuum of informational richness of anabolic analyte, from whole cells to lipids, DNA, RNA, and proteins.

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The design of a SIP study depends on the specific questions being addressed as well as several general considerations. First, the atomic proportion of the biomolecule of interest should be evaluated; mean cellular biomass, for example, has an approximate stoichiometry of CH1.66 O0.46 N0.14 [117], so tracking carbon atoms provides a higher potential signal than nitrogen atoms. Second, the source material for anabolic pathways determines the proportion of labeled atoms that are ultimately incorporated into biomolecules of interest. For example, an estimated 50% of ANME lipid carbon is derived from methane while the remainder comes from bicarbonate [55], making both precursors equally promising molecular probes of carbon incorporation. Third, experimentalists must evaluate the proclivity of potential atomic labels to exchange with the aqueous phase during metabolism. In particular, it is difficult to directly trace hydrogen atoms from reactant to biomolecular product, as many hydrogen bonds are relatively labile. Fourth, the times of biomolecule turnover will influence the concentration of the label and the duration of the experiment: DNA generally requires cell division to incorporate a label, while RNA, proteins, and lipids are continually regenerated [118]. Similarly, the duration of the experiment should be tuned to avoid nonspecific or otherwise undifferentiable cross-feeding, a process whereby an isotopic label moves among multiple trophic levels. Finally, technical considerations related to instrumental readout may limit experimental possibilities. The geometrical constraints of nanoSIMS detectors, for example, make simultaneous hydrogen and sulfur detection extremely challenging, and experiments should be designed with such limitations in mind. 9.4.2.1 Intact cells In 2001, investigators began tracking the methane dependent incorporation of 13 C and/or 15 N into individual cells with the newly available technique of secondary ion mass spectrometry (SIMS). To pinpoint the identity of these cells, FISH probes were developed to target 16S rRNA genes with varying degrees of phylogenetic specificity [119]. The combination of these tools – FISH-SIMS – facilitated the discovery of AOM-mediating consortia from methane seep sediments, initially through descriptions of natural isotopic abundances [15, 120], and later through the addition of isotopically enriched substrates [121]. Subsequently, FISH enhancements that bolstered signal (CARD-FISH[122, 123]), enriched for certain lineages (magneto-FISH [124]), or targeted functional genes (mRNA-FISH [125]), as well as more sensitive SIMS instruments (nanoSIMS [126]) have broadened the realm of empirical possibility. NanoSIMS counts ions of several predesignated mass/charge ratios at a resolution of ∼50 nm [127] across planes of material that are ablated away. By compiling these points and generating a 3D map of dozens or hundreds of closely associated cells, patterns of elemental uptake can be measured and the cross feeding relationships between different lineages can be inferred. FISH-nanoSIMS studies using stable isotope substrates have implicated ANME in nitrogen fixation [128], revealed consortia-

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wide differences in nitrate uptake based on distinct SRB lineages [109], proposed novel metabolic associations between consortia constituents [129], and demonstrated metabolic decoupling between ANME and SRB [74]. By extending the position of one of the instrument’s detectors and circumventing the 22 : 1 mass range restriction, simultaneous analysis of H, C, and N isotopes was recently performed [54]. In this mode, biomolecule exchangeability with the aqueous phase becomes an asset, as deuterium uptake reflects baseline activity levels that can permit more targeted questions of anabolic activity [54]. NanoSIMS has also helped clarify dynamics of metabolic activity in other seafloor studies, revealing carbon fixation by worm symbionts in shallow sediments [130] and demonstrating viable autotrophic cells in Pleistocene-aged sediments off the coast of Japan [131]. 9.4.2.2 Lipids Lipid-SIP tracks 13 C into phospholipid fatty acids and is among the more sensitive SIP techniques [132]. Because lipids are the primary components of cell membranes, this approach is sensitive to rapid changes in microbial populations [133]. While the phylogenetic resolution of lipids is lower than other SIP relevant biomolecules, ANME-1, -2, and -3 can be distinguished based on the presence of glycerol dialkyl glycerol tetraethers, crocetane, and particular pentamethylicosenes respectively [134, 135]. Sulfate-reducing partners are marked by varying ratios of C15 and C16 fatty acids [134]. Isotopic analysis of these phylogenetically constrained lipids allows researchers to determine the relative proportions of heterotrophic and autotrophic anabolism when samples are incubated with deuterated water (whose signal is incorporated into all metabolic activity) and 13 C-DIC (which is only incorporated into lipids in the case of autotrophy) [136]. When applied to ANME-1/HotSeep enrichments from Guaymas Basin sediment, this approach indicated that between 8–25% of both bacterial and archaeal lipid carbon comes from methane [137]. An earlier study found that proportion to be approximately 50% in ANME-2 dominated communities from Hydrate Ridge; partnering SRB were almost entirely autotrophic [55]. Subsequent studies have tracked deuterium from water into different ANME-1 lipids to clarify biosynthetic pathways and appreciate the nuances of distinct enrichment patterns among different lipids within the same lineage [138]. Continued study of moleculespecific biosynthetic processes, as informed by stable isotope additions, will help establish which molecules could serve as useful bellwethers of anabolic activity in environmental or microcosm settings. 9.4.2.3 DNA DNA-SIP experiments incubate samples with 13 C-labeled substrate for a period of time, extract genetic material, separate labeled and unlabeled DNA by ultracentrifugation in a CsCl solution, and characterize the enriched fraction [139]. The method is well positioned to link phylogenetic identity to a marker of growth in the form of

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genome replication, which is required in order for labeled substrate to be incorporated into the genetic material [140]. This prerequisite is both a liability (it can be slow and unevenly distributed) and a strength (it ensures that reproduction has occurred, if biomass expansion is a key focus). In the context of methane seep microbial communities, the several-month doubling times of ANME-SRB consortia would require extended experiments in order to recover a representative distribution of labeled DNA. Potentially of greater concern is the fact that all metabolic activity that occurs between cell divisions – the sum of a cell’s biochemical imprint on its surroundings – goes unseen by DNA-SIP. Few researchers have pursued DNA-SIP in the context of marine methane seeps, likely because they have been dissuaded by the long doubling times and the efficacy of FISH-SIMS in identifying organisms responsible for methane oxidation. However, in many locations, methane represents just a fraction of the electron donors that support sulfate reduction [141], so casting a wider net for alkanes that serve as reductants is an important undertaking. For example, a paired DNA- and RNA-SIP study of SRB from Amon mud volcano and Guaymas Basin sediment identified four clades that oxidized and incorporated 13 C from butane and dodecane [142]. In a study of aerobic methanotrophy in surface layer sediment at the Coal Oil Point seep off California, members of the Methylophaga and Methylophilaceae lineages – neither of which had been associated with methane oxidation – assimilated labeled carbon into their genomes after just three days of incubation with oxygen and 13 CH4 [143]. Other hydrocarbon metabolizing studies have targeted heavy DNA fractions from mixed communities incubated with benzene [144] and toluene [145] in an effort to specify which members could encode degradative pathways [146]. Despite its limitations, DNA-SIP study of methane seep communities could clarify the timescales over which various constituents replicate, offering data to complement qPCR or FISH based quantification efforts. When used in concert with other, more metabolism oriented techniques (see Sect. 9.4.2.4–9.4.2.5 below), it could reveal fundamental lineage or metabolism dependent relationships between activity and propagation. 9.4.2.4 RNA Because transcription is a hallmark of viable cells, RNA-SIP is a useful way to measure activity and, by sequencing and annotating isotopically labeled transcripts, the metabolically directed nature of that activity can be discerned [147, 148]. It is also substantially more sensitive than DNA-SIP: an early comparison found that 13 C accumulated in RNA nearly ten times as quickly as in DNA [149], while next generation sequencing platforms can detect transcripts with just 1.5 atom percent 13 C [150], or fully labeled RNA at mixing ratios as low as 0.001% [151]. This discrepancy was illustrated by the detection of ammonia oxidizing archaea from soil inoculum through labeled RNA; due to slow doubling times, no signal was found in the DNA [152]. Over time,

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cross feeding results in diffusion of a label across trophic levels of a community. If sampling points are chosen strategically, primary feeders can be identified first, followed by selective and generalist heterotrophs [153]. RNA-SIP played an important role in the identification of SRB actively engaged in nonmethane hydrocarbon degradation [142]: heavy transcripts were reverse transcribed and terminal restriction fragments were linked with Desulfosarcina lineages that dominated mud volcano sediment amended with ethane, propane, and butane [154]. Fortunato and Huber used biomass from hydrothermal vent fluid to assemble a metatranscriptome enriched in bicarbonate derived 13 C [155]. The approach revealed dynamic community responses to incubation conditions: Epsilonproteobacteria shifted from using oxygen as an electron acceptor to using nitrate at moderate temperatures, and Methanococcus-derived methane metabolism transcripts dominated at higher temperatures. The proven ability to generate a community-level metatranscriptome from seafloor habitats [155] and deep sediments [148] opens the door to similar investigations at methane seeps. 9.4.2.5 Proteins Proteins constitute the biochemical machinery of the cell, making the elucidation of actively produced proteins via SIP-proteomics an insightful window into the metabolic needs of microbes. In an environmental context, detection limits are estimated to be at approximately 10% label incorporation [156], and distinct incorporation values derived from multiple time points provide useful data on assimilation pathways and community dynamics [157]. Labeling can be conferred through substrates or common nutrients (such as 13 C-glucose or 15 NH3 ) or through the addition of labeled amino acids; the former approach typically generates more meaningful results, as biosynthetic pathways can be partially reconstructed, while the latter offers more predictable mass shifts and highly accurate quantitation [158]. Unlike the SIP methods described above, proteomics requires metagenomic knowledge – ideally derived from the same sample or environmental system – against which to search mass spectra results. Initial protein-SIP efforts examined interspecies interactions in toluene degrading cocultures [159], while more recent studies have moved into progressively more complex communities and environmental systems. Using sediment from Hydrate Ridge and Gulf of Mexico seep environments, Krüger et al. showed methane dependent uptake of labeled ammonium into proteins after only three weeks of incubation [160]. Bulk 15 N-enriched protein analysis as a measure of ANME biosynthesis was shown to be more sensitive than incorporation of 13 C from methane into lipids; the dominance of the Mcr protein in the ANME proteome may be a factor [161]. A SIP-metaproteomics study of Hydrate Ridge sediment detected thousands of proteins, including the full ‘reverse methanogenesis’ pathway attributed to ANME lineages. Multiple time points and 15 N enrichment values revealed that the addition of methane prompted ANME and SRB to synthesize proteins and led to decreases in transcription machinery as-

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sociated with other metabolisms linked to oxygen or nitrate reduction [162]. Dozens of Mcr orthologs produced during the course of the experiment also enabled the detection of abundant posttranslational modifications on Mcr, exposing an additional dimension of functional diversity with an as-yet-undetermined bearing on metabolic activity.

9.4.3 BONCAT Bioorthogonal noncanonical amino acid tagging (BONCAT) is a recently developed approach that allows protein synthesis in uncultured cells to be visualized [163– 166], co-opting anabolic processes in order to identify active organisms. BONCAT is based on the in vivo incorporation of a synthetic amino acid that exploits the substrate promiscuity of specific tRNA-amino acyl synthetases. Only two known artificial amino acids are able to hijack the natural translational machinery without recognized biases against specific taxonomies or physiologies or the need for genetic modification of the host cell: L-homopropargylglycine (HPG) and L-azidohomoalanine (AHA), which both replace L-methionine (Met) during protein synthesis [167]. In contrast to AHA, HPG is not prone to chemical transformation under the alkaline, highly sulfidic conditions typically encountered in marine sediments [165]. After incorporation into new proteins, the amino acid can be fluorescently detected via azide-alkyne click chemistry. This allows biosynthetically active cells to be detected, either while still alive or in the chemically fixed state [163–165]. When used in conjunction with rRNA-targeted FISH, BONCAT allows cell identity and protein synthesis activity to be linked (󳶳 Fig. 9.2). BONCAT has been demonstrated to correlate well with other, independent proxies of cell growth, specifically the incorporation of 35 S-methionine or 15 NH4 + often used in microautoradiography [166] or SIMS [165] experiments. The power of BONCAT for seep associated ecophysiology experiments was recently demonstrated in a study that tracked the translational activity of syntrophic consortia catalyzing AOM (󳶳 Fig. 9.2) [163]. By combining BONCAT, fluorescenceactivated cell sorting (FACS), whole genome amplification, and 16S rRNA gene sequencing, hundreds of individual, biosynthetically active archaeal-bacterial partnerships were identified. The study revealed that representatives of all major clades of archaeal methanotrophs occurring in a single methane seep sample were active under controlled incubation conditions. Unexpectedly, ∼14% of AOM consortia were anabolically active even in the absence of methane, either suggesting that energy sources other than methane oxidation could fuel the metabolism of this partnership or that energy storage compounds accumulated during times of plenty were mobilized. BONCAT-FACS also led to the discovery of a previously unrecognized interaction of ANME with members of the Verrucomicrobia, a poorly understood phylum that is widely distributed in marine sediment [168]. The nature of interactions between these two partners is as yet unknown.

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Fig. 9.2: Detection of protein-synthesizing cells in an ANME-SRB consortium. Sediment from the Hydrate Ridge methane seep site was incubated for 114 days in the presence of 50 µM HPG before being analyzed by BONCAT-FISH. Blue – DNA-stain DAPI. Red – an archaea-specific FISH probe. Green – translationally active cells as detected by BONCAT. Lower right shows an overlay of all channels. Scale bar equals 10 µm

BONCAT promises to be a valuable component of the activity based measurements toolbox. The approach is particularly attractive when studying substrates that are not available as isotope-labeled derivatives. In contrast to isotope-labeling studies, which require specialized instrumentation, BONCAT relies on standard epifluorescence microscopy. The approach is simple to establish, comparatively high throughput, and uses inexpensive reagents that are readily available from a number of vendors [164].

9.5 Key outstanding issues and challenges Over the last two decades, researchers have made substantial progress in uncovering the details of methane seep ecosystems through the lens of microbial metabolic activity. New metabolisms, subseafloor habitats, and interspecies interactions have been described, while analytical tools developed in service of seep related investigations have proved useful in other fields. Nonetheless, several obstacles currently obscure our understanding of key aspects of how seeps operate and what they mean for global systems. In particular, while work to date has focused largely on lab based microcosm incubations and the integrated microbial communities therein, future researchers face challenges in environmental realism and varying temporal and spatial scales.

9.5.1 Linking the lab with the real world Experiments conducted in the lab offer clear advantages of accessibility and controllability, and allow scientists to access a wide range of empirical tools. However, such work necessarily involves a degree of separation from the natural environment un-

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der investigation, and the severity and implications of that disconnect remain largely unaddressed. The role of pressure and temporal and spatial variability on metabolic activity of seep ecosystems warrants additional attention. 9.5.1.1 Pressure Given the range of depths at which marine methane seeps occur [169], pressure is an important variable that is frequently neglected in lab based experimental systems. When pressure has been considered, its impact has remained variable and sample dependent. Using Hydrate Ridge sediments from ∼800 m depth, methane partial pressures of 1.1 MPa led to a five-fold increase in sulfate reduction rates compared to atmospheric pressure [76]; a different study conducted at 9.0 MPa found an 80% AOM rate enhancement [49]. Experiments with Guaymas Basin sediment reported a general concordance between pressure and methane dependent sulfate reduction rate, but the relationship broke down at higher temperatures [170]. In experiments with seep sediment from the Japan Trench, however, methane driven sulfate reduction rates showed no relationship to changing pressure [171]. Investigations into the cause of any pressure related effects have similarly mixed results. Some researchers propose that rate increases are caused by greater methane solubility and bioavailability [172], while other experiments have demonstrated a six- to ten-fold AOM rate increase at 10 MPa when methane concentrations in the experimental and control (atmospheric pressure) incubations were held constant [141]. Given the range of published data, in situ pressure should be maintained throughout the sampling and experimental process for the most representative results. Activity associated tools focusing on different elements and molecules of the AOM system, or examining transcriptomic and proteomic responses to pressure, will also help to clarify the nuances of this environmentally relevant variable on seep communities. 9.5.1.2 Integrating across spatial scales One of the greatest challenges facing microbial ecologists is the integration of biochemical transformations at the nano- and microscale over larger areas to understand the impacts of kilometer-scale ecosystems. At the microbial level, synergistic relationships between neighboring cells can result in metabolite transfer or genome streamlining [173]; understanding these spatial relationships can help make sense of lineagespecific omics signals. There is also evidence that metabolic complementarity confers fitness advantages among co-occurring organisms, particularly in energy limited conditions [174], making colocalization an important trait for studies of biogeochemical cycles [175]. The spatial arrangement of, and metabolite transfer between, ANME and SRB within consortia has been a central focus of seep associated microbiology, and many of the nanoSIMS oriented studies described above (see Sect. 9.4.2.1) have been predicated on the preservation of these structures. While sample preparation and processing do

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not appear to irrevocably disrupt aggregates, any spatial associations that may be lost are not well constrained. Furthermore, the focused pursuit of the ANME-SRB coupling mechanism over the last several years has relegated other interspecies interactions to the sidelines. Given the density of redox boundaries in methane seep sediments and the diversity of seep communities, this system likely contains a number of unseen interactions that play important roles in ecosystem dynamics. Moving from a scale of dozens of microns to hundreds of meters – the necessary vantage distance from which to view a typical seep (󳶳 Fig. 9.3) – requires a number of poorly substantiated assumptions. Most notably, in situ methane concentrations (at various subseafloor depths, over meter scales horizontally, and as a function of time) are lacking. Understanding these variations, as well as methane- and sulfate-bearing subsurface fluid flow, is essential in determining how relevant microbial activity at seeps is for ecosystem services on a global scale. The current uncertainties have led to wide ranging estimates of the significance of global AOM in the carbon cycle and climate regulation [16, 176].

(a)

100 km

(c)

(d)

0.5 mm

5 cm

(e)

10 μm

(f)

10 μm

1m

(b)

Fig. 9.3: Integrating data at methane seeps across several orders of magnitude of spatial scale is a major challenge. Seagoing research expeditions target a limited segment of a geographical region (a), and submersible or robotic investigations happen on the landscape scale (b). Centimeterscale push cores (c) are retrieved, enabling investigations of porosity and grain packing via thin sections (d), phylogenetic identity of microbes via microscopy (e), or mineralogy via scanning electron microscopy (f). (Images courtesy of a: Google Earth; b: Victoria Orphan / Lisa Levin / Anthony Rathburn / WHOI; c, d: Jeffrey Marlow; e: Shawn McGlynn; f: Aude Picard)

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9.5.1.3 Temporal variations Another opportunity for improving lab based approximations of real-world settings is through consideration of temporal variation at methane seeps. Advective flow varies widely across different timescales, from hours to millennia. Benthic flux meters have captured shifts in fluid flow over the course of several weeks, from several meters per year upward to several centimeters per year downward, into the seafloor [28, 47]. Variations result from a complex combination of tidal forcing, compaction driven shifts in conduit geometry, hydrate dissociation, tectonic forces, local stratigraphy, sedimentation rate, and carbonate crust formation [28, 177]. Appreciating how localized concentrations of methane throughout a mound may change over time is important for two reasons. First, it bounds the extent of methane dependent microbial activity, providing a more realistic view of broader consumption: existing global estimates of AOM activity do not consider time dependent changes in activity [16, 176]. Second, examinations of temporal variation could demonstrate how community composition and metabolic potential respond to environmental changes. In other words, if it takes decades for a methanotrophic community to reestablish itself after a period of seep quiescence, then momentary methane concentrations could overestimate actual levels of methane consumption. Not all microbial activity at seeps is directly dependent upon methane, and as subsurface flow changes, different lineages may take advantage of shifting niches. A transplantation study of carbonate rocks from active seeps to seemingly quiescent sites showed a persistence of most taxa over 13 months, though relative abundances, particularly of putative sulfide oxidizers, did change [45]. The timescales over which microbial constituents colonize newly formed seeps – and how long it takes them to construct authigenic carbonate mounds – is unknown; sampling communities at seep complexes with a range of mound ages (e.g. Hydrate Ridge [178]), would help clarify colonization and succession processes.

9.6 Metabolic activity and tools of the future A number of emerging technologies are poised to facilitate major breakthroughs in activity based environmental microbiology. Some of these new tools are particularly well suited for ANME-SRB consortia, while others will be more generally applicable to uncultured systems. Below, we highlight several compelling opportunities on the horizon alongside specific questions related to seep hosted microbial activity that could thus become accessible.

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9.6.1 Preserving microscale spatial arrangements Analyzing microbial communities while maintaining their spatial configurations will clarify how different organisms interact. Reconstructing interspecies interactions from homogenized results of metaomics studies would allow researchers to propose interactions based on gene, transcript, or protein complementarity, but such inferences are difficult to substantiate, especially because a cell’s immediate diffusive environment severely restricts its relevant neighbors at any given time. In order to preserve spatial arrangements, one promising option could be to solidify a sample of sediment or carbonate rock with molten agar, and embed thin section slices in resin. This approach has been used on preconcentrated and density gradient separated consortia to provide detailed images of aggregate cross sections [179]. In concert with methods that retain microscale arrangements, imaging mass spectrometry (IMS) can generate two- and three-dimensional maps of protein [180] or lipid [181] markers. Matrix assisted laser desorption/ionization (MALDI) is a ‘soft ionization’ technique that mobilizes large molecules for analysis in a mass spectrometer without fragmenting them; this method, upstream of a time-of-flight mass analyzer, is the most common IMS configuration [182]. These methods have been used primarily in animal tissue analyses for biomedical purposes [183]; microbial applications have largely been for diagnostic purposes based on pure culture spectra [184]. The limiting factor that precludes mixed microbial community analysis has been spatial resolution of MALDI laser beams. Recently, however, resolution has improved to 1 µm [185], which is within the realm of usability for microbial consortia and biofilms. Determining the number of resolvable molecules from a given ‘pixel’ and their quantitation remain substantial obstacles [186].

9.6.2 Single-cell growth rate Growth rate in seep systems – as determined by increasing biomass – is most commonly determined through cell counts and/or nanoSIMS 15 N incorporation. These are both approximate methods that integrate over weeks, months, or years of incubation time; direct links between immediate, microscale environmental conditions such as methane concentration, pH, redox potential, or salinity are difficult to access. In order to determine single-cell or single-aggregate growth rates, micro- and nanoscale cantilevers can be used. These devices are narrow mechanical sensors; a fluid channel extends out and back along the perimeter in an elongated U-shape. By vibrating the cantilever, it is possible to correlate the resonant frequency to the mass of the material flowing through the channel [187]. Using a series array of cantilevers, recent designs have been able to track a single cell for a few minutes, and microbial cell

9.6 Metabolic activity and tools of the future | 245

growth rates of just 0.02 pg per hour (∼2% of a cell’s biomass [188]) can be detected. Since the flow is unidirectional, the throughput is orders of magnitude higher than previous iterations of the technology, allowing for a more representative sampling of a population [189]. To date, only eukaryotes and cultured bacteria have been tested, but presumably, mixed microbial systems could be examined as well. For example, maintaining appropriate conditions (activities of dissolved substances, temperature, etc.) through the flow path and depositing cells in an array of microchambers for single-cell sequencing would allow researchers to link identity and genetic makeup with growth rate under environmentally relevant parameters.

9.6.3 Microcalorimetry Heat production is widely believed to be the result of all metabolic reactions and a wasteful by-product of biological transformations [190]. Researchers have calculated ‘growth efficiency’ by combining experimental determinations of enthalpy and growth yields with calculations of standard enthalpy values [191]. Advancements in microfluidics and materials science have enabled smaller, more precise microcalorimeters. One recently developed device can measure heats of reaction from 3 nL volumes with 4.2 nW resolution, and was used to quantify heat generated by in vitro urease reactions [192]. Experiments with subseafloor crustal fluids comparing experimental and reference ampules in an isothermal calorimeter determined the temperature range of activity from heterotrophic sulfate reduction [193]. Associated Gibbs energy calculations yielded ∼ −13 kJ/mol sulfate, a seemingly viable value near previously proposed minimal limits [194]. While the theoretical basis and physiological meaning of biological heat production are not fully developed (for example, endergonic metabolism has been reported in M. barkeri [195]), the practice of measuring heat output through calorimetry is an intriguing capability [196, 197]. In particular, the approach is metabolism agnostic, as heat production from all metabolic reactions is integrated into a calorimetric measurement. Community-scale shifts in metabolic activity and substrate use can be observed over time, as demonstrated in a study of heterotrophy in tidal flat sediments [198]. The use of 13 C-labeled substrate linked electron acceptor depletion to changing heat outputs, while RNA-SIP revealed the microbial constituents involved. With heat production as a nearly universal product of metabolism, cryptic cycling in metal associated AOM [100] could be untangled. When coupled with activity measurements of specific metabolisms and cross referenced with thermodynamic calculations, the degree to which such metabolisms (e.g. methane oxidation, or nitrogen fixation) contribute to the overall complement of biological activity could be determined.

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9.6.4 Raman spectroscopy Microspectroscopy is an appealing technique because of its broad range of potential analytes, speed, noninvasive nature, single-cell resolution, and compatibility with other stains [199]. Raman spectroscopy excites a sample with photons of a given wavelength and detects those that undergo Stokes scattering, resulting in a lower energy, higher wave number emission [200]. The magnitude of the decreased energy is attributable in part to the masses of the atoms in the illuminated molecule. Because higher atomic masses lead to higher energies and lower wave numbers, stable isotope incorporation of > 10% can be discerned. In a study of groundwater biofilm, Raman spectroscopy was combined with FISH and 13 C amendments to reveal naphthalenedegrading Pseudomonas cells with substantial cell-to-cell variation [199]. Incorporation of D from D2 O was used as a general marker of activity to identify key players in a mouse microbiome sample and to separate them for downstream molecular analyses [201]. SIP-Raman is an important addition to the arsenal of activity based tools, but label-free spectroscopic methods could be especially helpful for in situ rate measurements. To the degree that specific molecules can be linked to certain absorption bands, the effect of integrated metabolic activity on localized concentrations can be queried. Infrared signatures of molecules like methane, ethane, and carbon dioxide would help researchers visualize and quantify how critical metabolites and greenhouse gases flow through seep systems. Strategies to accommodate such sensors at pressure and to ‘see through’ complex aqueous solutions will need to be developed, but spectroscopic techniques could allow researchers to study intact ecosystems in exciting new ways.

9.6.5 Replication rates from metagenomic data Computational treatment of metagenomic data can parse genome location-specific relative copy numbers to determine replication rates in a mixed microbial community. As a genome is copied during cell division, DNA polymerase moves from the origin of replication toward the terminus. Thus, genetic material recovered from rapidly growing populations will display a higher proportion of the origin-adjacent sequence. This approach was pioneered using gut microbiome samples [202]; when applied to environmental communities, investigators demonstrated that Candidate Phyla Radiation organisms [203] from a subsurface aquifer exhibited rare, and as yet unexplained, episodes of high growth rate [204]. Inferences of lineage-specific replication rates could open up a vast trove of metagenomic datasets to activity oriented inquiry, though interpretation of such results in a community context will require additional work.

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9.6.6 Single-cell omics The convergence of technical advances in cell sorting, microfluidics, and sequencing has created a new field of single-cell omics analyses. There are many advantages to pursuing molecular data from individual microbes in a mixed community: core genomes and coexisting subpopulations can be clarified [205], and novel diversity and new metabolic features can be uncovered [7]. In the context of activity based analyses, single-cell genomics, transcriptomics, and proteomics can all provide useful information. When paired with BONCAT-FACS [163] or SIP-Raman [201], individual genomes could be linked with biosynthetic activity, moving beyond FISH-nanoSIMS to provide a full genomic context for differentially active cells. High-throughput singlecell transcriptomics has thus far been restricted to eukaryotic samples [206], and early single-cell proteomic efforts are antibody based, limiting coverage to a few dozen proteins [207, 208]. In addition to technical difficulties associated with sequencing such small quantities of material, isolating cells in an intentional way remains challenging. Microwell dilution [209] or microfluidic droplet based methods [210] have minimal ability to discriminate between cells, while FACS can sort based on limited fluorescence associated information [211]. Integrating methods of isolating cells with traits of interest (e.g. possessing a certain gene, exhibiting high growth rates, or in close spatial association with other cells of interest) will be an important priority for single-cell analyses.

9.7 Conclusions In the decades since the discovery of seafloor methane seeps, the wide array of activity based tools has exposed a new world of metabolic possibility and revealed seeps as key capacitors in the biogeochemical processing of methane. The novel syntrophic coupling of ANME and SRB has been thoroughly explored, demonstrating a compelling case of energy sharing and electron transfer. Methanotrophic activity, in both sediments and carbonate rocks, is a substantial sink for methane that prevents vast quantities of a strong greenhouse molecule from entering the water column and, potentially, the atmosphere. However, substantial uncertainties remain, especially pertaining to the global impacts of seep associated methane consumption and the role of non-AOM metabolisms in community structure and broader ecosystem functioning. Planet-wide calculations of methane consumption have depended largely on scaling up lab based rates from point sources; understanding the nuances of subsurface flow and localized concentrations will drastically redraw the methane map. A wider set of microbial players is also well within reach [5, 6, 212]. Producers and consumers of higher hydrocar-

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bons (molecules which are found in high concentrations at methane seeps) have been largely ignored, and the ways in which AOM based primary production influences other trophic levels will no doubt reveal myriad other unforeseen interactions and metabolic networks. Methane seeps are remarkable habitats – one of the few examples of chemosynthetic oases on the seafloor – with a global reach. The tools described here will help researchers assess metabolic activity in all its forms and further clarify the role of microbial communities in driving marine methane seeps and shaping our planet.

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Barbara J. MacGregor, Beverly Flood, Jake Bailey, and Matthew Kanke

10 Multiplication is vexation: a genomic perspective on cell division and DNA replication in the large sulfur bacteria Multiplication is vexation, division is as bad; The rule of three doth puzzle me, and practice drives me mad. – Nursery rhyme

10.1 Abstract The large sulfur-oxidizing bacteria (LSB) are a morphologically diverse group of sulfide-oxidizing Beggiatoaceae (Gammaproteobacteria), including free-living and sheathed filaments, vacuolated and nonvacuolated species, and some of the largest known bacteria. The family Beggiatoaceae includes the genera Beggiatoa, Thioploca, Thiomargarita, Cand. Marithrix, and several other Candidatus groups. They are found at sulfidic/oxic interfaces in both marine and freshwater settings, including hydrothermal sites, and may store nitrate, elemental sulfur, polyphosphate, and/or carbohydrates, allowing them to tolerate environmental fluctuations or to migrate between sulfidic and oxidizing conditions. Only two species are in cultivation, both nonvacuolated Beggiatoa spp. from freshwater environments, but cells from several vacuolated species have been directly collected for genome sequencing. Comparison of these genomes with those of related Gammaproteobacteria reveals several changes that may have been key to the evolution of the LSB, in particular the loss of canonical genes for septum formation and for DNA replication initiation by DnaA. The evolutionary origin and mechanism of division of the central vacuoles found in some LSB remains a puzzle; we suggest that the acquisition of dynamin-family proteins may have been one key step. LSB genetic traits related to DNA replication have parallels to those in the Cyanobacteria, another morphologically diverse group with which the LSB appear to have a history of gene exchange. We conclude with a proposed model for the evolution of the LSB and suggest observations and experiments that could be used to test it.

https://doi.org/10.1515/9783110493672-010

262 | 10 Multiplication is vexation

10.2 Introduction The large sulfur-oxidizing bacteria (hereafter LSB) are a morphologically diverse group of Gammaproteobacteria (󳶳 Fig. 10.1) found near the boundary of sulfidic and oxic waters and sediments, in both marine and freshwater settings [1, 2]. Many of the marine morphotypes are distinguished by large central vacuoles, used by some (but not all [4]) species for storage of nitrate, presumably as an electron acceptor

(a)

(b)

(c)

(d)

Fig. 10.1: Large sulfur bacteria exhibit diverse morphologies and division characteristics. (a) Filamentous Beggiatoa-like large sulfur bacteria from Namibia. (b) Two-cell and four-cell Thiomargarita aggregates resulting from reductive cell division. (c) Ruptured Thiomargarita mother cell exhibiting apparent internal baeocytes. (d) Diverse morphologies of Thiomargarita from shelf sediments off Walvis Bay, Namibia. Scale bar in a = ∼ 600 μm; b and c = 100 μm; d = 1 mm

10.2 Introduction

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for sulfide oxidation [4]. Elemental sulfur is often stored in the periplasmic space, and glycogen [5] or polyhydroxybutyrate and polyphosphate are stored [6] in the cytoplasm. These storage capabilities are expected to enhance survival in the often shifting conditions under which these bacteria live, at sites including cold seeps, coastal upwelling zones, and hydrothermal vents and seeps. Some motile species shuttle between oxic and sulfidic conditions, alternately storing and consuming nitrate and sulfur [7, 8]. Cell division mechanisms in the LSB are varied; cells exhibiting reductive division (󳶳 Fig. 10.1b) and possible baeocyte formation (󳶳 Fig. 10.1c) have been identified microscopically. Because of their vacuoles and opaque white sulfur deposits (when present), filaments or single cells of some LSB can be seen with the naked eye, and collected at low magnification. They can be found in high concentrations where sulfide and nitrate concentrations are sufficient and reliable. Thus, although most LSB have not been cultivated, it has been possible to acquire sufficient material for genome sequencing and even proteomics [9]. Large size is also a drawback for investigators, however, because cell surfaces [10, 11] and sheath interiors [12] often host epibiont communities that make a sometimes overwhelming contribution to sequence libraries. At marine hydrothermal sites, sulfide is produced by sulfate-reducing bacteria in the anaerobic underlying sediment or mineral precipitate, and nitrate is acquired from the overlying water. The filamentous LSB studied in these environments have gliding motility [7, 13] and appear to migrate primarily within microbial mats, which they dominate in volume, although the diverse other species present may dominate numerically. For example, on the organic rich hydrothermally influenced sediments and precipitate chimneys of the Guaymas Basin, Candidatus Maribeggiatoa spp. form luxuriant orange and white mats (󳶳 Fig. 10.2). Recent (December 2016) observations suggest there may also be spherical LSB morphotypes at the borders of these mats (󳶳 Fig. 10.2; MacGregor, unpublished); it is not yet known whether or how these are motile. In vacuolated LSB, the vacuole appears to take up most of the central volume of the cells, with cytoplasm as a thin peripheral layer. This unusual geometry would seem to present a problem for cell division and DNA replication, among other cellular processes. Is the vacuole bifurcated during cell division, and if so, how? How many copies of the genome are present per cell, and how are these distributed among daughter cells? How is gene expression regulated among multiple genome copies and over long diffusional distances? How did vacuolated species evolve from nonvacuolated ones? Answering these questions has proven difficult because of a lack of LSB cultivars, but genomic data is starting to allow us to generate more hypotheses. We show here that comparison of the available genome sequences suggests that loss of several widely conserved genes for cell division and DNA replication functions may have been key to the diversification of the LSB. Following an overview of the avail-

264 | 10 Multiplication is vexation

(a)

Core 4869-21

Core 4869-02

Core 4868-18

Core 4868-13

1 mm

(b)

(c)

Fig. 10.2: Guaymas Basin microbial mats visually dominated by large sulfur bacteria. (a) Marker 6 site, Alvin dive 4563. Numerous polychaete worms can be seen, possibly grazing the mat, as well as several small crabs. The fluffy orange mat and its white fringe likely contain filamentous Cand. Maribeggiatoa spp.; the sparser bright white mat likely does not. (b) Cathedral Hill site, Alvin dive 4558. Fluffy white-bordered orange mats are seen both on the seafloor and draped over the mineral deposits of the mound. Orange and white Guaymas Basin LSB filaments belong to two (or more) different species with distinct habitat preferences [64]. (c) Large Guaymas Basin bacteria collected in December 2016. Morphotypes include thin and thick white filaments, thin pigmented filaments, and spherical and barrel-shaped cells as individuals or short (two- or three-cell) chains. The latter were collected from just outside of the mats, where temperature profiles indicated hydrothermal flow was still present. DNA sequencing is in progress (March 2017). Samples are named by Alvin dive and push core number

able LSB genome sequences, these changes will be discussed in detail. Parallels with regard to DNA replication with the Cyanobacteria, another morphologically diverse group, are then investigated. Finally, we propose a model for the evolution of the LSB, and suggest observations and experiments that might begin to test it.

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10.2.1 Overview of sequenced large sulfur bacteria The gammaproteobacterial LSB strains and species considered here are those for which relatively complete genomes are available. They have a wide range of habitats and morphologies. The two freshwater Beggiatoa are in pure culture, and one marine Beggiatoa is in coculture; the remaining genomes are from single cells, single filaments, or a large number of individual cells. Only two of them, B. leptomitiformis D-402 and Thioploca ingrica, are completely sequenced. Beggiatoa B18LDT and B. leptomitiformis D-402T are in pure culture. B. alba B18LD was isolated from freshwater rice paddy sediment [14, 15]. It grows as flexible filaments with gliding motility, and can disperse via production of necridia (sacrificial dead cells along a filament) and filament breakage. Polyhydroxybutyrate (PHB), polyphosphate, and sulfur inclusions have all been found, but no central vacuole has been reported. Its genome is estimated as 100% complete by single-copy gene complement [16], but it is not closed, consisting of one long and two short linear contigs. B. leptomitiformis D-402 is another nonvacuolated (so far as is known) species, isolated from a freshwater stream polluted by domestic sewage [17]. Its genome has been completely sequenced [18], and its physiology studied in detail. Morphologically, it is typically found as long gliding filaments, but under microaerophilic conditions its growth habit changed to short, highly motile filaments that dispersed throughout the medium. [19] So far as we are aware this has not been seen in other LSB, but laboratory observations remain sparse. Beggiatoa sp. 35Flor [20] is a marine isolate that was once thought to be in pure culture, but was later shown to be in coculture with a Pseudovibrio strain that could not be eliminated. It is found as filaments of ∼6 µm diameter, with a central vacuole (presumably for nitrate storage), and polyphosphate and sulfur inclusions. Its genome is estimated as 98% complete, based on single-copy genes [16]. The Candidatus Thiomargarita nelsonii Thio36 genome is from a single cell collected in coastal sediments of the Namibian upwelling zone, which receives large seasonal inputs of phytodetritus that produce anoxic, highly sulfidic conditions [16]. The cell was from a mucus-enclosed filament of cylindrical cells of ∼100 µm diameter. Genome completeness was estimated at 53% based on tRNAs, and 70% based on single-copy genes. Candidatus Thiomargarita nelsonii and Candidatus Thiomargarita nelsonii bud S10 are two assemblies of the same genome, obtained from a budding cell attached to a gastropod shell at the Hydrate Ridge methane seep (Oregon, USA) [21]. The bud S10 as-

266 | 10 Multiplication is vexation

sembly is the more complete, estimated as 86.53% complete from its single-copy gene complement, but misses some genes found in the first assembly, so both are included here. The species is dimorphic: elongated cells 1 mm long had spherical daughter cells at their tips, one of which was collected and washed for sequencing. The genome is notable for containing a high concentration of mobile elements such as introns, as well as genes for metacaspases, proteins that are rare in the sequenced members of most bacterial groups, including Gammaproteobacteria. In Cyanobacteria, metacaspases seem more common in species that can undergo differentiation to heterocysts. [22]. Thioploca ingrica grows as several trichomes (filaments) of 2.0–4.5 µm in diameter surrounded by a common sheath, within which the trichomes can migrate up and down. They do not have large vacuoles. The sequence was obtained from the metagenome of multiple trichomes collected from Lake Okotanpe (Japan) sediment [9]. Like B. leptomitiformis, this genome is completely closed. The Candidatus Maribeggiatoa sp. Orange Guaymas (hereafter BOGUAY, from its IMG abbreviation) was obtained from a single cleaned filament collected from a microbial mat overlying hydrothermally influenced sediments in the Guaymas Basin (Mexico) [23–25]. Pigmented filaments in these mats are ∼40 µm in diameter, with large central vacuoles. The genome is estimated as 98% complete [16], based on single-copy genes, but it is still in some 800 contigs.

10.3 Septation 10.3.1 The division and cell wall (dcw) gene cluster is fragmented in several of the large sulfur Beggiatoaceae compared to close relatives The genes and gene order of the division and cell wall (dcw) gene cluster are conserved in whole or in part across a wide phylogenetic range of bacteria (reviewed in [26]), including close relatives of the gammaproteobacterial LSB. Within the LSB, however, these genes are more dispersed, and in Thiomargarita spp. several appear to have been lost. Looking first at the ftsZ gene neighborhoods (󳶳 Fig. 10.3; 󳶳 Tab. 10.1), all the LSB genomes have shorter dcw clusters than species with closely related FtsZ genes. 󳶳 Fig. 10.3 also shows the positions of nearby putative transposases, directly repeated genes, and toxin/antitoxin genes, which could indicate previous genome rearrangements.

10.3 Septation

| 267

In genera other than Thiomargarita, all or most of the putative dcw genes not clustered with ftsZ can be found elsewhere in the genomes (󳶳 Fig. 10.4; 󳶳 Tab. 10.1). The exceptions are mraZ, not identified in any of the LSB (although it is found in closely related species), and ftsW, found in neither Thiomargarita spp. nor Thioploca ingrica. In the two completed circular genomes, T. ingrica and Beggiatoa, the dcw genes are dispersed throughout the genome compared to related Gammaproteobacteria (󳶳 Fig. 10.5). Transposase activity is one possible explanation for this rearrangement. For example, Piscirickettsia salmonis LF-89 has a complete dcw gene cluster. While it is annotated with 444 transposases, versus 178 for Beggiatoa leptomitiformis and 69 for Thioploca ingrica (󳶳 Fig. 10.5), just one of these is within 25 kb of a dcw gene, and it lies 17 ORFs past the end of the cluster. The incomplete Thioalkalivibrio paradoxus ARh 1 draft genome, the other close neighbor considered in detail (󳶳 Tab. 10.1), is annotated with just 36 transposases, with just one near a dcw gene, which likewise is found outside the main cluster. In contrast, at least one and usually several putative transposase genes are found near the dcw genes of all Beggiatoaceae investigated except Cand. Thiomargarita sp. Thio36 (󳶳 Tab. 10.1). It seems plausible that transposase activity may have contributed to the dispersal of the dcw cluster; or, from another perspective, that there was less selection against such dispersal within the LSB than other groups. Another possible contributing factor is that large bacteria with many genome copies may be subject to higher rates of recombination and rearrangement than bacteria with one or only a few genome copies, discussed further below (Section 10.5.3).

10.3.2 The available Thiomargarita genomes are missing genes for septum formation The pattern of genes apparently missing in Thiomargarita spp. becomes clear when mapped onto the known bacterial peptidoglycan biosynthesis and septation pathways (󳶳 Fig. 10.6). All or most of the available Beggiatoaceae genomes include putative genes for lateral (cell extending) peptidoglycan synthesis, and the cytoplasmic components of the early division apparatus (FtsZ, ZapA). However, the Thiomargarita genomes apparently lack the membrane components FtsA and ZipA, as well as most components of the late division ring assembly complex. These are partial genomes, so it is possible these genes were simply missed, but it seems significant that both follow the same pattern. It is also notable that whereas in most of the species shown, putative ftsQ and ftsA are found between ddl and ftsZ, in Thiomargarita these genes are directly adjacent, suggestive of a specific gene loss. This apparent loss of septal as opposed to lateral peptidoglycan synthesis in the species with more spherical morphology is in contrast to the pattern seen in some coccoid bacteria, such as Staphylococcus aureus, which only have the septal machinery and use it for all peptidoglycan production (reviewed in [27]).

268 | 10 Multiplication is vexation

FtsZ Marpu_2837 Marichromatium purpuratum 987 Thi970DRAFT_2120 Thiorhodovibrio sp. 970 TevJSym_aj00240 endosymbiont of Tevnia jerichonana (vent Tica) K257DRAFT_0145 Alteromonadaceae bacterium 2141T.STBD.0c.01a P897DRAFT_4753 Alteromonadaceae bacterium 2719K.STB50.0a.01 CUZ_03222 Sedimenticola sp. CUZ Ga0081741_11706 Sedimenticola sp. SIP–G1 NSS_00017740 Dechloromarinus chlorophilus NSS Ga0063159_01182 Cycloclasticus sp. SCGC_AC281–N15 Ga0077536_14422 Gammaproteobacteria sp. Q7A_1022 Methylophaga nitratireducens JAM1 MPL_01145 Methylophaga lonarensis MPL Q7C_863 Methylophaga frappieri JAM7 MAMP_03107 Methylophaga aminisulfidivorans MP Ga0070578_10871 Methylophaga sulfidovorans DSM 11578 G363DRAFT_01930 Thioalkalivibrio sp. ALSr1 D577DRAFT_02042 Thioalkalivibrio sp. AKL17 F790DRAFT_01107 Thioalkalivibrio sp. ALE20 D571DRAFT_00641 Thioalkalivibrio sp. ALJ24 TVNIR_2598 Thioalkalivibrio nitratireducens DSM 14787 M911_02725 Halorhodospira halochloris A ECTPHS_01904 Ectothiorhodospira sp. PHS–1 ThithiDRAFT_1491 Thioalkalivibrio thiocyanodenitrificans ARhD 1 Tgr7_0775 Thioalkalivibrio sulphidophilus HL–EbGR7 Ga0111141_10487 Acidithiobacillales bacterium SG8_45 Ga0111150_10665 Gammaproteobacteria bacterium SG8_11 Ga0111154_10824 Gammaproteobacteria bacterium SG8_47 LZ25DRAFT_01256 Thiohalospira halophila HL3 NA61DRAFT_00410 Thiohalomonas denitrificans HLD2 BOGUAY 00938_0727 Cand. Maribeggiatoa sp. Orange Guaymas Ga0060138_112953 Thioploca ingrica Ga0097846_114776, Ga0063879_01750 Cand. Thiomargarita nelsonii Ga0068239_10047 Beggiatoa sp. 35Flor BegalDRAFT_2908 Beggiatoa alba B18LD Ga0111282_1199 Beggiatoa leptomitiformis D–402 C200DRAFT_00782 Piscirickettsia salmonis LF–89 Ga0069574_100443 Piscirickettsia salmonis B1–32597 Ga0098222_112758 Legionella hackeliae ATCC 35250 Ga0124794_10779 Coxiellaceae bacterium CC99

0.01 (a)

Chromatiales unclassified Alteromonadales unclassified Thiotrichales unclassified

Thiotrichales

Chromatiales

A B C Chromatiales

D E F G H I J K L M N

Thiotrichales

Legionellales

Gammaproteobacteria Acidithiobacillia

Fig. 10.3: FtsZ inferred phylogeny and ftsZ gene neighborhoods. (a) Inferred phylogeny. Protein sequences were selected by BLASTP searches of the IMG/ER database [65], aligned using MUSCLE [66] in MEGA 5.2.2 [67], with minor adjustments to the alignments made manually, and neighbor joining used to select closest relatives. The final tree was produced using RAxML rapid bootstrapping [68] as implemented in ARB [69], using a random initial tree, the PROTGAMMA rate distribution and WAG amino acid substitution models, empirical amino acid frequencies, and branch optimization. The tree shown was the best of 25 runs. (b) Gene neighborhoods. Gene neighborhoods are from IMG/ER [65]. Full-length segments are 50 kb long, centered on the putative FtsZ genes. Putative dcw cluster genes are underlined. For Thiomargarita nelsonii, the gene neighborhood for the open reading frame (ORF) from the ‘bud S10’ assembly (Ga0063879_01750) is shown.

10.3 Septation

| 269

ftsZ Thioalkalivibrio thiocyanodenitrificans ARhD 1

⊗

SecA ArgJ

A

Thioalkalivibrio sulphidophilus HL–EbGR7

B C Thiohalospira halophilus HL3

D Thiohalomonas denitrificans HLD2

E

* *

Cand. Maribeggiatoa sp. Orange Guaymas

F

Thioploca ingrica

* G

Cand. Thiomargarita nelsonii

H Beggiatoa sp. 35Flor

I

Beggiatoa alba B18LD

⊗

* J

Beggiatoa leptomitiformis D–402

K

⊗⊗

**

Piscirickettsia salmonis LF–89 DnaA

L Piscirickettsia salmonis B1–32597 DnaA Legionella hackeliae xx

⊗ M

⊗

yy

⊗ N

⊗

*y x, (b)

Putative transposase, invertase, or protein with transposase DDE domain Putative toxin or antitoxin Direct repeats

murF

murG

ftsQ

mraY murD

murC

ftsL

ftsA ftsZ

ftsI

ftsW

mraZ mraW/rsmH

murE

murB ddl

lpxC

270 | 10 Multiplication is vexation

| 271

10.3 Septation

󳶣 Fig. 10.4: dcw cluster genes in the Beggiatoaceae. The two complete genomes, from Thioploca ingrica and Beggiatoa, are highlighted. The T. ingrica genome includes two additional putative MepM/NlpD genes on separate contigs (󳶳 Tab. 10.1). The Thio36 genome includes a second contig with fragments of a MurB-like and a Ddl-like ORF. Abbreviations: Thio36, Cand. Thiomargarita sp. Thio36; bud S10, Cand. Thiomargarita nelsonii bud S10; 35Flor, Beggiatoa sp. 35Flor; B18LD, Beggiatoa alba B18LD; B. lepto., Beggiatoa leptomitiformis D-402. ‘Thiomargarita nelsonii’ and ‘bud S10’ are two assemblies of the same genome [21]

Beggiatoa leptomitiformis D-402 (178 annotated transposases)

Piscirickettsia salmonis LF-89 (444 annotated transposases) mraW-ftsL-ftsI1-murE-murF-mraY-murD-ftsW-murG-murC1-ddl1-ftsQ-ftsA-ftsZ-lpxC 1

3100001

100001

3000001

murG-murC-murB-ddl-ftsQ-ftsA-ftsZ-lpxC

200001

2900001

ftsE

4100001

4200001

1

300001

100001 200001

4000001

2800001

300001

400001

3900001 400001

ftsW

3800001

zipA

2700001

500001

500001

ftsE-ftsX

3700001

murB-murC2

600001 2600001 700001

ftsK

3500001

murD

600001

mepM/nlpD

3600001

2500001 700001

800001

3400001 900001

2400001 800001

3300001 1000001 2300001

3200001

900001 1100001

ftsL-ftsI

murE

3100001

2200001 1200001

mraW

3000001

1300001

2100001

2900001 1400001

murF-mraY

1000001

ftsK zipA ddl2

1100001

2000001

1200001

2800001

ftsI2

1500001 1900001

1300001

2700001

mepM/nlpD

1600001

1800001

murB2

2600001 1700001

1 4700001 4800001

1700001

1800001 2400001

1900001 2300001 2200001

2100001

2000001

300001 400001

mraW-ftsL-murE-murF-mraY

4400001

ftsK

500001

4300001

600001

4200001

mepM/nlpD 2

700001

4100001

Thioploca ingrica (69 annotated transposases)

1500001

200001

4500001

ftsX

1600001

100001

4600001

2500001

1400001

800001

4000001

900001

3900001

ftsE-ftsX 1000001

3800001

1100001

3700001

zipA

1200001

3600001

murD-murG-murC-murB-ddl-ftsQ-ftsA-ftsZ-lpxC...mepM/nlpD 1 1300001

3500001

1400001

3400001

1500001

3300001

Putative transposase or invertase genes

1600001

3200001

1700001

3100001

1800001

3000001

Putative dcw and associated genes

mepM/nlpD 3

1900001

2900001 2000001

2800001 2100001

2700001 2600001

2200001 2500001 2400001 2300001

Other rings, from outside in: -- Genes on forward strand (colored by COG categories) -- Genes on reverse strand (colored by COG categories) -- RNA genes (tRNAs green, rRNAs red, other RNAs black) -- GC content, black -- GC skew, purple and green

Fig. 10.5: dcw cluster genes in complete Beggiatoaceae genomes and Piscirickettsia salmonis LF-89. Maps are from IMG/ER [65]

Description

BOGUAY

ThithiDRAFT

0662

1506

1505

1504

1503

0681

1502

1501

Inhibitor of Z ring formation

cell division transcriptional repressor

16S rRNA (cytosine(1402)N(4))-methyltransferase

membrane bound cell division leucine zipper septum protein

transpeptidase involved in septal peptidoglycan synthesis

DNA translocase/DNA segregation ATPase

UDP-Nacetylmuramoylalanyl-Dglutamate--2,6diaminopimelate ligase

UDP-N-acetylmuramoyltripeptide--D-alanyl-Dalanine ligase

zipA

mraZ

mraW

ftsL

ftsI (PBP3)

ftsK

murE

murF

2816

2815

2497

Not found

4689

4690

Not found

4026

Cand. Maribeggiatoa sp. Orange Guaymas

Thioalkalivibrio paradoxus (thiocyanodenitrificans) ARhD 1

Typical division and cell wall (dcw) cluster genes

Genome abbreviation for locus tags:

Gene

11339

11338

113536

113548

Not found

11337

11336

Not found

1142016, 107442 (short contig)

1142013

119464

103351

Not found

Not found

119463

Not found

Not found

Ga0097846

Ga006013 8

11889

Cand. Thiomargarita nelsonii

Thioploca ingrica

03603

03606

07524

07527

Not found

Not found

Not found

Not found

Not found

Not found

Ga0068239

Cand. Thiomargarita nelsonii bud S10

00042860

00026580

Not found

Not found

100447

100454

10873

10826

100267

100268

100284

100626

00073280, 00026570

Not found

Not found

10187

Ga0068239

Beggiatoa sp. 35Flor

Not found

Not found

Thi036DRAFT

Cand. Thiomargarita sp. Thio36

0742

1755

1738

1721

2094

2088

2087

3375

Not found

3357

3382

BegalDRAF T

Beggiatoa alba B18LD

112440

111003

11112, 113

11601, 600, 597, 596, 591-588

11605

112733

112734

112749, 750, 752

111142

111162

111163

Not found

113269

113262

Ga0111282

Beggiatoa leptomitiformis D-402

00773

00772

01331

00771

00770

00769

Not found

02009

02015

C200DRAFT

Piscirickettsia salmonis LF-89

Tab. 10.1: Predicted division and cell wall (dcw) cluster and related genes in the large sulfur Gammaproteobacteria and two species with closely related putative FtsZ genes. The dcw genes are in “standard” (E. coli) order. Boxes indicate groups of contiguous ORFs, with gaps where a putative gene is missing or found elsewhere in the genome. Most ORFs in each boxed group are contiguous; an exception is that BegalDRAFT_0595 and 0602 are not immediately adjacent. Black boxes indicate transposases on the same contig as the gene, within the region shown by IMG (±25 kb); above is upstream, below is downstream

272 | 10 Multiplication is vexation

1499

1939

1940

1498

1497

1496

1495

1494

1493

1492

1491

UDP-N-acetylmuramoyl-Lalanyl-D-glutamate synthetase

cell division transport system ATP-binding protein

cell division transport system permease protein

cell division-specific peptidoglycan biosynthesis regulator

undecaprenyldiphosphomuramoylpentapeptide beta-Nacetylglucosaminyltransfera se

UDP-N-acetylmuramate--Lalanine ligase

UDP-Nacetylenolpyruvoylglucosa mine reductase (UDP-Nacetylmuramate dehydrogenase)

D-alanine--D-alanine ligase

divisisome assembly protein, membrane anchored protein involved in growth of wall at septum

ATP-binding cell division FtsK recruitment protein

GTP-binding tubulin-like cell division protein

UDP-3-O-[3hydroxymyristoyl] Nacetylglucosamine deacetylase

Activator of murein hydrolase AmiC

murD

ftsE

ftsX

ftsW

murG

murC

murB

ddl

ftsQ

ftsA

ftsZ

lpxC

mepM /nlpD

1311

1473

1490

1500

phospho-Nacetylmuramoylpentapeptide-transferase

mraY

1173

0728

0727

0726

0725

0724

0723

0722

0721

0219

3484 (FtsX-like)

3485

2348 2353

2824

2818

105847, 114777

114779, 105849

112375, 112950, 11501

114776, 105846

Not found

01747

01749

01750

Not found

Not found

01751

114775, 105845

Not found

01752

01753

01754

114774, 105844

114773, 105843

114772, 105842

Not found

(Several FtsXlike permease family proteins annotated)

Not found

Not found

(Several FtsXlike permease family proteins annotated)

Not found

03602

Not found

109701, 124503, 118901

1142017, 107443

112952

112953

112954

112955

112956

112957, 113844

112958

112959

Not found

11745 (FtsX-like)

11746

112960

11340

Not found

00037340

00057350

Not found

Not found

00036830

00036820, 00046670

00046660

00048620

Not found

(Two FtsX-like permease family proteins annotated)

Not found

00057380, 00019370

00042850

102142?

10048

10047

10046

10045

10044

10737

103115

100330

104012

103029 (FtsX-like)

103122

103012

100446

3404

2909

2908

2907

2906

2905

2904

2903

2902

2900

0602

1376

0595 0596

1400

0743

11421

112260

1198

02993

00783

00782

00781

11100, 11101 1199

00780

00779

00472

00778

00777

00776

11102

11103

11104

11105

11106

11112, 111113

01243

01242

111657

00775

113483 113459 113454

00774

113043

112441

Tab. 10.1: Predicted division and cell wall (dcw) cluster and related genes in the large sulfur Gammaproteobacteria and two species with closely related putative FtsZ genes.

10.3 Septation | 273

2768

3158

2224

2562 (weak match)

0149

2817

Z-ring-associated protein A

N-acetylmuramoyl-Lalanine amidase

activator of murein hydrolases AmiA and AmiB

cell division protein FtsN

cell division protein FtsB

rod shape-determining protein RodA; cell elongation-specific peptidoglycan biosynthesis regulator RodA

zapA

amiC

envC

ftsN

ftsB

rodA

1738

1739

rod shape-determining protein MreC

rod shape-determining protein MreD

septum formation protein

mreC

mreD

maf (and maflike)

0872, 1197

1737

rod shape-determining protein MreB

mreB

2816

2769

Z-ring-associated protein B

zapB

PBP2 (pbpA)

1359

phospholipid-binding lipoprotein

mlaA

4922, 2017

0833

0832

0831

0835

0837

3036

4924 (weak match)

1651

0638

3644

3643

3642

Other cell division genes and gene clusters (transposases not shown)

01041

121942, 117783

111258, 111890

111966

111965

111964

111967

111968

111159

105687, 123551, 105024, 112008, 1144411, 101525

121353, 104681, 113112

00374, 01007

00948

00949

00950

113114 (MreB), 104683 (MreB)

104682, _113113

00947, 00765

00946, 07459

Not found

01006 (weak match)

Not found

121352, 113111

106942, 116202

Not found

Not found 1144412 (weak match), 105688 (weak match)

111708

Not found

01040

117784, 121943

Not found

01039

117785, 121944

112488 (weak match)

111343

113668

113669

113670

00070310, 00016040

00007520

10624, 106314

10282 (MreBCD all internal to different contigs) 103113 (MreBCD all internal to different contigs)

10059 (MreBCD all internal to different contigs)

00060880

10892 00062230 (MreB/Mrl family); 00068470 (scrap)

101042

104218

106913

100423

107714

10292

10291, 11151

11152

00007510

00019940 (possibly)

Not found

Not found

Not found

Not found

00010180

00010190

00010200

1311, 1996

1661

1662

1663

0708

0429

3016

1124

1292

0507

2649

2648

2647

11813, 11276

11297

11298

11299

111834

111238

112421

111243

112684

112978

112232

112233

112234

03015

03014

03013

03012

00302

00303

00457

01434 (weak match)

01422

02489

01040

01039

02118

Tab. 10.1: Predicted division and cell wall (dcw) cluster and related genes in the large sulfur Gammaproteobacteria and two species with closely related putative FtsZ genes.

274 | 10 Multiplication is vexation

10.3 Septation

| 275

Peptidoglycan biosynthesis FtsW

CYTOPLASMIC MurB

MurC MurD

MurE

FtsI septal PG

MurF

MraY MurG

lateral PG RodA

Ddl

PbpA (PBP2)

(a)

Early division ring assembly A

CM

ZipA

A

LEGEND All Beggiatoaceae Most Beggiatoaceae All but Thiomargarita All but Thiomargarita and Thioploca

(b)

Beggiatoa spp. only

Late division ring assembly

No Beggiatoaceae PG

AmiC EnvC I Q Q L B

E

CM

W

X A

A * all FtsN matches but Beggiatoa 35Flor are weak

dcw proteins not in figure MraZ: transcriptional repressor: not found in any MraW: 16S rRNA methylation: all but Thiomargarita bud S10 FtsL: membrane bound cell division leucine zipper septum protein all but Thiomargarita strains LpxC: UDP-3-O-[3-hydroxymyristoyl] N-acetylglucosamine deacetylase - all MepM/NlpD: Activator of murein hydrolase AmiC - all but Thiomargarita Thio36 (c)

Fig. 10.6: Predicted peptidoglycan biosynthesis and cell division pathways in the Beggiatoaceae. Adapted from Vicente et al. [70] CM, cytoplasmic membrane; PG, peptidoglycan layer

However, staining of Thiomargarita filaments with fluorescently labeled wheat germ agglutinin, which binds to peptidoglycan components [28], reveals what appear to be normal septa being synthesized from the cell periphery toward the center, across the vacuole (󳶳 Fig. 10.7). Whether the standard bacterial enzymes are used by Thiomargarita but have so far been missed in genome sequences, or a novel mechanism has evolved, are questions calling for future sequencing and experimental work.

276 | 10 Multiplication is vexation

Fig. 10.7: Thiomargarita cells with peptidoglycan labeled. Labeling of a chain of Thiomargarita sp. cells with wheat germ agglutinin (red) labels peptidoglycan and shows the apparent synthesis of the division septa in dividing cells (arrows). The third panel shows a three-dimensional rendering of a z-stack through wheat germ agglutinin-labeled cells showing the inside of a chain of dividing Thiomargarita cells. The concentration of newly synthesized peptidoglycan can be observed along the growing division septa (indicated by arrows). The cells were collected off of Walvis Bay, Namibia (23°00.009󸀠 , 14°04.117󸀠 , water depth: 125 m). Cells were labeled with wheat germ agglutinin, Texas Red®-X conjugate (Life Technologies). The z-stack was generated on a Nikon A1 multispectral confocal microscope with Nikon A1plus camera. Images were processed using Nikon NIS Elements version 4.51

10.3.3 What might substitute for ZipA and FtsA as an FtsZ membrane anchor in Thiomargarita septum formation? In E. coli, the FtsZ ring is tethered to the cell membrane via ZipA or FtsA [29], which interact with a conserved C-terminal FtsZ peptide. No homologs of these have been found in Thiomargarita genome sequences to date. Several FtsZ interaction partners have been considered as alternative anchors in other species (reviewed in [30]): FtsW in the Mycobacteria, which has a C-terminal extension unique to this group, but is a late recruit to the Z ring; EzrA, which is recruited early, but is typically a negative regulator of ring formation; SepF, a septation protein in gram positive bacteria; ClpX, a protease found in all or most bacteria, but not itself a membrane protein; and MinC,

10.3 Septation

| 277

which has an N-terminal FtsZ-interacting domain, but like EzrA is typically an inhibitor of Z-ring formation. No homologs of FtsW, EzrA, or SepF could be identified in the Thiomargarita genomes, but ClpX and MinC are present. Another possible FtsZ tether is MreB, part of the rod-shape determining MreBCD complex. Thermotoga maritima and E. coli MreB can both bind to cell membranes [31]: the T. maritima protein via a hydrophobic surface loop, and the E. coli protein via an N-terminal amphipathic helix that aligns parallel to the cell membrane. Most of the LSB genomes encode a putative MreB with a possible N-terminal amphipathic helix (󳶳 Fig. 10.8). The exceptions are Beggiatoa 35Flor, which has only a 2 aa helix predicted, although the helical wheel projection still shows a hydrophobic moment (not shown); and Thiomargarita sp. Thio36, whose putative mreB encodes only the C-terminal portion of the protein. The Thiomargarita nelsonii, BOGUAY, and Thioploca ingrica MreB ORFs are in neighborhoods similar to those of other Gammaproteobacteria, in possible operons with MreCD genes, while those in the other LSB are in apparently disrupted neighborhoods (󳶳 Fig. 10.9). In Beggiatoa 35Flor, the MreB, MreC, and MreD gene candidates are each interior to different contigs. The predicted MreC and

Cand. Maribeggiatoa Orange Guaymas

Cand. Thiomargarita nelsonii

Thioploca ingrica

(Ga0097846_113114)

(Ga0060138_111964)

Thiomargarita nelsonii bud S10

(BOGUAY_0831) L8

(Ga0063879_00950)

M1

L8

V5

M1 L6

L5

R4

L5

G4

A4

F9

[email protected]

F9

[email protected]

F2

G7

F9

[email protected]

L2

G7 R6

R6

K3 M1

F2

G7

R6 K3

I8

M1

L8

K3 M1

L5

L5

R4

R4 F9

[email protected]

F9

[email protected]

F2

G7

F2

Most to least hydrophilic Potentially charged

G7 R6 K3

Most to least hydrophobic

R6 K3

Beggiatoa alba B18LD

Beggiatoa leptomitiformis D-402

(BegalDRAFT_1663)

(Ga0111282_11299)

Fig. 10.8: Predicted N-terminal in-plane amphipathic helices in putative MreB proteins. Helices were predicted by AMPHIPASEEK [71, 72] from the complete sequences of the putative proteins, using the default settings. All helices shown are N-terminal. Arrows indicate the direction of the calculated amphipathy

278 | 10 Multiplication is vexation

MreB CUZ_00911 Sedimenticola sp. CUZ NSS_00028350 Dechloromarinus chlorophilus NSS Ga0081741_113393 Sedimenticola sp. SIP–G1 K257DRAFT_0660 Alteromonadaceae bacterium 2141T.STBD.0c.01a P897DRAFT_0384 Alteromonadaceae bacterium 2719K.STB50.0a.01 TevJSym_ao00750 endosymbiont of Tevnia jerichonana (vent Tica) Thi970DRAFT_3319 Thiorhodovibrio sp. 970 Thimo_2784 Thioflavicoccus mobilis 8321 C516DRAFT_03008 Arhodomonas aquaeolei DSM 8974 SPISAL_07125 Ectothiorhodospiraceae bacterium M19–40 EK23DRAFT_03499 Methylococaceae sp. 73a Ga0078419_10259 Methylococcaceae sp. B42 Q7C_1178 Methylophaga frappieri JAM7

BGS_0758 Beggiatoa sp. SS Ga0097846_113114 Ca. Thiomargarita nelsonii bud S10 BOGUAY_0831 Ca. Maribeggiatoa sp. Orange Guaymas Ga0060138_111964 Thioploca ingrica Ga0068239_10059 Beggiatoa sp. 35Flor BegalDRAFT_1663 Beggiatoa alba B18LD Ga0111282_11299 Beggiatoa leptomitiformis D–402 Ga0072641_13116 Acidiferrobacter a7 Ga0098512_113105 Aeromonas sobria CECT 4245 HMPREF1169_00064 Aeromonas veronii AER397 A941DRAFT_02213 Aeromonas veronii PhIn2 Ga0098526_11065 Aeromonas encheleia CECT 4342 Tola_0253 Tolumonas auensis DSM 9187 H027DRAFT1318 Tolumonas sp. BRL6–1 Ga0040625_10530 Agarivorans albus MKT_106 G462DRAFT_03573 Aliagarivorans taiwanensis DSM 22990 VN082462_34620 Vibrio navarrensis 08–2462 VVJY1305DRAFT_03444 Vibrio vulnificus JY_1305 Ga0069558_105095 Vibrio parahaemolyticus VPA–67 VME_20570 Vibrio harveyi 1DA3 VCV52_0411 Vibrio cholerae sv. O37_V52 VCHE09_0417 Vibrio cholerae HE–09 Ga0081957_10746 Gallaecimonas pentaromativorans YA_1 B3C1_18261 Gallaecimonas xiamenensis 3–C–1 Fbal_3464 Ferrimonas balearica DSM 9799 EK02DRAFT_3882 Ferrimonas marina DSM 16917 G506DRAFT_01862 Ferrimonas senticii DSM 18821 KT99_2134 Shewanella benthica KT99 Ga0077836_111803 Kangiella geojedonensis KCTC 23420

unclassified Alteromonadales unclassified Chromatiales A B C D E F G H I J K

Methylococcales

Thiotrichales

Acidiferrobacterales Aeromonadales

L

Alteromonadales

M Vibrionales

unclassified Alteromonadales Oceanospirillales

0.01

Gammaproteobacteria (a)

Fig. 10.9: Putative MreB proteins: Gammaproteobacterial type. (a) Inferred phylogeny. See legend to 󳶳 Fig. 10.3 for details of tree calculations. The box indicates the included Beggiatoaceae sequences, and letters indicate species for which gene neighborhoods are also shown. (b) Gene neighborhoods. Illustrations are from IMG/ER [65]. Full-length segments are 50 kb long, centered on the putative MreB genes. Putative magnetosome (mam) genes and features potentially related to gene exchange and rearrangement (transposases, directly repeated genes, restriction enzymes, toxins and antitoxins, etc.) are indicated

10.3 Septation

| 279

mreB A

Methylococaceae sp. 73a Ga0078419_10283 Endonuclease, Uma2 family

Ga0078419_10256 Endonuclease, Uma2 family

Methylococcaceae sp. B42

B

Ga0078419_10269 XerC integrase/recombinase

C

Methylophaga frappieri JAM7 Ga0063879_00973 Group II catalytic intron Ga0063879_00974 HNH endonuclease Ga0063879_00975 Reverse transcriptase

Ga0063879_00932 Restriction endonuclease

Ga0063879_00958 Putative restriction endonuclease

D

Cand. Thiomargarita nelsonii bud S10

Cand. Maribeggiatoa sp. Orange Guaymas x x ATP-binding cassettes subfamily B

E

⊗⊗

Ga0060138_111956 Endonuclease, Uma2 family

F

Thioploca ingrica Ga0060138_111951 Type I restriction enzyme M protein

G

Beggiatoa sp. 35Flor Ga0068239_10052 DDE superfamily endonuclease

⊗ H

Beggiatoa alba B18LD y y y DUF3131

I

Beggiatoa leptomitiformis D–402

⊗⊗

J

Acidiferrobacter a7 MSHA pilin biogenesis

Aeromonas sobria CECT 4245

K

Agarivorans albus MKT_106

L M

Vibrio vulnificus JY_1305

⊗

x, y

(b)

Putative transposase, invertase, or protein with transposase DDE domain Direct repeats

gatB

mreD

dat

gatA

PBP2 rodA

DUF493 lipB

gatC mreB mreC

rlpA PBP6

280 | 10 Multiplication is vexation

MreB/MamK (BOGUAY_2204) Ga0081909_115518 MreB Candidatus Magnetomorum sp. HK–1 OMM_00509 MreB Candidatus Magnetoglobus multicellularis Araruama McasDRAFT_01981 MreB Candidatus Magnetobacterium sp. MYR–1 DMR_40940 MamK Desulfovibrio magneticus RS–1 ME12555DRAFT_00003 MreB Lake Mendota epilimnion metagenome PSMK_08370 MreB Phycisphaera mikurensis NBRC Ga0104840_114103 MreB Halothermobacillus malaysiensis RA NPH_2401 conserved hypothetical protein Deltaproteobacterium NaphS2 Ga0073139_18135 MreB Delisea pulchra microbial communities affected by bleaching Ga0073305_0266 MreB Methanolobus vulcani PL 12/M MettiDRAFT_2438 MreB Methanolobus tindarius DSM 2278 Ga0073303_1845 MreB Methanolobus profundi Mob M Mpsy_0918 MreB Methanolobus psychrophilus R15 MG2_0177 MreB uncultured Marine Group II euryarchaeote Ga0111191_11095 MreB bacterium SM23_31 Ga0111191_11203 MreB bacterium SM23_31

BOGUAY 01051_2204 MreB Cand. Maribeggiatoa sp. Orange Guaymas Ga0060138_112692 MreB Thioploca ingrica Dlim_32010 Actin–like ATPase involved in cell morphogenesis Desulfonema limicola Ga0081909_1142510 MreB Candidatus Magnetomorum sp. HK–1

Ga0068239_104826 MreB Beggiatoa sp. 35Flor Thini_4345 MreB Thiothrix nivea JP2 DSM 5205 H261_19154 MamK MreB Magnetospirillum sp. SO–1 amb0965 MamK Actin–like ATPase Magnetospirillum magneticum AMB–1 Ga0077730_1116 MamK MreB Magnetospirillum magnetotacticum MS–1 Ga0081727_112336 MamK MreB Magnetospirillum gryphiswaldense MSR–1 Ga0072465_111017 MreB Magnetospira sp. QH–2 Mmc1_2259 Actin–like ATPase Magnetococcus sp. MC–1 MldDRAFT_1431 Actin–like ATPase Deltaproteobacterium MLMS–1 MldDRAFT_3284 Actin–like ATPase Deltaproteobacterium MLMS–1 DaAHT2_2571 Actin/actin family protein Desulfurivibrio alkaliphilus AHT2 DaAHT2_2569 Actin/actin family protein Desulfurivibrio alkaliphilus AHT2 MldDRAFT_1429 Actin/actin–like Deltaproteobacterium MLMS–1 Ga0073655_110339 MreB Opitutaceae bacterium EBPR Oter_3137 Actin/actin family protein Opitutus terrae PB90–1 Ga0081615_104734 MreB Opitutae 129 (UID2982) Ga0073125_113110 MreB Delisea pulchra microbial communities affected by bleaching VDG1235_3034 hypothetical protein Verrucomicrobiae bacterium DG1235

A B C D

E F G H

0.05

ARCHAEA Gammaproteobacteria Alphaproteobacteria Deltaproteobacteria Nitrospirae Verrucomicrobia Planctomycetes CFB group Unclassified

Magnetotactic species Magnetotactic species, ORF found near other putative magnetosome proteins

(a) Fig. 10.10: Putative MreB proteins: MamK-like. (a) Inferred phylogeny. See legend to 󳶳 Fig. 10.3 for details of tree calculations. Letters indicate species for which gene neighborhoods are also shown.

10.3 Septation |

281

mreB / mamK Cand. Magnetoglobus multicellularis Araruama mamP2EBAQMP1T

ferrous iron transport protein B mamO

A Desulfovibrio magneticus RS–1

⊗mamA mamB B

mamE

mamB

⊗⊗ ⊗

Halothermobacillus malaysiensis RA murQ

N-acetylmuramoyl-L-alanine amidase

C gyrB

Capsule assembly protein Wzi

recF

Deltaproteobacterium NaphS2 N-acetylmuramoyl-Lalanine amidase

⊗

integrase core domain protein

D antitoxin

Cand. Maribeggiatoa sp. Orange Guaymas N-6 DNA methylase

purC

E

Nucleoside-diphosphate-sugar epimerase

antitoxin LPS assembly

UDP-N-acetylmuramyl pentapeptide phosphotransferase/ UDP-N-acetylglucosamine-1-phosphate transferase

Thioploca ingrica

x

x

F dnaN

Beggiatoa sp. 35Flor G

⊗

⊗⊗

Magnetospirillum magneticum AMB–1

⊗

magnetic particle protein

mamM

H Membrane-fusion protein

(b)

Bacterial magnetic particle proteins

⊗

Putative transposase, invertase, or protein with transposase DDE domain

x

Direct repeats

Gammaproteobacteria Alphaproteobacteria Deltaproteobacteria CFB group

Fig. 10.10: Putative MreB proteins: MamK-like. (b) Gene neighborhoods. Illustrations are from IMG/ER [65]. Full-length segments are 50 kb long, centered on the putative MreB genes. Conserved gene clusters are underlined and features potentially related to gene exchange and rearrangement (transposases, directly repeated genes, restriction enzymes, toxins and antitoxins, etc.) are indicated

282 | 10 Multiplication is vexation

MreD genes from all of the available LSB genomes group together in phylogenetic trees (not shown); the closest neighbors vary, but nearly all are Gammaproteobacteria. BOGUAY, Beggiatoa 35Flor, and T. ingrica each also have a second possible MreB, which seems likely to have been acquired by horizontal gene transfer, possibly from magnetotactic bacteria (󳶳 Fig. 10.10). Magnetotactic bacteria are common in sediments [32], and in fact Desulfonema spp. may inhabit Thioploca sheaths [12], providing opportunities for gene exchange. MamK forms MreB-like filaments along which magnetosomes align (reviewed in [33]). None of these MamK-like MreB, nor any of the other sequences in the tree, are predicted to have N-terminal amphipathic helices. No other Mam proteins were found in the LSB; it is unlikely they produce magnetosomes. Multiple MreB-like proteins are also found in other bacteria (mentioned in [31]). All of the LSB except possibly Beggiatoa 35Flor and Thiomargarita sp. Thio36 therefore may possess MreBs that could link FtsZ to the cell membrane, and potentially play a role in cell division. If they have taken on this role, however, they have done so without diverging noticeably from other gammaproteobacterial MreB. Nor do the predicted FtsZ proteins appear unusual. Perhaps there have been significant changes at the regulatory level; for example elevated mreB expression.

10.3.4 How might central vacuoles divide? Little is known yet about the central vacuoles of the LSB: what sort of membrane bounds them, is it completely separated from the inner membrane or is it divided between daughter cells? The most straightforward hypothesis is that it arose as an invagination of the cell membrane, but is now a separate compartment, divided between daughter cells at the septation point. This simple scenario raises some mechanistic questions, however. If, as proposed above (Section 10.3.2), septal peptidoglycan is synthesized from the outside in, the inner membrane surrounding it will impinge on the vacuole membrane from all sides, unless this is constricting in the same place at (or before) the same time (see figure 21 in [34]). Perhaps the mechanical force of the septum alone is sufficient to deform the membrane, but there are other possibilities (not necessarily exclusive). If there is a FtsZ-like ring, it might push on the vacuole membrane (VM) at the same time as it pulls on the inner membrane (IM). In the invagination scenario, the inner face of the IM and the outer face of the VM are topologically equivalent, and could house the same FtsZ-interacting proteins. Alternatively, proteins within the vacuole lumen could be pulling it together. Whatever the driving force, as the septum closes, the existing VM and IM leaflets will need to be split and rejoined, presumably without losing vacuole or cytoplasmic contents. Several classes of membrane-fusing proteins are known, including dynamins and dynamin-like proteins, vesicle-inducing protein, and phage shock protein A, and possible examples of all of these are found in most LSB genomes. A preliminary survey of dynamin-like proteins in the LSB found three clusters, whose inferred phy-

10.4 Disruption of some gene clusters seems characteristic of marine LSB

| 283

logenies all suggest horizontal gene transfer. Shown here is one group represented in all sequenced LSB (󳶳 Fig. 10.11), with closest relatives among several other proteobacterial groups. The Conserved Domain Database (CDD [35]) predicts Dynamin_N (pfam 00350) domains in all except the Thiomargarita sp. Thio36 predicted protein, which is encoded at the end of a very short contig and likely incompletely sequenced. The amino acid sequences were checked for predicted coiled coil domains, which may be involved in protein-protein interactions in some dynamins [36]. A high probability of coiled-coil formation was found near the amino terminus of most of the LSB, with lower probabilities for Beggiatoa sp. 35Flor, B. leptomitiformis, and especially B. alba; the possibly truncated Thiomargarita Thio 36 sequence had none at all. All other sequences shown in the tree had little to no predicted N-terminal coiled coil probability. This dynamin-family protein may therefore have taken on a new function in the LSB, particularly in the vacuolate species. The gene neighborhoods are suggestive: in B. alba and B. leptomitiformis, a predicted cell-division checkpoint protein gene is found immediately downstream, whereas in Cand. Thiomargarita nelsonii, the unpublished Guaymas Basin ‘bin 4572_84’, and T. ingrica there is a putative MurJ flippase gene one ORF upstream. In the BOGUAY genome, as well as in the remaining species examined, the immediately neighboring ORFs have no obvious connection to cell division processes. Speculatively, this set of dynamin-family proteins may have adapted to different roles in cell division in morphologically different LSB.

10.4 Disruption of some gene clusters seems characteristic of marine LSB The dcw gene cluster is not the only one that is disrupted in the large sulfur bacteria. Considering only the relatively complete genomes, the predicted NADH dehydrogenase (nuo) genes are nearly contiguous in B. alba B18LD and B. leptomitiformis D-402, interrupted only by a putative antitoxin gene between nuoI and nuoJ (󳶳 Fig. 10.12a). In Thioploca ingrica, an antitoxin gene is found in the same position, but there are additional intervening ORFs and the nuoA-F and nuoG-N segments are at very different chromosomal positions (not shown). In the Thiomargarita nelsonii bud S10 genome assembly, the putative NuoA-M genes are found together, again with several intervening ORFs, but nuoN and a second nuoF are on separate contigs and cannot be joined directly to nuoA-M. Clues to possible rearrangement mechanisms in this species are putative xisHI genes (required for gene rearrangements leading to heterocyst formation in some Cyanobacteria [37]) just downstream of nuoM, and CRISPR-associated proteins several ORFs downstream of nuoN. The ATP synthase gene cluster is disrupted on a smaller scale (󳶳 Fig. 10.12b) in some LSB. It is contiguous in B. alba B18LD and B. leptomitiformis D-402, but there are two or three ORFs between atpA and atpG in T. ingrica and Cand. T. nelsonii bud S10

284 | 10 Multiplication is vexation

Putative dynamin-family protein 14 Nitrosomonas spp. 8 Nitrosomonas spp. Ga0105872_11533

Nitrosomonas marina Nm71

Ga0105871_10236

Nitrosomonas marina Nm22

Ga0111712_103115

Nitrosomonas sp. 51 Nm51

A

Ga0105860_103244

Nitrosomonas aestuarii Nm69

B

Bacteriovoracaceae bacterium EBPR Bin 340

C

Ga0073645_14582

Ga0097846_100394

Cand. Thiomargarita nelsonii

Thi036DRAFT_00058220** BOGUAY_2229 BGP_3843**

Cand. Maribeggiatoa sp. Orange Guaymas

Cand. Isobeggiatoa sp. PS

Ga0123547_11864 Ga0060138_113722

E F G

“Beggiatoa” sp. bin 4572_84

H

Thioploca ingrica

I

Beggiatoa alba B18LD

J

Ga0100909_11979

Beggiatoa leptomitiformis

K

Mmc1_0896**

Beggiatoa sp. 35Flor

Magnetococcus sp. MC-1

Ga0078419_138602 H035DRAFT_1776

MCA1019** Ga0131102_0017

N

Methylohalobius crimeensis 10Ki

O

Nitsa_1129**

Methylococcus capsulatus ATCC 19069

P

Methylococcus capsulatus str. Bath

Thiomicrospira sp. Milos_T2_DSM_13229

Thiomicrospira KP2

L868DRAFT_1860 Ga0104425_11314

M

Sulfurivirga caldicuralii DSM 17737

BS34DRAFT_2372 tkp2_00003740

L

Methylococcaceae sp. B42

H156DRAFT_2753*

NitteDRAFT_1487

D

BegalDRAFT_0549

Ga0068239_10617

NIS_0953**

Thiomargarita sp. Thio36

Thiomicrospira sp. Milos T1

Hydrogenimonas thermophila EP1-55-1

Nitratifractor salsuginis DSM 16511

Nitratiruptor sp. SB155-2

Nitratiruptor tergarcus DSM 16512 0.10

(a)

Gammaproteobacteria Betaproteobacteria Alphaproteobacteria Epsilonproteobacteria Deltaproteobacteria

Fig. 10.11: Predicted dynamin-family proteins. (a) Inferred phylogeny. See legend to 󳶳 Fig. 10.3 for details of tree calculations. Three Bacillus sequences were used to root the tree. Letters indicate species for which gene neighborhoods are also shown. All were annotated as dynamin family proteins except those with asterisks. *: translation elongation factors, GTPases. **: hypothetical or conserved hypothetical protein.

10.4 Disruption of some gene clusters seems characteristic of marine LSB

Predicted protein domains and coiled-coil probabilities

| 285

Putative dynamin-family protein gene neighborhoods

Nitrosomonas aestuarii Nm69

A

1 0.8 0.6 0.4 0.2 0

⊗

Nitrosomonas aestuarii Nm69 1 0.8 0.6 0.4 0.2 0

B

Bacteriovoracaceae bacterium EBPR Bin 340 1 0.8 0.6 0.4 0.2 0

C

Cand. Thiomargarita nelsonii

Ga0097846_100393 ribF riboflavin kinase/ FMN adenylyltransferase Ga0097846_100392 murJ, mviN putative peptidoglycan lipid II flippase

1 0.8 0.6 0.4 0.2 0

D

Thiomargarita sp. Thio36 1 0.8 0.6 0.4 0.2 0

E

Cand. Maribeggiatoa sp. Orange Guaymas

F

BOGUAY_2228 adenylyltransferase and sulfurtransferase

1 0.8 0.6 0.4 0.2 0

BOGUAY_2230 FAD/FMN-containing dehydrogenase

Cand. Isobeggiatoa sp. PS 1 0.8 0.6 0.4 0.2 0

G

Ga0123547_11863 ribF riboflavin kinase/ FMN adenylyltransferase Ga0123547_11862 murJ, mviN putative peptidoglycan lipid II flippase

“Beggiatoa” sp. bin 4572_84 1 0.8 0.6 0.4 0.2 0

H

Ga0060138_113723 ribF riboflavin kinase/ FMN adenylyltransferase Ga0060138_113724 murJ, mviN putative peptidoglycan lipid II flippase

Thioploca ingrica

I

1 0.8 0.6 0.4 0.2 0

Beggiatoa alba B18LD

J

BegalDRAFT_0548 cell division checkpoint GTPase YihA

1 0.8 0.6 0.4 0.2 0

Beggiatoa leptomitiformis

⊗⊗

1 0.8 0.6 0.4 0.2 0

K

Ga0111282_11980 cell division checkpoint GTPase YihA

Beggiatoa sp. 35Flor 1 0.8 0.6 0.4 0.2 0

L

Magnetococcus sp. MC-1 1 0.8 0.6 0.4 0.2 0

M

Methylococcaceae sp. B42

GtrA

1 0.8 0.6 0.4 0.2 0

N

Methylohalobius crimeensis 10Ki YfcA

1 0.8 0.6 0.4 0.2 0

O

Methylococcus capsulatus ATCC 19069 MotE

P

1 0.8 0.6 0.4 0.2 0

1 0.8 0.6 0.4 0.2 0

Dynamin_N domain COILS window width

(b)

window=14 window=21 window=28

Gammaproteobacteria Betaproteobacteria Alphaproteobacteria Deltaproteobacteria

⊗

Putative transposase, invertase, or protein with transposase DDE domain

(c)

Fig. 10.11: (b) Predicted conserved and coiled-coil domains. Protein domains were predicted using the Conserved Domain Database (CDD [35]) and coiled coil domains by COIL [73], with the default settings. (c) Gene neighborhoods. Illustrations are from IMG/ER [65]. Full-length segments are 50 kb long, centered on the putative dynamin-family protein genes. Predicted functions (if any) of ORFs immediately adjacent to these on the same strand are shown, and features potentially related to gene exchange and rearrangement (here just transposases) are also indicated

286 | 10 Multiplication is vexation

NADH dehydrogenase Beggiatoa alba B18LD

ABC D E F

G

HI

JK L

BegalDRAFT_1900 antitoxin, HigA family

M N

Ga0111282_112013 Antitoxin YwqK

Beggiatoa leptomitiformis D-402 ABC D E F

Thioploca ingrica

BegalDRAFT_1899 proteic killer suppression protein

BegalDRAFT_1887 Antitoxin YwqK

G

HI

JK L

Ga0111282_112002, 112001 Transposases

M N

Ga0060138_111000 YafN antitoxin Ga0060138_11999 Hyp prot Ga0060138_111002 DUF2442 Ga0060138_11998 Endonuclease, Uma2 family Ga0060138_111003 DUF4160 Ga0060138_11996 Protein N-acetyltransferase, RimJ/RimL family ABC D

E

F

Ga0060138_112535 YwqK antitoxin Ga0060138_112536 Cyclic nucleotide-binding domain-containing protein

G H

Thiomargarita nelsonii bud S10

I

JK L

M N

Ga0063879_03327 Uncharacterized SAM-binding protein YcdF Ga0063879_03325, 03324, 03322 Hypothetical proteins Ga0063879_03321 probable antitoxin Ga0063879_03317 YwqK antitoxin Ga0063879_03312 XisH Ga0063879_03311 XisI

ABCD E F1

G H I JKL

M

Ga0063879_05611 succinate dehydrogenase / fumarate reductase flavoprotein subunit Ga0063879_05609 AAA-like domain-containing protein

(a)

F2

ATP synthase N CRISPR-associated proteins

Beggiatoa alba B18LD I BEFH A G DC

Beggiatoa leptomitiformis D-402 Ga0111282_112002, 112001 Transposases

I BEFH A G DC Ga0060138_112551 riboflavin synthase alpha chain Ga0060138_112552 hypothetical protein

Thioploca ingrica

Thiomargarita nelsonii bud S10

(b)

I BEFH A

GDC Ga0063879_02406, 2407, 2408 hypothetical proteins

I BEFH A

G DC

Fig. 10.12: Putative (a) NADH dehydrogenase and (b) ATP synthase gene neighborhoods in relatively complete Beggiatoaceae genomes. Putative genes interrupting or flanking these gene clusters and those potentially related to gene exchange and rearrangement are labeled

respectively. The ribosomal protein clusters examined, however, are similar to those in related Gammaproteobacteria (not shown).

10.5 DNA replication |

287

10.5 DNA replication Bacterial DNA replication origins can often be recognized in genome sequences by: the location of dnaA (encoding chromosomal replication initiator DnaA), dnaN (encoding DNA polymerase III subunit beta), and several other characteristic genes; the distribution of nonamer DnaA binding sites; and a switch in the strandedness of GC skew, which occurs again at the terminus [38]. Exceptions to all of these features are known, however [39]. The commonly occurring origin and terminus regions are known as oriC and terC (‘C’ for chromosomal) respectively. GC skew is the preponderance of G over C in the lagging compared to the leading strand, apparently selected for because C is less susceptible to mutation on the leading strand. In bacteria with circular chromosomes, DNA replication typically proceeds bidirectionally from a single origin to a single terminus. Because DNA polymerase moves 5󸀠 to 3󸀠 , only one strand in each direction – the leading strand – is replicated continuously as the replication fork moves forward. The lagging strand is replicated in segments (Okazaki fragments) that are later joined together, and is therefore single stranded for a greater fraction of the replication time. This makes it more vulnerable to spontaneous deamination of C to U (or methylated C to T), and the mutations are repaired less efficiently at least in part because of the absence of an undamaged complementary strand [40].

10.5.1 LSB are lacking some or all of the typical gene features of bacterial replication origins From the available genome sequences, the large sulfur Gammaproteobacteria are missing some or most of the features of typical bacterial (especially gammaproteobacterial) origins of replication, compared to close relatives (󳶳 Fig. 10.13b, 󳶳 Tab. 10.2). B. alba B18LD and B. leptomitiformis D-402 are most similar to other bacteria, with yidC, rnpE, and rpmH on one strand and dnaN and dnaA on the other. yidD was not found in B. alba B18LD or several of the other species shown. The Beggiatoa sp. 35Flor genome does include the yidC cluster, but its putative dnaA and dnaN are each internal to separate contigs rather than adjacent. The remaining LSB genomes include a putative dnaN, and yidC nearby on the opposite strand, but lack a recognizable DnaA chromosomal replication initiation gene. Absence of dnaA was previously noted for T. ingrica [9]. For all of the LSB, other genes typical of bacterial oriCs (e.g. parAB) are either scattered or not found (󳶳 Tab. 10.2). DnaA is also apparently missing in several other, mostly symbiotic bacteria [41–43], and is not required for replication in all Cyanobacteria [44]; possible alternatives are discussed below (Section 10.5.4). Absence of DnaA is not necessarily a characteristic of large bacteria with multiple genome copies: the genome of Epulopiscium sp. N.t. morphotype B [45], a member of the largest known bacterial genus with many genome copies, has one contig ending

288 | 10 Multiplication is vexation DnaN Idiomarina loihiensis L2TR Idiomarina abyssalis KMM 227 Ga0113964 1062 Idiomarina sp. REDSEA–S27_B4 OS145_03295 Idiomarina baltica OS145 Ga0061064_1836 Idiomarina woesei DSM 27808 A10D4_08077 Idiomarina xiamenensis 10–D–4 A28LD_0538 Idiomarina sp. A28L cola_00002 Glaciecola sp. ANT9081 GARC_0069 Paraglaciecola (Glaciecola) arctica BSs20135 GLIP_0031 Aliiglaciecola (Glaciecola) lipolytica E C727DRAFT_00540 Gayadomonas joobiniege G7 Ga0081026_10784 Gammaproteobacteria bacterium Q1 F595DRAFT_0419 Perlucidibaca piscinae DSM 21586 Ga0077836_113 Kangiella geojedonensis KCTC 23420 BGP_4987 Cand. Isobeggiatoa sp. PS Ga0060138_11800 Thioploca ingrica Ga0097846_1123712, Ga0063879_00440 Cand. Thiomargarita nelsonii Thi036DRAFT_00024190 Thiomargarita sp. Thio36 BOGUAY 00696_0324 Cand. Maribeggiatoa sp. Orange Guaymas Ga0068239_104825 Beggiatoa sp. 35Flor BegalDRAFT_1134 Beggiatoa alba B18LD Ga0111282_112 Beggiatoa leptomitiformis D–402 A3GODRAFT_00657 Sedimenticola selenatireducens AK4OH1 DSM 17993 CUZ_00230 Sedimenticola sp. CUZ Ga0081741_112 Sedimenticola sp. SIP–G1 NSS_00034120 Dechloromarinus chlorophilus NSS Ga0077554_1417 Gammaproteobacteria sp. Ga0074134_14511 Gammaproteobacteria sp. S7S_00823 Alcanivorax pacificus W11–5 MUS1_00750 Marinomonas ushuaiensis DSM 15871 IL0002

Ga0073293_1497

Mmwyl1_0002

Alteromonadales

A B C D E F

Pseudomonadales Oceanospirillales

G H I J K L M N O P

Thiotrichales

unclassified

Q

Marinomonas sp. MWYL1

M320_pool_00021720

Marinomonas sp. GOBB3–320

D104_11045 Marinomonas

sp. D104

Marinomonas posidonica IVIA–Po–181 Marinomonas mediterranea MMB–1 Ga0061065_104106 Marinomonas fungiae JCM 18476 Ga0073128_14665 microbial metagenome of Delisea pulchra affected by bleaching disease MED121_10475 Marinomonas sp. MED121

Oceanospirillales

Mar181_0002

Marme_0002

unclassified Oceanospirillales

0.50 Gammaproteobacteria (a)

Unclassified

Fig. 10.13: (a) DnaN inferred phylogeny. See legend to 󳶳 Fig. 10.3 for details of tree calculations. The bracket indicates the included Beggiatoaceae sequences, and letters indicate species for which gene neighborhoods are also shown

10.5 DNA replication

|

289

dnaN Glaciecola sp. ANT9081 dnaA

A GARC_0059 DNA replication protein DnaC

GARC_0080 DNA polymerase I polA

Paraglaciecola arctica BSs20135

B GLIP_0043 DNA polymerase I polA

GLIP_0021 DNA replication protein DnaC

Aliiglaciecola lipolytica E

C

cola_04328 DNA replication protein DnaC

cola_00013 DNA polymerase I polA

Gayadomonas joobiniege G7

C727DRAFT_00552 cell division checkpoint GTPase YihA

C727DRAFT_00561 Cell division protein DamX

D Perlucidibaca piscinae DSM 21586

E Kangiella geojedonensis KCTC 23420

F Thioploca ingrica

G Cand. Thiomargarita nelsonii

No dnaA found

H Cand. Maribeggiatoa sp. Orange Guaymas

I

Beggiatoa sp. 35Flor

Ga0068239_104826 rod shape-determining protein MreB

J

⊗⊗

Beggiatoa alba B18LD

K

⊗*

Beggiatoa leptomitiformis D–402

L Sedimenticola selenatireducens DSM 17993

M Sedimenticola sp. CUZ

N Sedimenticola sp. SIP–G1

O Dechloromarinus chlorophilus NSS

P Alcanivorax pacificus W11–5

Q

⊗ BegalDRAFT_1131 DNA polymerase III, delta subunit

BegalDRAFT_1124 cell division protein

BegalDRAFT_1126 DNA polymerase III, epsilon subunit and related 3'-5' exonucleases

⊗ With dnaA

*

Putative transposase, invertase, or protein with transposase DDE domain Putative toxin or antitoxin

parB family chromosome partitioning protein parA chromosome partitioning protein gidB 16S rRNA m(7)G-527 methyltransferase gidA tRNA uridine 5-carboxymethylaminomethyl modification enzyme trmE tRNA modification GTPase yidC protein translocase subunit rnpA ribonuclease P protein component rpmH LSU ribosomal protein L34P

dnaA chromosomal replication initiator dnaN DNA polymerase III, beta subunit recF DNA replication and repair protein gyrB DNA gyrase subunit B

(b) Fig. 10.13: (b) dnaN gene neighborhoods. Gene neighborhoods are from IMG/ER [65]. Full-length segments are 50 kb long, centered on the putative DnaN genes. Putative oriC region genes are underlined

Description

112349

Membrane-anchored protein YidD

yidD

111490

111489

segregation and condensation protein A

segregation and condensation protein B

scpA

scpB

1609

112019

112018

11890

4673 (RecF/RecN/SMC N-term. domain)

111058

condensin subunit Smc

smc

1608

11321

1161

gyrB

recF

Not found

11800

115

0324

113

11801 Not found

4567

Not found

112

114

chromosomal replication initiator DNA polymerase III, beta subunit

11802

112385

11803, 11325

112415

111723

101663, 108507, 117797

108506, 101662, 117796

111812

113733, 105271

Not found

1123712

Not found

1123711, 104233

1123710, 104232

Not found

112379, 115548, 101783, 106408

111685, 122154

1164513, 110321

103171, 112616, 104385, 100772

03224

03225

01250

05475

07340, 05852

Not found

Not found

00439

00438

Not found

00437, 02503

04901

Not found

04091

01482, 01518, 00934, 00499 (fragment), 05995, 04957

111953, 11267, 112783, 11473

113074

07475

103173

Ga0063879

Cand. Thiomargarita nelsonii bud S10

103172, 112485, 112911, 113718/105222 (probably identical), 110794/119793 (probably identical), 108153/117423 (probably identical)

Ga0097846

Cand. Thiomargarita nelsonii

112564

Ga0060138

Thioploca ingrica

DNA replication and repair protein DNA gyrase subunit B

dnaN

dnaA

rpmH

0320

0319, 4719

1777

0322

112348

protein translocase subunit YidC

yidC/yajC or yidC/oxa1 family

111

112347

tRNA modification GTPase

trmE/mnmE

2372

0321

112343

tRNA uridine 5carboxymethylaminomethyl modification enzyme

gidA/mnmG

1756

2635, 3225

Not found

BOGUAY

Cand. Maribeggiatoa sp. Orange Guaymas

112350

112341

16S rRNA m(7)G-527 methyltransferase

gidB/rsmG

ribonuclease P protein component LSU ribosomal protein L34P

112340

chromosome partitioning protein

parA/soj or parA/mrp

rnpE

112339

parB/spo0J

Ga0077836

Kangiella geojedonensis KCTC 23420

chromosome partitioning protein, ParB family

Genome abbreviation for locus tags:

Gene

Not found

00029790, 00029810

32290

Not found

Not found

24190

Not found

Not found

Not found

Not found

Not found

10890

33570

59830

00057790, 00024150

11054

10552

10281

100455

Not found

104825

10734

11402

11403

11404

11405, 102814

10953

103922

10593

100143

Not found

Ga0068239

Thi036DR AFT Not found

Beggiatoa sp. 35Flor

Cand. Thiomarga rita sp. Thio36

1644

2146

1657

2330

Not found

1134

1135

1136

1137

Not found

1138, 3486

152

1142

224

1994

Not found

BegalDRAFT

Beggiatoa alba B18LD

113015

111875

11125

112754

Not found

112

111

113683

113682

113681

113680, 111669

11231

113614

112960

11815

Not found

Ga0111282

Beggiatoa leptomitiformis D-402

03917

03916

02858

00655

00656

00657

00658

Not found

Not found

02134

00660

00661

00679

00680

00684

00685

A3GODRAFT

Sedimenticola selenatireducens AK4OH1 DSM 17993

Tab. 10.2: Distribution of typical oriC region genes in large sulfur Gammaproteobacteria and two species with closely related putative DnaN genes. Bold outlines indicate ORFs contiguous in a given genome

290 | 10 Multiplication is vexation

10.5 DNA replication

GC Skew Diagrams for Finished Genomes

Idiomarina loihiensis L2TR

|

291

Average segment Skew shifts per 1 Gb length 37720

51

(F) Kangiella geojedonensis KCTC 23420

72497

19

(G) Thioploca ingrica

13865

310

(L) Beggiatoa leptomitiformis D-402

15294

201

(O) Sedimenticola sp. SIP-G1

101400

12

Marinomonas sp. MWYL1

62923

19

Marinomonas posidonica IVIA-Po-181

79414

19

1.0

Marinomonas mediterranea MMB-1

67727

20

-1.0

No dnaA found

0

GC skew

(c) Fig. 10.13: (c) GC skew of finished genomes. GC skew was calculated using a 10 kb moving window with 100 bp steps. Skew for a given window was calculated as (total_G_count)/(total_G_count + total_C_count), giving a value between 0 and 1. This is displayed as (0.5 − (skew_of_window ⋅ −1)), with G skew positive and C skew negative.

in dnaA (Epulo_02633) and another beginning in dnaN (Epulo_02570), each with the same collection of neighboring genes as close relatives in the Clostridia. Assuming these contigs are in fact linked, the presumed origin of DNA replication seems unexceptional in this species. Its ftsZ gene neighborhood is also comparable to that of other Clostridia (not shown).

10.5.2 The two complete LSB genomes lack GC skew The two complete LSB genome sequences, from B. leptomitiformis D-402 and Thioploca ingrica, both lack GC skew (󳶳 Fig. 10.13c), which, like the absence of dnaA, has already been pointed out in T. ingrica [9]. Absence of skew is relatively rare in most bacterial groups, as reviewed by Mackiewicz et al. for all complete bacterial genomes available in 2004 [39] and partially updated here (󳶳 Tab. 10.3). Endosymbionts appear several times in 󳶳 Tab. 10.3, and it seems intuitive that they might be subject to different selective pressures than free-living bacteria. However, not all endosymbionts lack skew. Of 37 complete circular endo-symbiont genomes currently (November 2016) in IMG/ER, 17 are strongly skewed by visual inspection of the provided chromosome maps, and at least five more are weakly skewed (not shown). There is no clear correlation between presence or absence of putative dnaA or dnaN, or whether they are adjacent on the chromosome, and the degree of

Deinococcus radiodurans R1 [39]

Phytoplasma asteris OY-M [39]

Wolbachia pipientis wMel [39]

Mollicutes

Alphaproteobacteria

Unknown

Varies among Streptomyces species

Streptomyces avermitilis MA-4680 [39]

Thermus/ Deinococcus

Unknown

Varies among Bifidobacterium species

Bifidobacterium longum NCC2705 [39]

Varies among Anaplasmataceae

Varies among Acholeplasmataceae: Phytoplasma spp. are mixed or partly mixed, Acholeplasma spp. are skewed

Varies among Deinococcus species

Unknown

Unknown

At least 2 genome copies needed for reassembly of fragmented chromosomes [75]

Unknown

Actinobacteria

All Aquificales lack skew

Aquifex aeolicus VF5 [39]

Ploidy

Aquificales

Current IMG database

Mackiewicz et al. (2004) [39]

Group

Tab. 10.3: Bacteria with chromosomes lacking skew. N/A, not applicable

WP0001, WP0978

PAM001

Ga0133345_141

SAV4316

BL0640

aq_322

DnaA gene

WP0934

PAM002

Ga0133345_142

SAV4317

BL0638

aq_1422

DnaN gene

No

Yes

Yes

Yes

Yes, with intervening hypothetical protein

No

DnaA and DnaN adjacent?

intracellular pathogens

Phytoplasma spp. are plant pathogens, Acholeplasma are animal commensals or pathogens

Highly resistant bacteria (dessication, radiation, acid)

Soil; linear chromosome

Human gut

Aquifex is hyperthermophilic; family is mostly(?) thermophilic

Notes

292 | 10 Multiplication is vexation

Polyploid [79]

Unknown

Lack skew or somewhat skewed

Lacks skew (at least some other Flavobacterium spp. are skewed)

Blattabacterium spp.

Flavobacterium branchiophilum FL-15

CFB group

Polyploid – average ∼20 origins and 14 termini per cell in exponential phase [77]

Pseudomonas putida DOT-T1E [76]

Gammaproteobacteria

Other sources

This is the only P. putida not clearly skewed. Does have dnaA opposite peak of cumulative skew.

gvip205

FBFL15_0506

None found

PP_0010

S6july2012_01293 (or other, in different strains)

Varies among cyanobacterial species; see 󳶳 Fig. 10.14

DnaA gene

Synechocystis sp. PCC 6803

Most cyanobacterial genomes lack skew; see 󳶳 Fig. 10.14

Ploidy

alr2009

Gloeobacter violaceus PCC 7421

Cyanobacteria

Current IMG database

Nostoc sp. PCC 7120

Mackiewicz et al. (2004) [39]

Group

Tab. 10.3: Bacteria with chromosomes lacking skew. N/A, not applicable.

FBFL15_1235

dnaN upstream of pheT in all sequenced genomes

PP_0011

S6july2012_01293

alr2010

gvip446

DnaN gene

No

N/A

Yes

No; dnaA upstream of photosystem II reaction center genes

Yes

No

DnaA and DnaN adjacent?

fish pathogen; genome rearrangements compared to close relatives noted [80]

cockroach endosymbionts

toluene degrader, isolated from wastewater [78]

Notes

10.5 DNA replication |

293

294 | 10 Multiplication is vexation Clade Morphology Gloeobacter violaceus PCC 7421

G

Unicellular

E E A

Unicellular Unicellular ** Filamentous

Leptolyngbya sp. O-77 Oscillatoria acuminata PCC 6304 Oscillatoria nigro-viridis PCC 7112

C3 A A

Filamentous Filamentous Filamentous

Geitlerinema sp. PCC 7407 *

D

Filamentous *

B3 B3 B3

Filamentous Filamentous Unicellular †

B2 B2 B2 B2 B2 B2

Unicellular Unicellular Baeocystous Unicellular Unicellular † Unicellular †

B2 B2 B2 B2

Unicellular Unicellular Unicellular Baeocystous

B2 B2 B2

Unicellular Unicellular Unicellular

B1

Baeocystous

B1 B1 B1 B1 B1 B1 B1 B1 B1 B1

Unicellular Heterocystous Heterocystous Heterocystous Heterocystous Heterocystous Heterocystous Heterocystous Heterocystous Heterocystous

B1 B1 B1 B1 B1 B1

Heterocystous Heterocystous Heterocystous Heterocystous Heterocystous Heterocystous

B1 C2 C2

Heterocystous Unicellular * Unicellular *

A C1

Filamentous Unicellular

Thermosynechococcus elongatus BP-1 Acaryochloris marina MBIC11017 ** Arthrospira platensis NIES-39

Crinalium epipsammum PCC 9333 Microcoleus sp. PCC 7113 Chamaesiphon minutus PCC 6605† Halothece sp. PCC 7418 Dactylococcopsis salina PCC 8305 Stanieria cyanosphaera PCC 7437 Synechocystis sp. PCC 6803 Cyanobacterium stanieri PCC 7202† Cyanobacterium aponinum PCC 10605† Geminocystis sp. NIES-3708 Microcystis aeruginosa NIES-843 Cyanothece sp. PCC 7822 Pleurocapsa sp. PCC 7327 Cyanothece sp. PCC 8801 Cyanothece sp. PCC 8802 Cyanothece sp. BH68, ATCC 51142 Chroococcidiopsis thermalis PCC 7203 Gloeocapsa sp. PCC 7428 Calothrix sp. 336/3 Calothrix sp. PCC 6303 Anabaena sp. 90 Anabaena cylindrica PCC 7122 Nostoc azollae 0708 (”Trichormus”) Nostoc sp. NIES-3756 Nostoc sp. PCC 7120 Nostoc sp. PCC 7524 Anabaena variabilis ATCC 29413 Nodularia spumigena CCY9414 Calothrix sp. PCC 7507 Nostoc punctiforme PCC 73102 Cylindrospermum stagnale PCC 7417 Nostoc piscinale CENA21 Nostoc sp. PCC 7107 Rivularia sp. PCC 7116 Synechococcus elongatus PCC 6301 * Synechococcus elongatus PCC 7942 * Trichodesmium erythraeum IMS101 Cyanobium gracile PCC 6307 Synechococcus sp. WH7803 *

C1

Unicellular *

Prochlorococcus sp. MIT0801* Prochlorococcus sp. NATL1A *

C1 C1

Unicellular * Unicellular *

P. marinus marinus CCMP1375 * P. marinus pastoris CCMP 1986 *

C1 C1

Unicellular * Unicellular *

DnaA required (Ohbayashi et al. 2016) DnaA not required (Ohbayashi et al. 2016) Origin not identifiable by skew or DnaA (Nakamura et al., 2002, Thermosynechococcus; Welsh et al., 2008, Cyanothece)

* ** †

Skewed Segmented (several major changes of skew) DnaA not found

Fig. 10.14: RpoC1 based inferred phylogeny and GC skew of completed cyanobacterial genomes. The tree was adapted from Bao et al. [74] to show only species for which complete genomes are available in IMG. All species for which no skew diagram is displayed lack visible skew

10.5 DNA replication |

GC skew diagrams for selected species Visual classification

Average segment length

Skew shifts per 1 Gb

Gloeobacter violaceus PCC 7421

14007

289

Thermosynechococcus elongatus BP-1

13101

417

16450

167

16604

171

13522

310

Oscillatoria acuminata PCC 6304

14953

221

Oscillatoria nigro-viridis PCC 7112

15702

187

14937

236

14362

230

12987

396

Cyanothece sp. BH68, ATCC 51142

13771

326

Nostoc azollae 0708 (”Trichormus”)

14029

265

14167

239

Nostoc piscinale CENA21

13903

259

Nostoc sp. PCC 7107

13934

267

Rivularia sp. PCC 7116

15396

201

Synechococcus elongatus PCC 6301

Skewed

13565

374

Skewed

13553

369

16914

159

15064

219

Skewed

21768

88

Skewed

16902

275

Skewed

17432

248

P. marinus marinus CCMP1375

Skewed

17329

236

P. marinus pastoris CCMP 1986

Skewed

17141

230

Acaryochloris marina MBIC11017

Segmented

Arthrospira platensis NIES-39 Leptolyngbya sp. O-77

Geitlerinema sp. PCC 7407

Skewed

Stanieria cyanosphaera PCC 7437 Synechocystis sp. PCC 6803

Cylindrospermum stagnale PCC 7417

Synechococcus elongatus PCC 7942

Trichodesmium erythraeum IMS101

Cyanobium gracile PCC 6307 Synechococcus sp. WH7803 Prochlorococcus sp. MIT0801 Prochlorococcus sp. NATL1A

DnaA required (Ohbayashi et al. 2016) DnaA not required (Ohbayashi et al. 2016) Origin not identifiable by skew or DnaA (Nakamura et al., 2002, Thermosynechococcus; Welsh et al., 2008, Cyanothece)

1.0 0

GC skew

-1.0

Fig. 10.14: RpoC1 based inferred phylogeny and GC skew of completed cyanobacterial genomes. The tree was adapted from Bao et al. [74] to show only species for which complete genomes are available in IMG. All species for which no skew diagram is displayed lack visible skew

295

296 | 10 Multiplication is vexation

GC skew. These endosymbionts include Proteobacteria (Alpha, Beta, Gamma, and Delta), Flavobacteria, and Actinobacteria. The sample is too small to distinguish phylogenetic trends, and there can be variation within even closely related groups. For example, of eight Wolbachia (Alphaproteobacteria) genomes, three (IMG genome IDs 2597490173, 637000340, and 2524023201) appear strongly skewed, one (IMG genome ID 2540341172) weakly skewed, and four (IMG genome IDs 642555168, 637000339, 2540341171, 2588253760) mixed. Such varied clusters could yield clues to mechanisms governing skew, although endosymbionts would likely not be the most tractable group to work with.

10.5.3 Parallels to Cyanobacteria While in most bacterial lineages GC skew seems to be the rule, among finished cyanobacterial genomes it is the exception, visually evident only in the unicellular marine Prochlorococcus and Synechococcus spp., the filamentous Geitlerinema sp. PCC 7407, and to a lesser extent the unicellular Cyanobium gracile PCC 6307 and Acaryochloris marina MBIC11017 (󳶳 Fig. 10.14). Even in these species it is considerably weaker than in many other bacteria (e.g. the non-Beggiatoaceae species in 󳶳 Fig. 10.13c). Average segment length and the frequency of shifts in skew show no clear patterns, except that Synechococcus sp. WH7803 is the most strongly skewed by both appearance and numbers. It was noted over a decade ago that neither patterns of GC skew nor putative DnaA binding sites could be used to locate possible replication origins in Thermosynechococcus elongatus BP-1 [46] or Cyanothece sp. ATCC 51142 [47] as they can in other bacteria. Several complete cyanobacterial genomes lack a recognizable dnaA (Chamaesiphon minutus PCC 6605, Cyanobacterium stanieri PCC 7202, and C. aponinum PCC 10605). Even when present, DnaA may not always be required for DNA replication. Richter et al. [48] found that a dnaA mutant of the Synechocystis sp. PCC 6803 exhibited wild-type growth; an apparent DnaA gene was present (38% identical to E. coli), but DnaA boxes were not found. More recently, Ohbayashi et al. [44] showed that Synechocystis sp. PCC 6803 and Nostoc azollae dnaA mutants have no apparent growth defect or compensatory mutations, whereas Synechococcus elongatus PCC 7942 dnaA mutants gave rise to revertants in which a plasmid (and its replication machinery) were integrated into the chromosome. This strain also has the most skewed chromosome of the three (󳶳 Fig. 10.14), suggesting a possible connection between skew and DnaA-dependent replication, however work with additional species would be needed to confirm this idea. There is some correlation of skew with low genome copy number in the Cyanobacteria (󳶳 Tab. 10.4), but this has only been measured for a few species. As genome copy number can vary with strain and growth conditions [49–51] as well as growth stage [52, 53], and the same methods were not used in all cases, these comparisons

10.5 DNA replication

|

297

dnaA 3700001

1

100001

3600001

3700001 200001

3500001

38000011

100001 200001

3600001 300001

300001

3500001 3400001

400001

400001 3400001

3300001

500001

500001 3300001

3200001

600001

600001 3200001

3100001

700001

700001 3100001 3000001 800001

rRNA

800001 3000001

2900001

900001

900001 2900001

2800001

1000001

1000001 2800001

2700001

rRNA

1100001

1100001 2700001

2600001

1200001

1200001 2600001

2500001

1300001

1300001

2400001

1400001

2300001

1500001 2200001

1600001 2100001

1700001 2000001

1900001

1800001

2500001

dnaA

1400001 2400001 1500001 2300001 1600001

2200001 1700001

2100001 2000001

rRNA

Bdellovibrio bacteriovorus HD100

1900001

1800001

Bdellovibrio bacteriovorus 109J

Rings, from outside in: -- Genes on forward strand (colored by COG categories) -- Genes on reverse strand (colored by COG categories) -- RNA genes (tRNAs green, rRNAs red, other RNAs black) -- GC content, black -- GC skew, purple and green Fig. 10.15: Example of a skewed and a ‘segmented’ genome. Maps are from IMG/ER [65]

may not be reliable. However, it seems reasonable that a species with high genome copy number might undergo more frequent genomic recombination and rearrangement, which could remove the selection for skew by enhancing repair of mutations and/or erase it by rearranging DNA segments of differing GC skew. Even with selection for skew, the genome would likely take some time to reequilibrate after such a rearrangement; this could account for ‘patchwork’ genomes such as that of Bdellovibrio 109J (four segments) compared to Bdellovibrio HD100 (two segments) (󳶳 Fig. 10.15). Fluorescence microscopy of the spherical Synechocystis sp. PCC 6803 [54] revealed several chromosomal copies per cell; neither they nor the thylakoids seem to separate until just before final septation, and chromosome copies are not necessarily distributed evenly among daughter cells. It was suggested these and some other Cyanobacteria may not have a mechanism for strict chromosome partitioning. In the rodshaped Synechococcus elongatus PCC 7942, by contrast, when multiple chromosome copies are present they are generally spaced evenly along the long axis of the cell, alternating with carboxysomes, and partitioned evenly between daughter cells upon septation [55]. The position of the DNA replication origin has been investigated in this species [56]. It was stated that no clear boundary in GC skew could be identified, although our analysis suggests otherwise (󳶳 Fig. 10.14), where a cluster of possible DnaA boxes was identified upstream of dnaN. This region was very slightly overrepresented in next generation sequencing of BrdU-labeled DNA from a log-phase culture. The origin neighborhood is expected to be present in more copies per cell on average than the terminus region because cells typically divide shortly after the terminus is replicated; the weakness of this signal in S. elongatus PCC 7942 could result from asynchronous

syc0449_d

Skewed

Skewed

Synechococcus elongatus PCC 6301

Synechococcus sp. WH7803

Synpcc7942_1100

N/A

Anacy_0001

Mixed

Mixed

N/A

Mixed

N/A

N/A

Mixed

Synechococcus elongatus PCC 7942

Anabaena variabilis ATCC 29413 (Kutz, UTEX 1444)

Anabaena (Nostoc) sp. PCC 7120

Chrysosporum (Aphanizomenon) ovalisporum ILC-164

Synechocystis PCC 6803 (“Kasuza strain”)

Synechocystis PCC 6803, “GT strain” or “Vermaas strain”

Synechocystis PCC 6803, “motile strain” or “Moscow strain”

Anabaena cylindrica Lemm. (=PCC 7122)

N/A

S6july2012_01293

N/A

alr2009

Ava_0001

N/A

N/A

Skewed with mixing

Microcystis sp. HUB524

SynWH7803_0702

N/A

N/A

Prochlorococcus (natural Pacific Ocean populations)

DnaA

GC skew, main chromosome

Strain

Anacy_0016; Anacy_1219 (split by transposase)

N/A

N/A

S6july2012_00310

N/A

alr2010

Ava_0002

Synpcc7942_0001 (short upstream TMH)

N/A

SynWH7803_0001

syc1496_c

N/A

Chromosomal DnaN

No (0001 and 0016 close but on opposite strands)

N/A

N/A

No

N/A

Yes

Yes

No

N/A

No

No

N/A

DnaN and DnaA adjacent?

FtsQ-FtsZ

N/A

N/A

RsmH–FtsQ-FtsZ

N/A

FtsQ-FtsZ

FtsQ-FtsZ

MlbA-Ddl-FtsQ-FtsZ

N/A

MlbA-Ddl-FtsQ-FtsZ

MlbA-Ddl-FtsQ-FtsZ

N/A

FtsZ region predicted genes

Heterocystous

Unicellular

∼ 60 [49], ∼ 15 [89]

25 [90]

Unicellular

∼ 42 [49], ∼ 21 [89]

Heterocystous

∼ 8 in vegetative cells; much higher in akinetes [52]

Unicellular

Heterocystous [88]

∼ 8 [87]

12 [50], 10 [89]

Heterocystous

Unicellular

5–8 [86]

Unicellular ∼ 4 [49]

Unicellular

Unicellular

Unicellular

Morphology [81]

1–10 [85]

2–4 [84]

1–6 [83]

1 [82]

Genome copies per cell

Tab. 10.4: GC skew and DnaN genes in finished cyanobacterial genomes with genome copy number data. N/A, not applicable (genomes not sequenced)

298 | 10 Multiplication is vexation

10.5 DNA replication

|

299

replication either among the chromosomal copies in each cell, or across the population. It was proposed that this might be related to the absence of GC skew, but a mechanism was not suggested.

10.5.4 Alternatives to DnaA Several alternative DNA replication mechanisms are known in bacteria. The plasmidtype can sometimes serve as origins of replication in chromosomes (in the studied cases always secondary chromosomes), with the addition of some mechanism of tying their replication to the cell cycle. For example, the human pathogen Burkholderia cenocepacia J2315 has three chromosomes, one with a bacterial-type replication origin and two with plasmid-like RepA origins, including putative RepA and partition (Par) genes and RepA binding sites. [57] Like oriC of main chromosomes, these are located at points of inflection in GC skew. It appears that replication of the chromosomes is initiated sequentially, allowing each system to partition its daughter chromosomes to the poles properly, but the mechanism has not yet been identified. As a second example, chromosome 1 of Vibrio cholerae N16961 has a bacterial-type oriC, while ori2 on chromosome 2 is controlled by DNA-binding protein RctB. Cell-cycle control of chromosome 2 replication requires binding of RctB to a site on chromosome 1, but details of this interaction have not yet been worked out [58]. No obvious matches to RctB could be found in the LSB genomes (not shown). Finally, as mentioned above (Section 10.5.3), Synechococcus elongatus PCC 7942 dnaA mutants that integrated a plasmid into their chromosome were able to continue growth; although this was a laboratory situation, it does suggest that mutational changes of origin are possible. Another alternative, inducible stable DNA replication (iSDR) is known from cells induced for the SOS response [59] [60]. It does not require DnaA, transcription, or translation, and can continue for several hours. Initiation of iSDR uses a DNA primer and requires the recombinase activity of RecA and the helicase activity of RecBC or RecBCD, both of which are essential for the RecBCD pathway of homologous recombination. In E. coli, the iSDR origins oriM1A and oriM1B are in the oriC region, while oriM2 is near terC. However, we are not aware of any reports of bacteria growing for extended periods using this system. Candidate genes for all of the Rec proteins can be found in all of the LSB genomes (not shown). A third mechanism, constitutive stable DNA replication (cSDR), uses an RNA primer and is activated in E. coli RNase HI mutants [61]; RNaseH normally degrades the DNA in RNA-DNA hybrids, so this would not be expected to work when RNaseH is functional. Whatever the replication initiation mechanism, selection for GC skew should still be in effect, unless there are several (or perhaps shifting) origins per chromosome. As far as we are aware, DNA polymerization is always unidirectional, so one strand will always be more subject to C deamination. As mentioned above (Section 10.5.3),

300 | 10 Multiplication is vexation

frequent genome rearrangement and/or recombination could erase this signal. Alternatively, these organisms could use some especially effective form of DNA repair.

10.6 Summary and perspectives The LSB genomes available to date suggest this group differs in several respects from related Gammaproteobacteria: their oriC regions have a reduced complement of genes, and some are apparently missing DnaA; the two for which complete genomes are available lack GC skew; some genes commonly clustered in other species are dispersed, particularly in Thiomargarita; and the Thiomargarita genomes lack much of the expected cell division machinery, in particular the known genes for septal (as opposed to lateral) peptidoglycan biosynthesis. More genomic sequences and considerably more observational and experimental evidence will be needed to understand the evolution of this morphologically diverse group. The evidence to date suggests a possible set of events, although their order is not entirely clear and there were likely other accompanying changes. The DnaAtype origins of replication found in B. alba and B. leptomitiformis apparently were replaced by an alternate form, perhaps acquired from a plasmid or phage. Relaxed copy-number control via alternate origins may have favored multiple genome copies, which in turn may better support large cell size, both by increasing and decentralizing gene expression and by decreasing the need for precise chromosome partitioning. Multiple genome copies could also contribute to loss of GC skew, with more frequent gene rearrangements and/or higher rates of recombinational repair. At some stage of this evolution, vacuoles developed in the marine forms (perhaps by invagination of the inner membrane), allowing storage of nitrate and perhaps other compounds. This can be an advantage in environments with shifting oxic/sulfidic boundaries, but is an apparent challenge for chromosome partitioning and cell division. Changes in the septation mechanism may have been required to allow very large cells to form and divide. Potential additional advantages of large size and multiple genome copies are increased surface area for epibionts, which might temporarily or permanently take over host cell functions, and possession of backup copies of genes, allowing some to be mutated or replaced while still maintaining essential cell functions [45, 62]. What observations and experiments might lead to better understanding of the development of the LSB? Additional genomic sequences will of course help with inference of a timeline; it will be particularly interesting to see at what stages horizontal gene transfer (or close interspecies relationships) may have supported evolution of new cell morphologies. Most LSB have resisted cultivation to date, but can sometimes be kept alive in the laboratory, so microscopy will likely be the primary experimental approach, at least in the near future. For example, mechanisms of cell growth and division can be investigated by following incorporation of labeled peptidoglycan precursors [63] to identify sites of lateral and septal cell wall growth, or by antibody label-

References |

301

ing of FtsZ and MreB to identify which (if either) forms a cell division ring in different LSB. Questions of genome copy number, the synchrony (or otherwise) of chromosome replication, and the location of replication origins in species lacking DnaA might be investigated by looking for foci of incorporation of nucleic acid analogues (bromodeoxyuridine or ethynyl deoxyuridine), either by fluorescence microscopy or by DNA capture and sequencing. Acknowledgment: The authors thank William Margolin for suggestions and reagents, and Han Bao and Cheryl Kerfeld for sharing an updated cyanobacterial phylogeny in advance of publication. We thank the University Imaging Centers and Guillermo Marqués at the University of Minnesota for imaging support. The UNC Microbiome and Marine Sciences Lunch Bunch groups are thanked for helpful questions, and Alecia Septer and Lisa Nigro for comments on the manuscript. The Guaymas Basin project was funded by NSF OCE 0647633 and 1357238.

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Weishu Zhao and Xiao Xiang

11 Life in multiextreme environments: cross-stress response in Thermococcales 11.1 Abstract Thermococcales are commonly the dominant microbes in black chimney environments of deep sea hydrothermal vents, which are one of the most extreme environments on the Earth and are similar to the early Earth. Thermococcales have adapted to the dramatic environmental fluctuations of physical and chemical factors in hydrothermal vents. They usually have wide growth ranges but small genomes. Studies of Thermococcales adaptation to various extreme environments find that some of special metabolic pathways they possess are related to adaptation to multiextreme environments. These metabolic pathways include: compatible solutes, energetic metabolisms, membrane lipid components, amino acid metabolism, and antioxidation pathways. Exploring these pathways further reveals that Thermococcales may have common adaptation strategies to multiextreme environments. Studying common adaptation mechanisms of Thermococcales will expose the survival strategies of microorganisms in low energy, high temperature extreme environments such as the deep biosphere or extraterrestrial domains. It will also provide precious research models and ideas for investigating metabolic traits of early life and further enhance our understanding of the origin of life. Last but not least, it will also offer theoretical references and biological materials to synthetic biology research and industrial applications.

11.2 Introduction and background A deep sea hydrothermal vent system is a kind of ecosystem dependent on chemical energy rather than solar energy. It is also one of the most extreme environments and is an ideal choice to study the enzymes, metabolic pathways, physiological traits and adaptations to multiextreme environments of various extreme microbes [1]. The system is considered to be similar to the early Earth, therefore it is regarded as an accessible way to explore early life and also the origin of life [1]. Thus, since deep sea hydrothermal vents were first discovered by scientists of the Scripps Institution of Oceanography near Galápagos Rift in 1977 [2], they have become hotspots for research. Oxidative, alkaline and cold seawater cools down and reacts with reduced, acid and hot vent fluid, which contributes to the hydrothermal vent environment. The formation of a hydrothermal vent results in enormous physical and chemical fluctuations https://doi.org/10.1515/9783110493672-011

308 | 11 Life in multiextreme environments: cross-stress response in Thermococcales

and steep gradients of physical and chemical factors nearby, including temperature, pH, salinity and redox level [1]. The physical and chemical gradients provide essential nutrients and energy for chemotrophic microorganisms, while also posing a challenge for the organisms living in the deep sea hydrothermal vent environment. The primary problem for the survival of organisms in hydrothermal vents is how to cope with the extreme and dramatically fluctuant physical and chemical environments. In deep sea hydrothermal vents, chemotrophic microorganisms are the primary producers, which support the entire deep sea hydrothermal vent system [3, 4]. Microorganisms in the system usually have extreme traits such as being thermophilic (thermotolerant), piezophilic (piezotolerant), acidophilic (acidolerant), alkaliphilic (alkalitolerant) and so on. The hyperthermophilic archea Thermococcales are the most representative microorganisms among them, and are near the root of the phylogenetic tree. Many environmental metagenomic and transcriptomic studies have shown that Thermococcales are the predominant microorganisms in eutrophic aquatic geothermal environments [5, 6]. They are a ubiquitous in deep sea hydrothermal vent systems as representatives of hyperthermophilic microorganisms and are dominant in hydrothermal chimneys, plumes and hydrothermal vent sediments. Moreover, environmental metatranscriptome analysis has shown that Thermococcales are the exclusive active microbe community at the depths of 1,600 m below the sea floor [7]. Thus, studying how Thermococcales deal with dramatically fluctuant physical and chemical environments can help reveal the strategies of microbes for adapting to multiextreme environments and also the developing processes of early organisms. Thermococcales are also the sources for the development of various extreme enzymes.

11.3 Basic characteristics of Thermococcales As typical microorganisms in hydrothermal vents, Thermococcales exist only in geothermal aquatic environmenst [8]. Three genera belong to Thermococcales: Thermococcus, Pyrococcus and Palaeoococcus. They have extremely high growth temperatures and wide growth ranges, especially the Thermococcus genus. Their genome sizes are small; all genomes of Thermococcales are smaller than 2.3 Mb as far as we know. Thermococcales have relatively simple metabolism pathways due to small genomes, which make them the ideal case study for research on adaptation to multiextreme environments in hydrothermal vents.

11.3.1 Physiological traits In 1983, Zillig isolated and identified the first Thermococcales strain, Thermococcus celer [9], on a shallow volcano in Italy. According to the List of Prokaryotic names with standing in Nomenclature (LPSN), isolated and identified Thermococcales so far

11.3 Basic characteristics of Thermococcales | 309

include 31 Thermococcus species, six Pyrococcus species and only three Palaeococcus strains. All Thermococcales were isolated from hyperthermal geothermal environments. Thus, they all belong to hyperthermophiles and their optimum growth temperature is above 75 °C. Up to now, most of the Thermococcales isolates have been found to live independently while only two, T. cleftensis and T. paralvinellae [10], establish symbioses with macroorganisms, which are Paralvinella sp. oolychaetes from deep sea hydrothermal vents [11]. The list of isolated Thermococcales is shown in 󳶳 Tab. 11.1. 11.3.1.1 Temperature range The major differences between the three genera of Thermococcales are the range of growth temperature and optimum growth temperature. Thermococcus have the widest range of growth temperatures, which are generally from 55 °C to 95 °C; the ranges usually cover more than 40 °C and are far larger than the common ranges of other microorganisms. The optimum growth temperatures of Thermococcus are from 75–88 °C and the highest temperatures are around 100 °C. Among all Thermococcus strains, T. eurythermalis A501 [12] is the most significant one in terms of growth temperature range. The growth temperature range at atmospheric pressure of T. eurythermalis A501 is 50– 100 °C, which covers 50 °C. So far, it has the widest growth temperature range among all known Thermococcales strains. Growth temperature ranges of Palaeococcus are smaller than those of Thermococcus and their highest growth temperatures are lower than 90 °C. While the optimum growth temperatures of known Palaeococcus are 80– 90 °C, fewer of them have been isolated. Compared with these two genera, Pyrococcus have higher optimum growth temperatures and the highest growth temperatures. The optimum growth temperatures are 95–100 °C and the highest growth temperatures are above 100 °C. Correspondingly, the lowest growth temperatures of Pyrococcus are relatively high; usually around 70 °C [8]. 11.3.1.2 pH range For Thermococcales strains, differences in pH ranges are quite large. However, most of the strains are neutrophiles; the optimum pH ranges are 6–8. Only two strains are alkalophiles; they are T. alcaliphilus [13] and T. acidaminovorans [14]. The optimum growth pH of both is 9. Many Thermococcales strains have relatively wide pH growth ranges. For example, the growth pH range of P. glycovorans is 2.5–9.5, which covers 7 pH units. This is unusual for both acidophiles and alkalophiles. P. glycovorans has the widest pH growth range among all Thermococcales strains. 11.3.1.3 Pressure range The response of Thermococcales strains to high hydrostatic pressures is also a hot research topic. Many Thermococcales strains were isolated from deep sea hydrothermal vents. Their original living habitat is at the depths of 2,000 m and the pressure of the

Environment location (Deep)

Italy (0 m) Kraternaya cove (0 m) Middle Okinawa Trough (0 m) Mexican west coast (2,600 m) North Fiji Basin (2,000 m) Izu-Bonin and South Mariana Trough areas, Western Pacific ocean Vulcano Island, Italy (0 m) 21°N, East Pacific Rise Guaymas Basin (2,000 m) Guaymas Basin (2,000 m) Submarine hot vents, New Zealand (0 m) Submarine hot vents, New Zealand (0 m) Juan de Fuca Ridge (2,200 m) Snakepit, Mid-Atlantic Ridge (3,550 m) New Zealand (0 m) Palaeochori Bay, Milos, Greece (9.4 m) Mid-Okinawa Trough (1,394 m) Italy (0 m) Lucrino Naples a– Porto di Levante (0 m) Samotlor oil reservoir (1,799–2,287 m) Lake Taupo area of North Island (0 m) Guaymas Basin (2,000 m)

Year

1983 1990 1995 1996 1996 1996

1997 1997 1998 1998 1998 1998 1998 1999 1999 2000 2000 2001 2001 2001 2001 2003

Species

T. celer T. stetteri T. profundus T. chitonophagus T. fumicolans T. peptonophilus

T. alcaliphilus T. hydrothermalis T. aggregans T. guaymasensis T. gorgonarius T. pacificus T. barossii T. barophilus T. zilligii T. aegaeus T. siculi T. acidaminovorans T. litoralis T. sibiricus T. waiotapuensis T. gammatolerans

SMHV DSHV DSHV DSHV SMHV SMHV DSHV DSHV FHS SMHV DSHV SMHV SMHV HTOR FHS DSHV

MSF MSF DSHV DSHV DSHV DSHV

Type

56 55 60 56 68 70 60 48 55 70 50 65 55 40 60 55

85 85 88 88 80–88 80–88 82.5 85 75–80 85 85 85 85 78 85 88

90 100 94 90 95 95 92 95 85 90 93 93 98 88 90 95

93 94 90 93 103 100

H

6.5 3.5 5.6 5.6 5.8 6.0 4.0 4.5 5.4 4.5 5.0 5.0 4.0 5.8 5.0 4.0

– 5.7 4.5 3.5 4.5 4.0

L

88 75 80 85 85 85–90

O

L – 55 50 60 70 60

pH

Temperature (°) O

9.0 6.0 7.0 7.2 6.5–7.2 6.5 6.5–7.5 7.0 7.4 6.0 7.0 9.0 6.0 7.3 7.0 6.0

5.8 6.5 7.5 6.7 8.0 6.0

H

10.5 9.5 7.9 8.1 8.5 8.0 9.0 9.5 9.2 7.5 9.0 9.5 8.0 9.0 8.0 8.5

– 7.2 8.5 9.0 9.5 8.0 1.0 2.0 – – 1.0 1.0 1.0 1.0 0.0 2.0 1.0 1.0 1.8 0.5 0.0 1.0

– – 1.0 0.8 0.6 1.0

L

2.0–3.0 0.3 4.0 2.0 2.0–3.0 – 1.8–2.0 0.5 2.0

2.0–3.0 3.0–4.0 1.8 1.8 2–3.5 2–3.5

3.8 2.5 2.0 2.0 1.3–2.6 3.0

O

NaCl (%) H

6.0 8.0 – – 6.0 6.0 4.0 4.0 1.2 8.0 4.0 6.0 6.5 7.0 1.4 5.0

– – 6.0 8.0 4.0 5.0

– – – – – – – 0.1 – – – – – – – –

– – – – – 30

L

– – – – – – – 40 – – – – – – – –

– – – – – 45

O

Pressure (MPa)

– – – – – – – 80 – – – – – – – –

– – – – – 60

H

Tab. 11.1: List of characterized Thermococcales isolates. “T.”: Thermococcus; “Pyr”: Pyrococcus; “Pal”: Palaeococcus; MSF, Marine solfataric fields; DSHV, Deep-sea hydrothermal vent; SMHV, Submarine hot vents; HTOR, High-temperature oil reservoir; FHS, Freshwater hot spring; THS, Terrestrial hot spring; GW, Geothermal well. L: lowest boundary of growth condition; O: optimal growth condition; H: highest boundary of growth condition. “–”, not reported

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Environment location (Deep)

Mid-Atlantic Ridge Suiyo Seamount, Izu-Bonin Arc (1,380 m) Kodakara island, Japan (0 m) Suiyo Seamount, Izu-Bonin Arc (1,380 m) Rainbow site, the Mid-Atlantic Ridge (2,300 m) Sarah Spring area, East Pacific Rise (2,700 m) Paralvinella sp. polychaete worms (2,350 m) Paralvinella sp. polychaete worms (2,200 m) East Pacific Rise (2,633 m) Porto di Levante (0 m) – East Pacific Rise (2,650 m) North East Pacific Ocean (1,395 m) Ashadze, Mid Atlantic Ridge (4,100 m) Ogasawara Trough (1,338 m) Southern Tyrrhenian Sea (0 m) East Pacific Ocean (2,737 m)

Year

2004 2005 2005 2007 2007 2013 2014 2014 2014 1986 1988 1999 1999 2011 2000 2006 2013

Species

T. atlanticus T. coalescens T. kodakarensis T. celericrescens T. thioreducens T. prieurii T. cleftensis T. paralvinellae T. nautili Pyr. furiosus Pyr. woesei Pyr. glycovorans Pyr. horikoshii Pyr. yayanosii Pal. ferrophilus Pal. helgesonii Pal. pacificus

DSHV DSHV THS DSHV DSHV DSHV DSHV DSHV DSHV MSF – DSHV DSHV DSHV DSHV GW DSHV

Type H 95 90 100 85 94 95 94 91 95 105 105 104 102 108 88 85 90

5.0 5.2 5.0 5.6 5.0 4.0 – 5.0 4.0 5.0 – 2.5 5.0 6.0 4.0 5.0 5.0

L

85 87 85 80 83–85 80 88 82 87.5 100 100 95 98 98 83 80 80

O

L 70 57 60 50 55 60 55 50 55 65 – 75 80 80 60 45 50

pH

Temperature (°) O 6.0 6.5 6.5 7.0 7.0 7.0 – 8.0 7.0 7.0 6.5 7.5 7.0 7.5–8.0 6.0 6.5 7.0

H 8.0 8.7 9.0 8.3 8.5 8.0 – 8.3 9.0 9.0 – 9.5 8.0 9.5 8.0 8.0 8.0

1.5 1.5 1.0 1.0 1.0 1.0 – 1.0 1.0 0.5 – 1.7 1.0 2.5 2.0 0.5 1.0

L

2.3 2.5 3.0 3.0 3.0 2.0 – 3.2 2.0 2.0 3.0 2.6 2.4 3.5 4.7 2.8 3.0

O

NaCl (%) H 4.6 4.5 5.0 4.5 5.0 5.0 – 6.5 4.0 5.0 – 5.2 5.0 5.5 7.3 6.0 4.0

– – – – – – – – – 0.1 – 0.1 0.1 20.0 0.1 – 0.1

L

– – – – – – – – – 10 – 10 0.1 52 30 – 30

O

Pressure (MPa)

– – – – – – – – – 30 – 45 40 120 60 – 80

H

Tab. 11.1: List of characterized Thermococcales isolates. “T.”: Thermococcus; “Pyr”: Pyrococcus; “Pal”: Palaeococcus; MSF, Marine solfataric fields; DSHV, Deep-sea hydrothermal vent; SMHV, Submarine hot vents; HTOR, High-temperature oil reservoir; FHS, Freshwater hot spring; THS, Terrestrial hot spring; GW, Geothermal well. L: lowest boundary of growth condition; O: optimal growth condition; H: highest boundary of growth condition. “–”, not reported

11.3 Basic characteristics of Thermococcales | 311

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habitat is mostly above 20 MPa. Thus, most Thermococcales strains are model strains for studying pressure adaptation in piezophilic hyperthermophiles. Many studies focus on the pressure adaptation of Pyrococcus. P. furiosus [15], P. glycovorans [16] and P. yayanosii [17] are all piezophiles. P. yayanosii CH1is the first and only reported obligate piezophilic hyperthermophilic archaeon [17], isolated from depths of 4,100 m at the Mid-Atlantic Ridge. The growth pressure range of P. yayanosii CH1 is 15–130 MPa and the optimum growth pressure is 52 MPa [17, 18]. P. yayanosii CH1 is a precious material for studying the piezophilic mechanism of hyperthermophilic archaea. Compared with Pyrococcus, fewer research efforts focus on pressure adaptations of the genus Thermococcus; so fat only T. barophilus is reported to be piezophilic [19]. The optimum pressure is 40 MPa and the growth pressure range is 0.1–80 MPa. T. eurythermlais is found to be conditional piezophilic archaeon; the optimum pressure is 0.1–30 MPa at the optimum growth temperature of 85 °C, while the optimum pressure is 10–20 MPa at the higher temperature of 95 °C [12]. Research on the relationship between temperature and pressure influence on T. peptonophilus indicated that the optimum growth temperature of T. peptonophilus became 85 °C at 30 MPa and the optimum temperature switched to 90–95 °C at 45 MPa [20]. Aside from T. barophilus, T. eurythermlais and T. peptonophilus, there are no other studies on pressure responses of Thermococcus. Among three newly isolated species in Palaeococcus, both P. ferrophilus [21] and P. pacificus [22] are piezophiles; optimum growth pressures for both are 30 MPa and growth pressure ranges are 0.1–60 MPa and 0.1–80 MPa respectively. 11.3.1.4 Salinity range In general, Thermococcales have narrow growth salinity ranges because most of the strains are isolated from marine environments and salinity does not change dramatically around hydrothermal vents [1]. Thus, common optimum salinity ranges of Thermococcales are 2–4%, which is close to the average salinity of seawater (3%). The optimum salinity ranges of two strains, T. zilligii [11] and T. waiotapuensis [23], which were isolated from freshwater, are 0.3% and 0.5% respectively. 11.3.1.5 General metabolic characteristics Isolated strains of Thermococcales so far have similar physiological and metabolic characteristics. They are all obligate anaerobic and organic chemoheterotrophic microorganisms [8]. Most of the strains require complex organics, such as yeast extract, peptone, tryptone and so on, and some can use starch, sucrose, acetate and so on as their sole carbon source [12]. These strains are all sulfur-reducing archaea and the growth of some depends on elemental sulfur. However, while most strains do not depend on elemental sulfur, it can stimulate their growth significantly. Thermococcales produce H2 S in the presence of element sulfur, while they produce H2 without elemental sulfur. As a novel clean energy, hydrogen has received great attention. Thermococ-

11.3 Basic characteristics of Thermococcales |

313

cales have a special metabolic trait that means they can use sugar, formic acid, carbon monoxide and so forth as simple organic substrates to produce H2 . Therefore, modifying Thermococcales strains by genetic tools to increase productivity of hydrogen and then using the hydrogen for energy production has become a hot topic in applied research [24–27]. Moreover, for Thermococcus and Pyrococcus, the development of hyperthermal industrial enzymes based on the hyperthermophilic traits of Thermococcales has been carried out for long time, but such research for Palaeococcus has just began.

11.3.2 Genome traits With the development of sequencing technology, genome analysis has provided more abundant information for studying metabolisms, functions and usages of microorganisms and greatly enhances our understanding of them. We have acquired whole genome sequences for quite a few Thermococcus strains. We got the complete sequence map of T. kodakarensis KOD1 in 2005. Subsequently, complete genome maps of 18 strains of Thermococcus were acquired based on records of NCBI. Fourteen of them belong to identified species while the other four have not been identified. For Pyrococcus, seven strains have complete genome maps; four of them have been identified but the other three have not. However, research into the genome of Palaeococcus is not sufficient; only two strains have complete genome maps. The list of genome traits of Thermococcales is shown in 󳶳 Tab. 11.2. Complete genomes of Thermococcales show that genome sizes of these hyperthermophilic archaea are between 1.7 to 2.3 Mb. For Thermococcus and Palaeococcus, the genome sizes of different strains are quite different. The genome sizes of Pyrococcus strains are usually smaller, at are around 1.7–1.9 Mb. For Thermococcales, the biggest known genome belongs to T. celericrescens DSM 17994 (2.34 Mb) [28] and the smallest genome is that of P. yayanosii CH1 (1.72 Mb) [29]. There are two type genomes regarding G+C content; one is around 40% and the other is above 50% [12, 30]. Coding genes make up a high proportion in all these genomes; sequence-coding proteins account for 95% of the whole genome sequence. Analysis of complete genome maps can be used to predict potential metabolic pathways and functions and thus provide clues to elucidate these pathways and functions. Metabolic pathways and functions of T. kodakarensis KOD1, T. gammatolerans EJ3 [30] and T. onnurineus NA1 [31] have been predicted based on whole genome sequences. For T. gammatolerans EJ3, results of proteome analysis further prove and amend the annotations of genome [32, 33]. For T. onnurineus NA1, differentiation analysis of the proteome with different C1 compounds was adopted to explain the potential mutative relationships among all main metabolic pathways and functions and the meaning of sulfur metabolism to the strain [34, 35].

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Tab. 11.2: List of current complete genomes of Thermococcales. “T.”: Thermococcus; “Pyr.”: Pyrococcus; “Pal.”: Palaeococcus Strain name

Genome size (bp)

G+C

Predicted proteins

Plasmid

T. barophilus MP T. gammatolerans EJ3 T. kodakarensis KOD1 T. litoralis DSM 5473 T. nautili 30–1 T. onnurineus NA1 T. sibiricus MM 739 T. peptonophilus JCM 9653 T. guaymasensis DSM 11113 T. zilligii AN1 T. thioreducens DSM 14981 T. celericrescens DSM 17994 T. cleftensis CL1 T. paralvinellae ES1 T. sp. 4557 T. sp. AM4 T. sp. 2319x1 T. sp. CDGS Pyr. furiosus DSM 3638 Pyr. horikoshii OT3 Pyr. abysii GE5 Pyr. yayanosii CH1 Pyr. sp. NA2 Pyr. sp. ST04 Pyr. sp. NCB100 Pal. ferrophilus DSM 13482 Pal. pacificus DY20341

2,064,237 2,045,438 2,088,737 2,309,423 1,976,356 1,847,607 1,845,800 1,847,000 1,920,914 1,764,559 2,052,483 2,337,139 1,950,313 1,957,742 2,011,320 2,086,428 1,961,221 1,928,800 1,908,256 1,738,505 1,765,118 1,716,818 1,861,320 1,736,890 1,977,130 2,206,430 1,859,370

0.42 0.54 0.52 0.43 0.55 0.51 0.40 0.52 0.53 0.55 0.54 0.54 0.56 0.40 0.56 0.55 0.45 0.51 0.41 0.42 0.45 0.52 0.43 0.42 0.45 0.54 0.43

2,115 2,111 2,231 2,292 2,127 1,936 1,912 1,917 2,021 1,785 2,083 2,332 1,785 2,054 2,076 2,225 2,042 1,867 1,989 1,802 1,846 1,795 1,920 1,793 1,997 2,235 1,945

1 0 0 0 3 0 0 1 0 0 0 0 0 0 0 0 0 0 0 0 1 0 0 0 0 0 0

11.4 Adaptation related pathways of Thermococcales Hyperthermophilic archaea have adapted to the dramatically fluctuant physical and chemical conditions in hydrothermal vents. Like other hyperthermophiles, Thermococcales have reverse gyrase, generating positive supercoiling in DNA [36], and chaperonins [37], such as heat shock proteins and thermosome to enhance stability of the molecular structure. Besides these, Thermococcales have some other special metabolic pathways, which help them to adapt to multiextreme environment stresses. These pathways include: special ATP synthase driven by a sodium ion and energy providing system, amino acid metabolism, synthesis and accumulation of various compatible solutes, special synthetic pathways, and composition of membrane lipids and anaerobic antioxidative pathways. Energy metabolisms and amino acid metabolisms are two of most important pathways among all of them. For energy metabolism, Ther-

11.4 Adaptation related pathways of Thermococcales |

315

mococcales can acquire ATP from diverse pathways and can synthesize ATP with extremely low energy quanta in hyperthermal environments [38], which gives them a great advantage in hydrothermal vent environments. Moreover, most Thermococcales strains can use amino acids as their sole carbon source. Hydrothermal vents contain abundant amino acids and polypeptides, which come from abiotic synthesis and degradation of macroorganisms [1]. Thermococcales reduce their energy consumption by using these amino acids and polypeptides directly, since synthesis of amino acids is an important part of energy consumption under low energy conditions [39]. Besides these, Thermococcales synthesize and accumulate diverse compatible solutes as cell protectors, adjust motility and stability of cell membranes by changing the composition of membrane lipids and get rid of extra electrons in the pathways through anaerobic antioxidation. These special metabolic pathways help Thermococcales to survive under the multiextreme environments of hydrothermal vent.

11.4.1 Energy conversion ATP synthases on the membrane are present in every domain, and all these ATP synthases have a common ancestor. For Thermococcales strains, energy production is carried out by A1 A0 ATP synthase driven by sodium ions, which is different from general ATP synthases driven by protons in bacteria and eukaryotes. This A1 A0 ATP synthase driven by sodium ions was first reported in P. furiosus [40]. Besides this special ATP synthase, obligate anaerobic strains of Thermococcales usually do not have complicated respiratory chains but a kind of membrane bound complex: membrane bound hydrogenase (MBH) and membrane bound oxidoreductase (MBX). They use Na+ /H+ antiporters of these membrane bound hydrogenase complexes to form Na+ gradients across membranes and thus drive A1 A0 ATP synthase. All strains of Thermococcus and Pyrococcus contain MBH and MBX in their genomes. Both T. sibiricus and T. barophilus have two copies of membrane bound the hydrogenase complex [41]. MBH and MBX are highly homologous [42], but they have a critical difference. MBH has [Ni-Fe] hydrogenase activity where the electron is transferred to H+ to produce H2 . However, MBX lacks two critical amino acids at the catalytic site of [Ni-Fe] hydrogenase and thus loses hydrogenase activity, so the electron is transferred to NADP+ and produces NADPH redox [43]. In culture with elemental sulfur, Thermococcales strains use sulfur as the final electron acceptor. MBX and NADPH sulfur oxidoreductase (NSR) are the main participators and essential parts of sulfur metabolism, and the electron is transferred to elemental sulfur to produce H2 S through NADPH redox, produced by MBX [44]. While in culture without elemental sulfur, Thermococcales strains use protons, through [Ni-Fe] hydrogenase in MBH, to transfer electrons to H+ and produce H2 . Both MBH and MBX can pump H+ or Na+ out of the cell during electronic transfer and help form the electrochemical gradient and produce ATP by ATP synthase.

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Besides simple electronic transport chains comprised of MBH and MBX, some Thermococcales strains, such as T. onnurineus, can also produce energy through the formate metabolic pathway [45]. This pathway is the simplest and most ancient reported energy conversion pathway. Although formate can react with water and produce carbonate and hydrogen in anaerobic environments, the standard Gibbs free energy of this reaction is very low (∆G ř = +1.3 kJ/mol). Therefore, it was considered impossible for the pathway to provide enough energy to support the growth of microorganisms. Additionally, before the discovery of Thermococcales, no individual strain was found to be able to produce carbonate and hydrogen using formate. Thermococcales can produce hydrogen through oxidation of formate. The actual Gibbs free energy of this process in the strain is supposed to be −8 to −20 kJ/mol [45].The completed formate oxidative gene cluster includes formate dehydrogenase, the membrane bound hydrogenase complex and Na+ /H+ antiporters. Formate produces H+ and carbon dioxide via the formate dehydrogenase complex. The membrane bound hydrogenase complex turns H+ to H2 , pumps H+ out and forms an extracellular proton gradient [24]. Next, cation/proton antiporters turn the proton gradient into a Na+ gradient. The Na+ gradient drives ATP synthase to synthesize ATP and provide energy for the cell [34]. Extracellular formate depends on formate transporters to enter the cell. Intracellular formic acid is supposed to be produced by pyruvate under catalysis of the predicted pyruvate-formate lyase (PFL), which could produce both formic acid and acetyl-CoA [46]. An homologous protein of PFL has been found only in Thermococcus, which may be related to production pathways for intracellular formic acid. However, the function of this protein needs further experimental proof. So far, every part of Thermococcales’ specific energy conversion system that has been revealed depends on sodium. Although the detailed mechanism of these complexes remains unclear, more and more proteome and transcriptome data have recently shown that energy conversion is related to adaptation to different environmental stresses. One point to highlight is that piezophilic research on the obligate piezophile P. yayanosii CH1 and the piezophile T. barophilus MP found that energy conversion was significantly upregulated under high hydrostatic pressure [47, 48].

11.4.2 Amino acid metabolism The biosynthesis pathways of amino acids are quite different in different Thermococcales strains, though other major metabolic pathways are highly conserved in Thermococcales. The biosynthesis of eight amino acids can be predicted based on genomes for all known Thermococcales strains, including alanine (Ala), asparate (Asp), asparagine (Asn), glutamate (Glu), glutamine (Gln), glycine (Gly), serine (Ser), and threonine (Thr). In some Thermococcales strains, the pathways of seven amino acids can be predicted: Cysteine (Cys), Methionine (Met), Tryptophan (Trp), Lysine (Lys), Arginine (Arg), Histidine (His) and Tyrosine (Tyr). Completed biosynthesis pathways of

11.4 Adaptation related pathways of Thermococcales |

317

the rest of the amino acids cannot be found in known genomes of Thermococcales. However, essential amino acid requirement tests of T. gammatolerans and P. abyssi found that whether a particular amino acid is essential is not directly correlated with whether its synthetic pathway can be predicted in the genome. The genome of P. abyss contains completed synthetic pathway of Thr, but Thr is still required as an essential amino acid in culture [49]. On the other hand, although biosynthesis pathways of Pro and Ile have not been predicted based on the genome of T. gammatolerans, there is no need to add these two amino acids into the medium as essential amino acids [33]. This data indicates that Thermococcales may have some special amino acid biosynthesis or transfer pathways, which need further investigation. The transcriptomic and proteomic studies of Thermococcales under different environmental stresses reveal that biosynthesis and utilization of amino acids play a key role in environmental adaptation in Thermococcales. In a study of T. kodakaraensis comparing protein expression differentiation under heat, oxidative stress and high salinity through 2-D gel based proteomic analysis discovered 59, 42 and 29 differentially expressed proteins under these three conditions respectively [50]. KEGG pathway enrichment revealed upregulation of amino acid biosynthesis as a response to all three stresses [50]. Additionally, amino acid biosynthesis was found to be related to pressure adaptation in the study of piezophiles in Thermococcales. A study on the obligate piezophile P. yayanosii CH1 reported significant differential responses of amino acid metabolism under both high and low hydrostatic pressure stresses [48]. For T. barophilus MP, biosynthesis of His, Asp, Thr and Pro is closely related to response to high and low hydrostatic pressure stresses [47]. Additionally, the test of essential amino acid utilization under different pressure conditions showed that three extra amino acids, Asn, Cys and Tyr, were required under high hydrostatic pressure in T. barophilus MP [51]. The relationship between essential amino acid utilization and responses to environmental stresses is still under investigation. It was speculated that energy-providing limitation under stress conditions may contribute to the relationship. Although the mechanisms remain unclear, it is reasonable to speculate that biosynthesis of amino acids is one of the common adaptation pathways Thermococcales use to deal with multiple stresses according to current knowledge.

11.4.3 Compatible solute Compatible solute is a kind of organic compound with small molecular weight, which accumulates inside cells when the microorganism is dealing with extreme environmental stresses. It helps keep intracellular osmotic pressure stable, protects the stability of intracellular macromolecules and thus maintains cell activity [52, 53]. Studies of Thermococcales show that compatible solutes are related to some stresses like heat, high salinity and high hydrostatic pressure.

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There are three main kinds of compatible solute: sugar and its derivatives, amino acids and their derivatives and polyol-phosphodiesters [54]. The accumulation of intracellular compatible solute mainly depends on de novo synthesis and acquisition from the medium [52, 53]. (a) Sugar and its derivatives. Thermophiles mainly use Trehalose (Tre) and Mannosylglycerate (MG) or Glucosyloglycerate (GG) as compatible solutes [55]. Among them, Trehalose is regarded as a general protectant against various environmental stresses; it has been reported under high salinity, heat, acid and oxidative stress. Hyperthermophiles tend to use MG or GG as cell protectants under heat stress because MG and GG have greater ability than Trehalose to protect proteins from heat [56]. Nevertheless, Trehalose is also accumulated and used as a compatible solute in some Thermococcales. (b) Amino acids and their derivatives. Glutamate, Proline, Glycine betaine and Ectoine are usually used as compatible solutes to cope with osmotic pressure in mesophiles. Hyperthermophiles often use Glutamate as a compatible solute [57]. Besides Glutamate, Thermococcales also use Aspartate as a compatible solute [56]. (c) Polyol-phosphodiesters. It is common that polyol-phosphodiesters such as cyclic-2,3-bisdiphosphoglycerate (cDPG), diglycerol phosphate (DGP) and dimyoinositol-phosphate (DIP) are used as compatible solutes in hyperthermophilic archaea and bacteria. cDPG and DGP were reported in a few strains as related to regulation of osmotic pressure, while more strains use DIP as a compatible solute. DIP was reported in the hyperthermophilic archaea Thermococcales Methanococcus, Pyrodictium and Archaeoglobus; the optimum growth temperature for all of them is over 80 °C. DIP was also reported in the hyperthermophilic bacteria Aquifex and Thermotoga [57]. DIP, Glu, Asp, MG and Trehalose arethe main compatible solutes accumulated and used in Thermcoccoales as reported. According to the data from comparative genomes, all Thermococcales strains have synthetic pathways of Glu and Asp and the key synthetic enzyme IMP of DIP. The key genes in the synthesis of MG, MPGS and MPGP exist in the genomes of most Thermococcales, but are not present in some strains like T. kodakarensis, T. onnurineus and T. eurythermalis. The key gene for synthesis of Trehalose, TreT [58], is only present in T. eurythermalis, T. litoralis, T. sibiricus, T. barophilus, P. furiosus and P. horikoshii. Moreover, only T. litoralis was found to use β-galactopyranosyl-5-hydroxylysine (GalHL) as a compatible solute [56]; it was not reported in other Thermococcales strains. Experimental data show that compatible solutes in Thermococcales are usually accumulated under heat, high osmotic pressure and high hydrostatic pressure [59]. In general, different compatible solutes are used under different stresses, but they also have some kind of synergistic effect. In T. kodakarensis, accumulation of DIP under heat stress is 20 times that under optimum conditions, and accumulation of Asp under

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hyperosmosis stress is 4.3 times that under optimum conditions; but after knocking out the key synthetic enzymes of DIP, T. kodakarensis accumulated more Asp to take over the function of DIP under heat stress [60]. In P. furiosus, MG is accumulated under heat stress while DIP is accumulated under hyperosmosis conditions [61]. When knocking out the key synthetic enzymes of MG and DIP separately, it was found that the function of DIP could be completely replaced by MG under heat stress with similar growth yields to the wild type; but when MG was knocked out, the strain could still grow with accumulation of DIP and Asp under high salinity, albeit with worse growth than the wild type [61]. In P. ferrophilus, the only Palaeococcus strain that has undergone further study, MG is accumulated under both heat and hyperosmosis stresses but Glu is only accumulated under heat stress and Asp is only accumulated under hyperosmosis conditions [62]. Besides the heat and hyperosmotic stresses, significant accumulation of MG is also reported in T. barophilus under both high and low hydrostatic pressure stresses in a study of pressure adaption [59]. Transcriptomic and proteomic data of the obligate piezophile P. yayanosii show that the expression of key genes in the synthesis of cDPG was upregulated under high hydrostatic pressure [48], which indicates that accumulation of cDPG may be related to high pressure adaptation, although it has not been detected under other stresses in Thermococcales. An experimental test is required to prove this. Researchers have studied the function of compatible solutes in environmental applications for decades and found variable compatible solutes and their synthetic pathways in response to different environmental stresses. More and more studies suggest that using compatible solutes may be a part of common adaptation strategies to deal with environmental stresses. However, the mechanisms of how compatible solutes regulate osmotic pressure and how they protect cell and macromolecular functions remain unknown and need further study.

11.4.4 Composition of membrane lipid The membrane lipid composition and synthetic pathways of bacteria and archaea are completely different. The composition of membrane lipids of Thermococcales is relatively simpler than that of other archaea. Membrane lipid composition of Thermococcales is composed of sn-2,3-diphytanylglycerol diether, Archaeol, which consists of 20 carbons, and the glycerol dibiphytanyl glycerol tetraether, which consists of 40 carbons [63]. There are two important elements in the synthesis of archaeon membrane lipids: glycerol-1-phosphate (G1P) and isopentenyl diphosphate units, which form geranylgeranyl diphosphate (GGPP, C20 ). Synthesis of G1P comes from the archaeon glycolytic pathway and synthesis of GGPP comes from the mevalonate (MVA) pathway. G1P and GGPP synthesize Archaeol (C20 ) and Archaeol units further synthesize GDGT0 (C40 ) [63–65]. The known membrane lipids of Thermococcales only contain C20 and C40 chains, but membrane lipids of other archaea can have C25 and C80 [66].

320 | 11 Life in multiextreme environments: cross-stress response in Thermococcales

Although the study of archaeon membrane lipids is still in its early stage, we still get some interesting information about archaeon membrane lipids based on current studies. The composition of membrane lipids is closely related to environmental adaptations in archaea. There are two major responses of archaeaon membrane lipids to environmental stresses: (A) change in amount of cyclohexane groups and (B) change of composition of membrane lipids with different-length carbonic chains. Hyperthermophilic Thermococcales are not reported to have cyclohexane groups, hence the research focus on membrane lipid content of different-length carbonic chains. Research on the membrane lipids of T. kodakaraensis showed that the relative content of C20 and C40 is different under heat stress and different cell growth phases. The relative content of C40 increased at the high temperature of 93 °C, while the relative content of C20 increased at the low temperature of 60 °C, compared with the optimum growth temperature of 85 °C. The content of C20 also increased in the exponential phase compared to the stationary phase, while that of C40 increased in the stationary phase [66]. Besides the chain length, the content of polar membrane lipids is related to cell growth phase and supply quality of phosphate [67]. A pressure adaptation study on piezophilic T. barophilus revealed that the composition of membrane lipids is related to a response to pressure stress. The relative content of C20 increased under high hydrostatic pressure and cold stress while the content of C40 increased under low hydrostatic pressure and heat stress. The content of unsaturated nonpolar membrane lipids, which contain irregular polyisoprenoid carbon skeletons, is also influenced by temperature and pressure changes [68]. All these results imply that biosynthesis pathways of membranes are related to responses of Thermococcales to various stresses. However, since the synthetic pathway of lipids in Thermococcales has not been completely revealed, further investigation is required to expose the direct function of membrane lipid synthesis on adaptation.

11.4.5 Antioxidant pathway As obligate anaerobic microorganisms, Thermococcales are different from aerobic microorganisms in that they do not use oxygen as a final electronic acceptor. However, Thermococcales also use antioxidation to remove the damage caused by extra electrons derived from metabolism. The key enzyme used by Thermococcales is superoxide reductase (SOR), which was first reported in P. furiosus [69]. In contrast to superoxide dismutase (SOD) in aerobic organisms, SOR reduces peroxide ions, produced during metabolism, to hydrogen peroxide, and then reduces hydrogen peroxide to water using peroxidase rubrerythrin (Rbr) [70]. Rbr is a specific enzyme in anaerobic microorganisms. The electron carriers during this process are rubredoxin (Rd) and NADP rubredoxin oxidoreductase (NROR) [71]. Besides these, Thermococcales have homologs of flavodiiron (FDP), which belongs to the antioxidation system of aerobic

11.5 Common adaptation strategies for different stresses

| 321

microorganisms. For example, researchers found two predicted FDP proteins FdpA (PF0751) and FdpB (PF0694) in the genome of P. furiosus. However, their exact physiological function is not clear. Obligate anaerobic Thermococcales are reported that can endure traces of oxygen using this anaerobic system. P. furiosus still grew well in the presence of 8% (V/V) oxygen and the productivity of basic metabolic production such as formate was not affected. However, the productivity of hydrogen was reduced by 50%. Knocking out SOR and FbpA raised the sensitivity of P. furiosus to oxygen, which proves that these two proteins have important functions in antioxidation. Knocking out Rd, which is involved in electron transfer, did not significantly affect the strain’s oxygen sensitivity, because other intracellular electronic vectors, such as ferredoxin, NADPH, NADH and so on could replace it [72]. Besides the oxygen stress, anaerobic antioxidation of Thermococcales is also correlated with other environmental stresses. In T. kodakarensis, the redox related proteins peroxiredoxin and oxidoreductase all upregulate under heat stress, and the redox related proteins peroxiredoxin, ferredoxin and thioredoxin reductase all upregulate under hyperosmosis conditions [50].

11.5 Common adaptation strategies for different stresses Since Thermococcales are typical hyperthermophilic microorganisms, the initial studies on Thermococcales mainly focused on the hyperthermophilic mechanisms of adaptation. Researchers tried to expose the mechanisms by searching for the hyperthermophilic enzymes and genes. Then, with more and more piezophilic strains reported in Thermococcales, pressure studies became a new hot topic. However, research on piezophilic strains attempting to find some specific piezophilic genes remain unclear. Researchers gradually discovered that genes reported to be related to piezophilic traits were also reported in adaptation studies of other extreme environmental factors [73]. Thus, it is proposed that pressure adaptation may not be independent but instead associated with adaptation to other stresses [68]. It is possible to devise a hypothesis that proposes a common adaptation strategy for coping with different environmental stresses. So far, most of the studies still focus on hyperthermal traits, and few studies focus on other stresses like high hydrostatic pressure, high salinity and oxidation. Only a few studies refer to cold adaptation and no studies thus far have targeted pH adaptation. Typical hydrothermal vents, where Thermococcales live, are usually associated with dramatic pH fluctuations, a nd most of the neutrophilic Thermococcales strains have wide pH growth ranges. pH conditions are one of the important extreme environments in the original habitats of Thermococcales, and thus is an important environmental factor to help us understand their adaptation strategies to multiple stresses. However, current research on the topic is far from sufficient.

322 | 11 Life in multiextreme environments: cross-stress response in Thermococcales

11.5.1 Common adaptation strategy It is possible to speculate that Thermococcales may have common adaptation strategies for multiple stresses since they have small genomes but wide growth ranges. Based on current data, there is overlapping of their responses to various environmental stresses. Although current research is not complete, what we know about their responses under heat, cold, high hydrostatic pressure, hyperosmosis and oxidative stresses has already provided a large amount of information about adaptation strategies for multiple stresses in Thermococcales. Statistics on the current physiological, biochemical, omic data related to environmental adaptation suggests adaptation is related to some metabolic pathways, including information processes, compatible solutes, chaperonin, amino acid metabolism, membrane lipid synthesis, energy conversion, motility and antioxidation systems (󳶳 Fig. 11.1, 󳶳 Tab. 11.3). Apart from membrane lipid synthesis and energy conversion, studies of which are just beginning, information processes in material metabolism, compatible solutes, chaperonin and amino acid metabolism are involved in all tested adaptations of Thermococcales. This implies that Thermococcales, with their small genomes but wide growth ranges, may have common adaptation strategies to multiple stresses. However, direct evidence of a common adaptation strategy to multiple stresses in one particular strain of Thermococcales remains to be provided.

Fig. 11.1: Stress-response related pathways in Thermococcales

Replication, Transcription, Translation DIP Trehalose MG Glu & Asp cDPG Biosynthesis Protease & Peptidase Transporters GDGT-0(C40) Archaeol(C20) MBH & MBX ATP synthase MCP Flagellum SOR Oxidoreductase

Information process Compatible solute

(9)

(4) (4)

(1) (4) (1) (4) (1) (4) (6) (7)

(9)

(6) (7)

(1) (4) (1) (3) (1) (3) (2) (3) (2) (3)

Heat

(5) (9)

Cold

Stress Responses

(8)

(7) (12) (12) (8) (12) (8) (12)

(8)

(8)

(11)

(12)

High Pressure

(4) (4)

(4) (4)

(4) (3) (3) (2) (3) (2) (3)

Hyperosmotic

(4) (10) (4) (10)

(4)

(4) (4) (4)

(4)

Oxidative

Heat Shock Response by the Hyperthermophilic Archaeon Pyrococcus furiosus [37] Compatible Solutes of the Hyperthermophile Palaeococcus ferrophilus: Osmoadaptation and Thermoadaptation in the Order Thermococcales [62] Effects of temperature, salinity, and medium composition on compatible solute accumulation by Thermococcus spp. [56] Proteome profiling of heat, oxidative and salt stress responses in Thermococcus kodakarensis KOD1 [50] A Mutant Chaperonin That Is Functional at Lower Temperatures Enables Hyperthermophilic Archaea To Grow under Cold-Stress Conditions [74] Effect of growth temperature and growth phase on the lipid composition of the archaeal membrane from Thermococcus kodakaraensis [66] Membrane homeoviscous adaptation in the piezo-hyperthermophilic archaeon Thermococcus barophilus [68] Genome expression of Thermococcus barophilus and Thermococcus kodakarensis in response to different hydrostatic pressure conditions [47] Cold Shock of a Hyperthermophilic Archaeon: Pyrococcus furiosus Exhibits Multiple Responses to a Suboptimal Growth Temperature with a Key Role for Membrane-Bound Glycoproteins [75] (10) Mechanism of oxygen detoxification by the surprisingly oxygen-tolerant hyperthermophilic archaeon, Pyrococcus furiosus [72] (11) Molecular chaperone accumulation as a function of stress evidences adaptation to high hydrostatic pressure in the piezophilic archaeon Thermococcus barophilus [59] (12) High hydrostatic pressure adaptive strategies in an obligate piezophile Pyrococcus yayanosii [48]

(1) (2) (3) (4) (5) (6) (7) (8) (9)

Antioxidation

Motility

Energy metabolism

Membrane lipids

Amino acids metabolism

Function

Pathway

Tab. 11.3: Current reported pathways related to Thermococclaes adaptation

11.5 Common adaptation strategies for different stresses | 323

324 | 11 Life in multiextreme environments: cross-stress response in Thermococcales

11.5.2 Summary of adaptation strategy To sum up, the main strategies that Thermococcales use to deal with multiple stresses include: (a) having various ATP acquiring pathways and the ability to use low energy quanta [38] to ensure enough energy to support big biomass in deep environments with low energetic supply; (b) taking full advantage of sufficient amino acids in the original habitat to avoid enormous energy consumption by de novo synthesis [39]; (c) synthesizing and accumulating various compatible solutes and chaperonin as cell protectants to maintain intracellular homeostasis and prevent macromolecules from deconstruction caused by stresses; (d) changing membrane lipid composition to adjust fluidity and stability of the cell membrane; and (e) using anaerobic antioxidation to remove extra electrons in metabolism. Among them, information process, amino acid metabolism and membrane lipid composition are important compositions of substance metabolism; energy conversion and mobility systems are related to the production and consumption of intracellular energy; and antioxidation is the link between substance metabolism and energy metabolism (󳶳 Fig. 11.1). These special metabolic pathways help Thermococcales to survive under mult-stress environments near hydrothermal vents. We anticipate that more studies will focus on the common adaptation strategy of Thermococcales to multiple stresses. Thermococcales are the predominant microorganisms in the early-Earth-like environment of hydrothermal vents, and are near the root of the phylogenetic tree. Whether the common adaptation strategy of Thermococcales is general to other microorganisms and how the common adaptation works in dealing with changes in diverse environmental conditions will be the main topics for study and discussion. Linking our knowledge of common adaptation in Thermococcales to environmental changes during the history of the Earth will provide precious models and research insights for the study of early life traits and help us acquire a better understanding of the origin of life.

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Index 13 C

230, 237, 245, 246 radiotracer 228, 229, 231 16S rRNA 149 16S rRNA gene 235, 239 35Flor 265, 271, 277, 282, 283, 287 35 S radioisotopes 232 14 CH 4

A abiogenic hydrocarbons 109 abiogenic methane 126 abiogenic methane formation 109 abiotic reactions 181, 203, 210 accretionary prism 57 acetoclastic pathway 6 acetogenesis 128 Acharax 12, 144 Achromatium volutans 92 acidic 90 Acidobacteria 92, 149 acidophiles 308 Acridine Orange Direct Count 148 Actinobacteria 149 activity product 184 adaptation 88 adaptation strategies 322 advection-diffusion-reaction models 225 advective fluid flow 61 advective methane flux 20 aerobic 10 aerobic methane oxidation 11, 20, 228 aerobic methanogenesis 6 aerobic methanotrophic bacteria 21 aioA-like functional genes 98 alkaliphiles 308 alkaliphilic 124 alkane molecules 237 Alphaproteobacteria 201 Alviniconcha 196, 202 Alvinocaris longirostris 145 amino acids 316 ammonia – monooxygenase 187, 207 – oxidation 191, 206 – oxidation rates 206 – oxidizing archaea (AOA) 204, 207 – oxidizing bacteria (AOB) 204, 207 https://doi.org/10.1515/9783110493672-012

ammonia oxidizing archaea 237 ammonium transporter 187 amo 187 amt 187 anaerobic 10 anaerobic methane oxidation 10 Anaerobic methane-oxidizing Archaea (ANME) 129 anaerobic methanotroph 223, 231, 236, 241 anaerobic oxidation of methane 223 – metal-associated 233 – nitrate-associated 231, 233 – nitrite-associated 233 – sulfate-dependent 227, 230 anammox 186, 197, 198, 201 ANME 64, 153 annual CH4 flux 21 anthropogenic methane 23 anthropogenic methane source 23 anthropogenic warming 15 antioxidants 320 AODC 148 AOM 10, 11, 20, 64, 70 AOM community 20 archaea 6, 10, 87 archaeal diversity 152 Archaeoglobales 93 Archean age 4 archive 2, 12 Arcobacter 93 Arctic 18 arctic 17 Arctic Ocean 18, 19, 23 arsenate 96 arsenic 90, 96 arsenite 96 arsenotrophy 98 asphalt 2 As-sulfides 91 Atlantis Massif 116 atmosphere 14, 21, 23 atmospheric concentration 22 atmospheric methane 13, 19, 21–23 atmospheric methane sources 21 atmospheric temperature 16 ATP synthase 283, 286, 315

332 | Index

authigenic carbonate 12 Automated Ribosomal Intergenic Spacer Analysis 92 autotrophic bacteria 198 Awaruite 109 B B18LD 265, 271, 283, 287 back arc 57 back-arc basins 162 Bacteria 87 Bacteroidetes 93 baeocyte 263 Bathyarchaeota 152 bathymetric mapping 55 Bathymodiolus 12, 144 Beggiatoa 65, 200, 208, 261, 262, 265, 277, 282, 283, 287 Beggiatoa alba 265, 271, 283, 287, 300 Beggiatoa leptomitiformis 265–267, 271, 283, 287, 291, 300 Beggiatoaceae 261, 266, 267, 271, 275, 278, 286, 288, 296 Betaproteobacteria 127, 201 Big Bacteria 65 bin 4572_84 283 bioavailability 96 biogeochemical processes 88 biological filters 73 biology 5 biomarkers 111, 117 Bioorthogonal non-canonical amino acid tagging 239, 247 biosynthesis 195, 316 bivalve 11 Black Sea 232 black smoker 1 Blake Plateau 19 Blake Ridge formation 19 BOGUAY 266, 277, 282, 283 brine 4 brine seep 2, 4 bubble 9, 19–21, 23 bud S10 265, 268, 271, 283 Burkholderia cenocepacia 299 C Ca2+ 12 Cabeço de Vide 113, 119

Ca-carbonate 12 calcium 12 Calvin–Benson-Bassham (CBB) pathway 166 Calyptogena 144 cantilevers 244 cap carbonate 5 carbon dioxide 20, 98 carbon monoxide 115 carbon reservoir 14 carbon sequestration 130 carbon sink 12 carbon source 20 carbonate 2, 12, 20 carbonate deposit 12 carbonate mineral 12 carbonate precipitation 12 carbonate reduction pathway 6 carbonate reefs 144 carbonates 141, 226, 234, 244 cell division 261–263, 275, 282, 283, 300, 301 Cell Quantification 233 Central Indian Ridge (CIR) 162, 163 CH4 liberation 19, 22 CH4 sources 7 CH4 flux rate 20 chemical disequilibrium 180, 184 chemical gradient 4 chemical-biological coupling 5 chemoautotrophic 3 chemoautotrophy 86 chemodenitrification 197, 210 chemolithoautotrophy 165 chemosynthesis 107 chemosynthetic 2, 10 chemosynthetic bacteria 63 chemosynthetic ecosystems 139 chemosynthetic microorganisms 86 chemotrophic microorganisms 308 China Ocean Mineral Resources Research and Development Association (COMRA) 165 Chlorobium 149 Chloroflexi 149 clasts 54 clathrate 7 clathrate gun hypothesis 13 clay dehydration 61 climate 2, 23 climate change 13, 14, 21, 23, 88, 98 climate modeling 16

Index | 333

climate warming 18 closed system 189 Clostridiales 128 CO2 6, 20 CO2 emissions 89 CO2 fixation 64 CO2− 3 12 Coast Range Ophiolite Microbial Observatory (CROMO) 112, 120 coccolithophorid 6 cold seep 1, 139 cold seepage 23 cold vent 1 comammox 204, 210 Comamonadaceae 127 Conchocele 144 consortium 10 continental margin 2, 16, 17, 108, 116, 139 continental slope 2, 16, 17 Convergent Margin 58 coupled isotopes 190, 206 crab 12 crater 3 crude-oil 2 cryovolcanism 130 cSDR 299 cultivation 92 culture independent 224, 233 Cyanobacteria 261, 264, 266, 287, 293, 296 Cytophaga-Flavobacterium 92 D D-402 265, 271, 283, 287, 291 dcw gene cluster 267, 268, 271, 283 deep-sea mining 162 Deferrisoma palaeochoriense 92 degradation 10 δ13 C 7 δD 7 deltaic sediment 58 Denaturing Gradient Gel Electrophoresis 92 denitrification 197 – chemoautotrophic 184, 201 – rates 199 destabilization 10, 16 Desulfacinum hydrothermale 92 Desulfotomaculum 128 deuterium 236, 246 – monodeuterated methane 228

diapiric intrusion 61 diffusion 225 dimethylsulfide 6 dimethylsulfoniopropionate 6 disproportionation 11 dissimilatory nitrate reduction to ammonium (DNRA) 202 dissociation rate 16 dissolved amino acids 99 dissolved inorganic carbon (DIC) 126, 229, 236 dissolved organic matter 90 dissolved organic nitrogen 181, 211 dissolved organic sulfur 90 division and cell wall (dcw) gene cluster 266 DMSP 6 DNA 192 DNA replication 261, 263, 287, 291, 296, 297, 299 DnaA 261, 287, 291–293, 296–301 dnaN 287, 289, 291, 293, 297 DNRA 186, 200, 202 DON 211 dynamin 261, 282, 284 E East Pacific Rise 162 Eel River Basin 226, 228 elevated temperatures 88 Emiliania huxleyi 6 Enceladus 130 endemic fauna 2 endo-symbiont 11, 291 endo-symbiotic 11 energy accumulation 15 energy sources 179 ephemeral 4 Epsilonproteobacteria 93, 173, 200 equilibrium constant 184 extraterrestrial 4 extreme ecosystems 85 extremophiles 86 F fault slip events 66 faulting 58 faunal community 10 ferrous iron 109 Firmicutes 149 Fischer-Tropsch-type reactions 107

334 | Index

fissure 2, 10 flair 9 flaire 18 fluid 1 fluid advection 4 fluid conduit 4 fluid flow 2, 12 fluorescence in situ hybridization 234, 235, 239, 246 flux 4 forearc 116 forearc basin 57 formate 109 formate dehydrogenase 316 Fourier transform ion cyclotron resonance 90 FtsA 267, 276 ftsW 267 FtsZ 266–268, 272–274, 276, 277, 282, 298, 301 ftsZ 266–268, 291 functional genes 192, 207, 208 fungi 211 G Gakkel Ridge 162 Gammaproteobacteria 93, 173, 261, 262, 266, 267, 272–274, 277, 286, 287, 290, 293, 300 gas hydrate 7, 9, 17 gas hydrate deposit 8, 16 gas hydrate destabilization 16, 18–20, 22, 23 gas hydrate distribution 16 gas hydrate formation 17 gas hydrate inventory 23 gas hydrate reservoir 17, 23 gas hydrate stability 12 gas hydrate stability zone 8 GC skew 287, 291, 294–300 gdh 187 genetic markers 207 genome copy number 296, 298, 301 genomics 223 geochemical gradients 183 geochemistry 225 geofluids 10 geomicrobiology 85 GH formation 9 GHSZ 8 GHSZ retreat 15 Gibbs free energy 184

gln 187, 196 global climate 14 global gas hydrate deposits 17, 18 global gas hydrate inventory 17 global inventory 17 global methane budget 21 global methane emission 21 global methane sources 23 global warming 1, 10, 13, 14, 22, 23 glt 187, 196 glutamate dehydrogenase 187 glutamate synthase 187 glutamine synthetase 187 green house effect 13 green house gas 1, 13, 14, 22 GS-GOGAT 187 Guaymas Basin 230, 237, 241, 263, 264, 266, 283, 301 Gulf of Mexico 230, 238 H Håkon Mosby 4, 69 halophilic 4 heat capacity 14 heat propagation 15 heat reservoir 14 heat shock proteins 314 heat transfer 14, 18, 23 heat transport 15 Hellenic Volcanic arc 89 heterogeneous catalysts 109 heterotrophic bacteria 197 heterotrophs 99 high pressure 241 homoacetogenesis 126, 152 HS− 11 Hydrate Ridge 225, 230, 234, 236, 238, 241, 243, 265 hydrate-bound methane 22 hydraulic activity 3 hydrocarbon reservoir 58 hydrocarbons 1, 229, 248 hydrofracturing 59, 139 hydrogen 109 hydrogen sulfide 11 hydrogenase 315 hydrogenotrophy 114 hydrothermal 263, 266 hydrothermal activity 91

Index | 335

hydrothermal plume 167 hydrothermal sulfides 165 hydrothermal vents 85, 162, 307 hydroxylamine 181 Hyperalkaline 111 hyperalkaline conditions 111 hyperosmosis 319 hypersaline 4 hyperthermophiles 95, 318 hyperthermophilic archaea 92 hyperthermophlic microorganisms 308 I incubations – ex situ 199, 206 – in situ 199 – isotope labeling 199 intact polar lipids 208 International Continental Drilling Program (ICDP) 118 International Seabed Authority (ISA) 163, 164 interspecies interactions 244 inventory 9 iron reducers 92 iSDR 299 isotopic equilibration 191 isotopic fractionation 188, 199 Izu-Bonin-Mariana subduction zone 111 K kerogens 5 kinetic barriers 183 kinetic isotope fractionation 6 Kuenenia 201 L ladderane lipids 202 Lake Okotanpe 266 large sulfur-oxidizing bacteria 262 last glacial maximum 13 last universal common ancestor 130 laterites 112 ligands 96 Longqi 172 Lost City Hydrothermal Field 111, 116 LSB 261–267, 277, 282, 283, 287, 291, 299, 300 M magnesium 12 MamK 280, 282

Mariana forearc 118 Maribeggiatoa 263, 264, 266 marine sediment 14 marine shallow-water hydrothermal vents 85 Mars 130 mass spectrometry 238, 244 membrane lipids 319 metabolic activity 223 – anabolism 233 – catabolism 227 – observational approaches 225 metacaspase 266 metagenome 246 metagenomics 117, 207, 208 metal 96 metatranscriptomics 207, 208 methane 1, 109, 139 methane bubble 9 methane budget 1, 22 methane diffusion 19 methane emission 1, 23 methane flux 12, 20 methane isotopologues 109 methane metabolism 227 – complete methane oxidation 229 – methane activation 227 – methanogenesis 231 – methanotrophy 229 – net methane consumption 231 methane release 17 methane seep 2, 21, 223 methane seep sediments 234 methane sink 21, 22 methane transport 19 methane venting 18 methane-consuming microbes 21 methanogenesis 64, 126, 231 methanogens 152, 227 methanol 6 methanotrophy 11, 114, 126 methyl coenzyme M reductase 231, 238 methylamine 6 methylated amines 6 methylphosphonate 6 Mg2+ 12 microbes 2, 10 microbial colonization 226 microbial community 10 microbial community composition 88

336 | Index

microbial diversity 88 microbial filter 23 microbial methane 6 microbial methane filter 20, 21, 23 microbial methane oxidation potential 19 microbial sulfur cycling 166 microbial water column methane filter 22 microcalorimetry 245 micronutrient 96 microscopy 240 Mid-Atlantic Ridge 109, 162, 312 Mid-Cayman Rise 117 mid-ocean ridges 108 mixing 180, 189, 205 – endmember fluids 180 – in the subsurface 180 – low temperature fluids 180 – mixing line 181 – non-conservative behavior 198 mixotrophy 229 MOx 20 MOx activity 20 MOx bacteria 20 mraZ 267 MreB 277, 278, 280, 282, 301 mRNA 192 mud breccia 54, 60, 66, 69, 77 mud diapirism 58 mud pools 54 mud volcanism 10 mud volcano 2, 4, 53 multi-cellular organisms 211 multiple stress factors 324 multiple-extreme environments 307 mussels 12 N NADH dehydrogenase 283, 286 Namibian upwelling 265 natural gas 2 natural laboratory 92 natural methane emission 23 natural methane source 23 natural seep 23 New Caledonia 118 nif 187, 194 nir 186 nirK 191 nirS 191

nitrate – dissimilatory reduction to ammonium (see also DNRA) 186, 196 – reductase 201 – reductase, assimilatory (nas) 187 – reductase, dissimilatory (nar) 186, 201 – reductase, periplasmic (nap) 191, 196, 201 – reduction 186 – reduction, assimilatory (nas) 196 nitric oxide 181 – reductase 186, 201 nitrification 186, 191, 204 – heterotrophic 209, 210 nitrite – oxidation 191, 210 – oxidizing bacteria 204 – oxidoreductase 187 – reductase, assimilatory (nirB) 187 – reductase, dissimilatory (nir) 186, 201 – reductase, to ammonium (nrf) 186 nitrogen 179 nitrogen assimilation 187, 195 nitrogen fixation 193, 235 nitrogen sources 188, 191 nitrogenase 187, 193 – alternative 194, 210 Nitrospirae 204, 208 nitrous oxide 181 – reductase 186, 201 nor 186 nos 186 nrf 186 nxr 187, 208 O ocean acidification 20 Oceanic Methane Paradox 6 oil field 3 oil seep 2 olivine 109 omics 192 ophiolites 108 organic matter 10, 20 oriC 287, 289, 290, 299, 300 origin 261, 287, 291, 297, 299 origin of life 87, 108, 324 origins of life 179, 209 overpressured sediments 57 oxygen 20

Index |

P Palaeochori Bay 88 Paleocene–Eocene Thermal Maximum 13 Par 299 parAB 287 past climate 13 Pearl River 144 peptidoglycan 267, 275, 276, 282, 300 peridotite 114 permafrost 17, 23 Permian–Triassic extinction event 13 PETM 13 petroleum 2 pH gradients 98 phase boundary 8 phase separation 96 piezophiles 308 Planctomycetes 149 plant-microbe interactions 116 plate tectonics 108 pockmark 2, 3 polymetallic sulfides 162 porewater 19 potential chemical energy 180 preindustrial times 22 pressure adaptation 312 prokaryotic 6 Prony Bay 118 pyrolysis 191 pyroxene 109 Q quantitative PCR 192, 234 R Rainbow vent field 109 Raman spectroscopy 111, 246, 247 Rayleigh distillation 189 RctB 299 reactive intermediates 184 rearrangement 267, 278, 281, 283, 284, 286, 297, 300 recombination 267, 297, 299, 300 redox reactions 183 reduced chemical species 180 reductive tricarboxylic acid (rTCA) cycle 166 remineralization 181, 187 RepA 299 residence time 13

337

reversed tricarboxylic acid cycle 162 Ridgeia piscesae 202 Riftia pachyptila 195 rnpE 287 rpmH 287 RT-qPCR 192 S salt diapir 58 Samail Ophiolite 111, 116 Santa Elena Ophiolite 113 Scalindua 201 scaly-foot gastropod 164 SCS 139 sea floor mapping 69 sea level rise 13 sea-air methane flux 19 seafloor 23 secondary ion mass spectrometry 235, 239, 244 sediment 20 sediment translocation 3 sediment-water interface 11 seep chemistry 5 seep fauna 10 seep morphology 1 seeps 1 septum 261, 282 septum formation 267 Serpentine soils 122 serpentinite 58 serpentinite springs 111 serpentinization 107 Serpentinomonas 127 Shinkaia crosnieri 145 side scan sonar 55 Single Cell Growth Rate 244 Single Cell Omics 247 slope failure 4 smectite-illite transformation 61 SMTZ 148, 151 SO2− 4 11 South China Sea 139 Southeast Indian Ridge (SEIR) 163 Southwest Indian Ridge (SWIR) 161, 163 Spathi Bay 88 spatial scales 241, 244 Spirochaetes 149 stable isotope probing 234 – DNA 236

338 | Index

– experiment set-up 235 – lipids 236 – proteins 238 – RNA 237, 245 stable isotope signature 6 stable isotopes 188 storativity 113 subduction zone 4, 62 subseafloor sediments 238 subsurface fluid flow 243 Subsurface lithoautotrophic microbial ecosystems (SLiMEs) 126 sulfate 11 sulfate reducing bacteria 223, 238, 241, 245 sulfate reduction 3, 92 sulfate reduction rates 232 sulfate-reducing bacteria 6, 10 sulfide 2, 200, 202 sulfide flux 12 sulfide oxidation 10 sulfide-oxidizing bacteria 12, 200 sulfidic 90 sulfur catabolism 232 sulfur cycle 117 sulfur oxidizers 173 sulfur oxidizing bacteria 226 Sulfurovum 93 sulphide oxidation 64 SUP05 201 superoxide dismutase 320 superoxide reductase 320 surface layer 8 SWIR 161 symbiont 11 symbiosis 86 symbiotic bacteria 11 Synechococcus elongatus PCC 7942 296–299 syntrophic 10 T Tablelands ophiolite 113 tectonic faults 10 Tekirova ophiolite 115 terminal electron acceptor 10 terrestrial methane sources 21 Thaumarchaeota 152 thermal capacity 14 thermal conduction 15 thermal decomposition 181 thermocline 8

Thermococcales 308 Thermococcus 93 thermodynamic potential 184 thermogenic methane 5 thermophiles 95, 308 thermophilic organisms 92 Thermoplasmatales 93 Thermoproteales 93 thermosome 314 Thio36 265, 267, 271, 277, 282, 283 Thiomargarita 261, 262, 265–268, 271, 275–277, 282, 283, 300 Thiomargarita spp. 266, 267 Thiomicrospira 92 Thioploca 65, 92 Thioploca ingrica 265–267, 271, 277, 282, 283, 287, 291 toxicity 96 trace methane oxidation 230 travertine 115 tritiated methane 228 tube worm 11, 12 turbidites 148 U ultrabasic conditions 111 ultramafic intrusions 108 ultramafic rocks 107 underthrusting 62 V vacuole 263, 265, 275, 282 Verrucomicrobia 149, 239 Vestimentiferan 70 Vibrio cholerae N16961 299 vitrinite reflectance 61 volcanism 89 W water column 20 white mats 91 Woesearchaeota 173 Y yidC 287 yidD 287 Z Zambales ophiolite 119 ZipA 267, 276