Microbial Life in the Cryosphere and Its Feedback on Global Change 9783110497083, 9783110496451

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Susanne Liebner, Lars Ganzert (Eds.) Microbial Life in the Cryosphere and Its Feedback on Global Change Life in Extreme Environments

Life in Extreme Environments

Series Editor Dirk Wagner

Volume 7

Microbial Life in the Cryosphere and Its Feedback on Global Change Edited by Susanne Liebner and Lars Ganzert

Editors Prof. Dr. Susanne Liebner GFZ German Research Centre for Geosciences, Helmholtz Centre Potsdam Section Geomicrobiology Telegrafenberg 14473 Potsdam, Germany [email protected]

Series Editor Prof. Dr. Dirk Wagner GFZ German Research Centre for Geosciences, Helmholtz Centre Potsdam Section Geomicrobiology Telegrafenberg 14473 Potsdam, Germany [email protected]

Dr. Lars Ganzert GFZ German Research Centre for Geosciences, Helmholtz Centre Potsdam Section Geomicrobiology Telegrafenberg 14473 Potsdam, Germany [email protected]

ISBN 978-3-11-049645-1 e-ISBN (PDF) 978-3-11-049708-3 e-ISBN (EPUB) 978-3-11-049390-0 ISSN 2197-9227 Library of Congress Control Number 2020950604 Bibliographic information published by the Deutsche Nationalbibliothek The Deutsche Nationalbibliothek lists this publication in the Deutsche Nationalbibliografie; detailed bibliographic data are available in the Internet at http://dnb.dnb.de. © 2021 Walter de Gruyter GmbH, Berlin/Boston Cover image: Axel Kitte, Helmholtz Centre Potsdam, GFZ German Research Centre for Geosciences Typesetting: Compuscript Ltd., Shannon, Ireland Printing and binding: CPI books GmbH, Leck www.degruyter.com

Preface The cryosphere describes environments where water appears in a frozen form. It includes permafrost, glaciers, ice sheets, and sea ice and is currently more affected by Global Change than most other regions of the Earth. In the cryosphere, limited water availability and subzero temperatures cause extreme conditions for all kinds of life, but many microorganisms can cope with this extremely well. The cryosphere’s microbiota displays an unexpectedly large genetic reservoir and taxonomic as well as functional diversity, which, however, we are still only beginning to map. Moreover, microbial communities influence the response mechanisms of the cryosphere towards Global Change. Altered patterns of seasonal temperatures and precipitation are already occurring in many cryotic ecosystems, affecting microbial element cycling, adaptation strategies, and physiology. Activation of nutrients by thawing permafrost and increasing thaw layer thickness, as well as erosion, for example, renders nutrient stocks accessible to microbial activities. Also, glacier melt and retreat stimulate microbial life, which in turn influences albedo and surface temperatures. In this context, the functional resilience of microbial communities in the cryosphere is of major interest and concern. Particularly important is the ability of microorganisms, and microbial communities as a whole, to respond to changes in their surroundings by intracellular regulation and population shifts within functional and ecological niches, respectively. Research on microbial life exposed to permanent freeze or seasonal freeze-thaw cycles has led to astonishing findings about microbial versatility, adaptation, and diversity. Microorganisms thrive in cold habitats and new sequencing techniques have produced large amounts of genomic, metagenomic, and metatranscriptomic data that allow insights into the fascinating microbial ecology and physiology at low and subzero temperatures. Moreover, some of the frozen ecosystems such as permafrost in Siberia, Canada, and Alaska constitute major global carbon and nitrogen storages but can also act as sources of the greenhouse gases methane and nitrous oxide impacting our climate globally. We hope that with this book, you find exciting insights into the life strategies of single microbial species and the entire microbiome in prominent ecosystems of the cryosphere. You will learn about microbial climate feedbacks and multiple response mechanisms to environmental change and how this potentially changes biogeochemical cycles of the cryosphere and beyond. Susanne Liebner and Lars Ganzert

https://doi.org/10.1515/9783110497083-202

Volume published in the series Volume 1 Jens Kallmeyer, Dirk Wagner (Eds.) Micobial Life of the Deep Biosphere ISBN 978-3-11-030009-3 Volume 2 Corien Bakermans (Ed.) Micobial Evolution under Extreme Conditions ISBN 978-3-11-033506-4 Volume 3 Annette Summers Engel (Ed.) Micobial 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 Volume 5 Jens Kallmeyer (Ed.) Life at Vents and Seeps ISBN 978-3-11-049475-4 Volume 6 Natuschka Lee (Ed.) Biotechnological Applications of Extremophilic Microorganisms 978-3-11-042773-8 Volume 8 Étienne Yergeau Advanced Techniques for Studying Microorganisms in Extreme Environments ISBN: 978-3-11-052464-2

Contents Preface   v Contributing authors 

 xv

Susanne Liebner and Christian Knoblauch 1 Microbial carbon dioxide and methane turnover and the permafrost carbon  1 feedback   1 1.1 The permafrost carbon feedback   3 1.2 Carbon decomposition and GHG release from thawing permafrost    1.3 Microbial consumption of methane in permafrost environments  7  8 1.4 Temperature response of methane oxidizers   10 1.5 Interaction of methane oxidizers and mosses      1.6 Methane oxidation in submarine permafrost 13  16 1.7 Summary   18 References  Rachel Mackelprang, Nesliah Tas and Mark Waldrop 2 Functional response of microbial communities to permafrost thaw   27 2.1 Introduction   28 2.2 Activity in permafrost      2.3 Response to thaw 30  35 2.4 Future directions   36 References  Rhiannon Mondav 3 Genomic inventory of permafrost microorganisms   43 3.1 Introduction   43 3.2 A very recent history   45 3.3 (Meta)metagenomics of permafrost      3.4 Future directions 49  50 3.5 Summary   50 References 

 27

 43

Marcus A. Horn and Stefanie A. Hetz 4 Microbial nitrogen cycling in permafrost soils: implications for atmospheric  53 chemistry   53 4.1 Introduction   61 4.2 Microbial processes of the N-cycle   61 4.2.1 Biological nitrogen fixation (BNF) and associated organisms   68 4.2.2 Nitrification and associated organisms 

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4.2.3 Dissimilatory nitrate reduction and associated organisms   86 4.2.4 Anaerobic ammonia oxidation   88 4.2.5 Ammonification   92 4.3 Conclusions      References 92

 76

Dirk Wagner 5 Methanosarcina soligelidi – a multi-resistant archaeon  113 from Siberian permafrost   113 5.1 Introduction  5.2 Isolation and characterization of Methanosarcina  114 soligelidi SMA-21  5.3 Tolerance of Methanosarcina soligelidi against environmental  116 stress   117 5.3.1 Response to subfreezing temperatures      5.3.2 Desiccation tolerance 119  122 5.3.3 Tolerance against salt stress   124 5.3.4 Oxygen sensitivity   125 5.3.5 Radiation resistance  5.4 Survival of Methanosarcina soligelidi under Mars analog conditions –  128 implication for potential life on Mars  5.4.1 Survival of Methanosarcina soligelidi under Martian thermo-physical  129 conditions  5.4.2 Methane production activity under simulated Martian subsurface  133 conditions   134 5.5 Conclusion   135 Acknowledgments      References 135 Lada Petrovskaya, Ksenia Novototskaya-Vlasova, Anastasia Komolova and Elizaveta Rivkina 6 Biochemical adaptations to the permafrost environment:  141 lipolytic enzymes from Psychrobacter cryohalolentis K5T   141 6.1 Introduction   142 6.2 Psychrobacter cryohalolentis K5T and its lipolytic enzymes      6.2.1 EstPc 142  143 6.2.2 Lip1Pc   144 6.2.3 Lip2Pc   144 6.2.4 AT877  6.3 Ecological adaptations of lipolytic enzymes from  146 P. cryohalolentis K5T   148 6.4 Conclusions 

Contents 

Acknowledgments   149 References 

 149

Beat Frey  153 7 Microbial ecology of mountain permafrost: The Alps   153 7.1 Microbes in alpine permafrost   153 7.2 Climate change in low-latitude permafrost regions      7.3 The alpine permafrost microbiome 155  155 7.3.1 Bacteria   157 7.3.2 Archaea   162 7.3.3 Fungi   163 7.4 Future perspectives      Acknowledgments 166  166 References  Sizhong Yang 8 Microbial ecology of Alpine frozen ground:  173 the Qinghai-Tibet Plateau   173 8.1 Introduction  8.2 Soil organic carbon (SOC) and greenhouse gas (GHG) emission from  174 Tibetan frozen ground   175 ­8.3 Microbial enumeration in Tibetan permafrost soil   175 8.4 Novel species from the Tibetan frozen ground soils      8.4.1 Bacterial isolates 176  178 8.4.2 Archaeal isolates   180 8.5 Microbial taxonomic composition   180 8.5.1 Bacterial community   181 8.5.2 Methanotrophic community      8.5.3 Archaeal community 181  184 8.5.4 Fungi diversity   184 8.6 Shaping environmental factors   186 8.7 Habitat filtering vs. geographical distance   186 8.8 Microbial response/feedback to climate change   188 8.9 Conclusion and perspectives   189 Acknowledgments   189 References  Sophie Crevecoeur and Connie Lovejoy  197 9 Microbial assemblies of Arctic freshwater systems      9.1 Characterization of Arctic freshwater systems 197  198 9.2 Microbial component of the Arctic diversity   199 9.2.1 Bacteria 

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9.2.2 Archaea  9.2.3 Protists  9.2.4 Viruses 

 201  202  203

 204 9.3 Microbial assembly rules    9.4 Microbial biogeography  204 9.5 Environmental control of microbial communities   207 9.6 Impact of climate change      References 207

 206

Riitta Nissinen and Manoj Kumar  213 10 Plant-associated microbes in the Arctic      10.1 Arctic vegetation 213  214 10.2 Plant microbiomes   215 10.3 Plant-associated fungi in the Arctic   218 10.4 Bacterial communities in Arctic plants      10.5 Nitrogen-fixing bacteria 220 10.6 Arctic plant microbiomes are endemic to cold climates  222 and show seasonal dynamics   222 References  Ewa Poniecka and Elizabeth A. Bagshaw  227 11 The cryoconite biome   227 11.1 Cryoconite biota   228 11.1.1 Molecular analysis of biota      11.1.2 Photoautotrophs 229  229 11.1.3 Heterotrophs   231 11.1.4 Anaerobes   233 11.2 Biogeochemistry      11.2.1 Nutrients 233  233 11.2.2 Pollutants   234 11.3 Future   235 References  Stefanie Lutz and James A. Bradley  239 12 Glacial surfaces: functions and biogeography    12.1 Introduction  239  239 12.2 Life on glacial surfaces   239 12.2.1 Supraglacial habitats   241 12.2.2 Prokaryotes   241 12.2.3 Eukaryotes      12.2.4 Viruses 243  243 12.2.5 The biogeography of supraglacial ecosystems 

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Contents 

 244 12.3 Glacial functions   244 12.3.1 Carbon and nutrient cycling and export   245 12.3.2 Albedo and melting   245 12.3.3 Biotechnology   245 12.4 Future outlooks  12.4.1 Impact of global climate change on glacier surface functions   246 12.4.2 Future challenges and questions      Acknowledgments 247  247 References 

 245

James A. Bradley  253 13 Microbial dynamics in forefield soils following glacier retreat      13.1 Introduction 253  253 13.2 Soil development following glacier retreat   255 13.3 Microbial communities in glacier forefield soils  13.4 Microorganisms are drivers of glacier forefield biogeochemical  256 cycles   258 13.5 Future climate change   259 13.6 Future challenges   261 Acknowledgments      References 261 Index 

 265

Contributing authors Elizabeth Bagshaw School of Earth and Environmental Sciences Cardiff University Main Building, Park Place CF10 3AT Cardiff, UK [email protected]

Marcus A. Horn Institute of Microbiology, Leibniz University Hannover Herrenhäuser Strasse 2 30419 Hannover, Germany [email protected]

James A. Bradley School of Geography Queen Mary University of London London, E1 4NS, United Kingdom AND GFZ German Research Centre for Geosciences, Helmholtz Centre Potsdam Section Interface Geochemistry Telegrafenberg 14473 Potsdam, Germany [email protected]

Christian Knoblauch Institute of Soil Sciences, Universität Hamburg Allende Platz 2 20146 Hamburg, Germany [email protected]

Sophie Crevecoeur Canada Centre for Inland Waters, Water Science and Technology Branch, Watershed Hydrology and Ecology Research Division, Environment and Climate Change Canada, Burlington, Ontario, Canada [email protected] Beat Frey Swiss Federal Research Institute WSL Zürcherstrasse 111 8903 Birmensdorf, Switzerland [email protected] Stefanie A. Hetz Institute of Microbiology, Leibniz University Hannover Herrenhäuser Strasse 2 30419 Hannover, Germany [email protected]

Anastasia Komolova Institute of Physicochemical and Biological Problems in Soil Science, Russian Academy of Sciences 142290 Pushchino, Moscow Region, Russian Federation [email protected] Manoj Kumar Natural Resources Institute Finland Eteläranta 55 96300 Rovaniemi, Finland [email protected] Susanne Liebner GFZ German Research Centre for Geosciences, Helmholtz Centre Potsdam Section Geomicrobiology Telegrafenberg 14473 Potsdam, Germany [email protected] Connie Lovejoy Département de Biologie, Takuvik Joint International Laboratory, Institut de Biologie Intégrative et des Systèmes (IBIS) and QuébecOcéan, Université Laval, Québec, Canada [email protected]

xvi 

 Contributing authors

Stefanie Lutz GFZ German Research Centre for Geosciences, Helmholtz Centre Potsdam Section Interface Geochemistry Telegrafenberg 14473 Potsdam, Germany [email protected]

Elizaveta Rivkina Institute of Physicochemical and Biological Problems in Soil Science, Russian Academy of Sciences 142290 Pushchino, Moscow Region, Russian Federation [email protected]

Rachel Mackelprang California State University Northridge 18111 Nordoff St. 91330 Northridge, CA, USA [email protected]

Neslihan Tas Lawrence Berkeley National Laboratory 1 Cyclotron Road, MS 70A-3317 94720 Berkeley, CA, USA [email protected]

Rhiannon Mondav Uppsala University, Department of Ecology and Genetics, Evolutionsbiologiskt Centrum (EBC) Norbyvägen 18 D 75236 Uppsala, Sweden [email protected]

Dirk Wagner GFZ German Research Centre for Geosciences, Helmholtz Centre Potsdam Section Geomicrobiology Telegrafenberg 14473 Potsdam, Germany [email protected]

Riitta Nissinen University of Jyväskylä, Department of Biological and Environmental Science P.O. Box 35 40014 Jyväskylä, Finland [email protected] Ksenia Novototskaya-Vlasova Institute of Physicochemical and Biological Problems in Soil Science, Russian Academy of Sciences 142290 Pushchino, Moscow Region, Russian Federation [email protected] Lada Petrovskaya Shemyakin & Ovchinnikov Institute of Bioorganic Chemistry, Russian Academy of Sciences ul. Miklukho-Maklaya 16/10 117997 Moscow, Russian Federation [email protected] Ewa Poniecka School of Earth and Environmental Sciences Cardiff University Main Building, Park Place CF10 3AT Cardiff, UK [email protected]

Mark Waldrop USGS 345 Middlefield Road 94025 Menlo Park, CA, USA [email protected] Sizhong Yang GFZ German Research Center for Geosciences, Helmholtz Centre Potsdam Section Geomicrobiology 14473 Potsdam, Germany [email protected] AND State Key Laboratory of Frozen Soil Engineering, Northwest Institute of Eco-Environment and Resources, Chinese Academy of Sciences Lanzhou 730000, China

Susanne Liebner and Christian Knoblauch

1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback 1.1 The permafrost carbon feedback Permafrost is perennially frozen ground, which characterizes soils of the high northern latitudes. Above the permafrost, a shallow surface soil layer, the so-called active layer, thaws during the short summer season. Permafrost soils store between 1100 and 1500 Gt of organic carbon (OC) [1]. While covering only 15% of the exposed land surface of the northern hemisphere, permafrost soils represent approximately 50% of the global soil carbon storage and contain about twice as much carbon as currently in the atmosphere. The sources of OC in the permafrost are fossil plant remains that have not been mineralized in the active layer. Permafrost affected soils receive fresh organic matter from the current vegetation, which in the tundra are mainly mosses, lichens, grasses, and dwarf-shrubs. Hence, the highest carbon concentrations are found in surface soils even though substantial amounts are also preserved in the widespread so-called Late Pleistocene Ice Complex or Yedoma permafrost deposits of northeastern Siberia, Alaska, and northwestern Canada. These 12 to 80 kyr old deposits, which are several tens of meters thick, contain, on average, 1.2 to 4.8% OC derived from plant debris and grass roots of the tundra steppe [2, 3]. Warming shows pronounced effects on near-surface permafrost (Fig. 1.1), and future climate change is expected to reduce its current extent substantially through active layer deepening. In detail, current model simulations indicate that about 10–20% of near-surface permafrost was lost due to surface temperature increase between 1960 and 2000 [4], and between 10% and 65% of near-surface permafrost is predicted to disappear until the year 2100 [5, 6]. The amplification of surface warming due to the mobilization of carbon from thawing permafrost and the associated release of the greenhouse gases (GHGs) carbon dioxide (CO2) and methane (CH4) is called the permafrost carbon feedback (PCF) [7]. The PCF strength and timing are estimated to add carbon emissions in the range of 5.7 ± 4% of the anthropogenic carbon release by 2100 and between 0.05 and 0.39°C by 2300 [8, 9]. It has been suggested that accounting for additional GHG emissions from this mobilized soil OC (SOC) could reduce permissible anthropogenic fossil fuel emission budgets by up to 25% to limit warming to 2°C above pre-industrial levels by 2100 [10, 11]. However, neither global climate projections nor climate policies adequately account for emissions from thawing permafrost [12]. The large OC stocks in tundra ecosystems imply that the Arctic has been a sink for atmospheric CO2 over long periods of time. Global warming has the potential to change permafrost landscapes from a carbon sink into a carbon source, but the https://doi.org/10.1515/9783110497083-001

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 1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback

Fig. 1.1: Degrading permafrost landscape in northern Scandinavia. The picture shows the collapse of a palsa mound, which already created a thaw pond surrounding it (on the right).

response of the Arctic carbon cycle to current changes in permafrost areas is still uncertain [13, 14]. This is also due to the lack of observational data from the vast landscapes in northern Russia that are characterized by cold and deep permafrost. The strong seasonal variability of carbon fluxes in Arctic ecosystems [15–17] and the lack of sufficient studies on multi-annual GHG budgets still prevent a clear picture of the current source-sink term of the Arctic tundra. An increasing release of CO2 from ­permafrost-affected landscapes, e.g. due to deeper permafrost thawing, has been shown [17]. However, warming of Arctic ecosystems may also increase the uptake of atmospheric CO2 due to a longer growing season [18]. In addition to the consequences of permafrost warming on soil temperature, its effect on soil hydrology will be of imminent importance because increasing soil moisture favors the formation of anaerobic, which means oxygen-free, soil conditions and, thereby, the production of CH4 [19], a GHG 28 to 45 times more potent than CO2 on a 100-year time horizon [20, 21]. Since permafrost impedes soil drainage, water-saturated soils with anaerobic conditions are widespread and contribute substantially to atmospheric CH4 concentrations [22, 23]. Despite low ­temperatures, which reduce organic matter decomposition rates in permafrost-affected soils, the Arctic tundra is still a substantial CH4 source, releasing about 19 Tg CH4 yr−1 or, almost 10% of the total global natural wetland emissions [24].

1.2 Carbon decomposition and GHG release from thawing permafrost 

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The biological production of methane is due to methanogenic archaea (methanogens), which belong to the phylum Euryarchaeota. They require anoxic conditions and low redox potentials to be active. Methanogenic taxa were also proposed outside the phylum of Euryarchaeota [25, 26], but their methanogenic physiology remains to be proven. Methanogenic archaea use H2/CO2 (hydrogenotrophic) and acetate (acetoclastic) as main substrates. While most methanogens are capable of hydrogenotrophic methanogenesis, acetoclastic methanogenesis is restricted to the taxa Methanosaeta (also known as Methanothrix; obligate acetogenic) and Methanosarcina (facultative acetogenic) [27]. Before being released into the atmosphere CH4 produced in anoxic soils and sediments has to migrate through oxygen-containing environments such as aerobic soil horizons, the water column of lakes, or air-filled tissues in vascular plants. At the boundary between anoxic and oxic soil conditions where both CH4, migrating from anoxic soil layers below, and oxygen, diffusing from the atmosphere into the soil, are present, aerobic CH4-oxidizing bacteria (methanotrophs) thrive [28, 29] (see also section 1.3). Even at low oxygen concentrations, methanotrophic bacteria oxidize CH4 to CO2 [30], thereby reducing CH4 emissions into the atmosphere.

1.2 Carbon decomposition and GHG release from thawing permafrost To improve projections on how much and how fast OC from permafrost can be released as trace gases from the northern tundra, it is necessary to better understand the degradability of organic matter in the different permafrost deposits that are projected to thaw in the near future. Data on permafrost organic matter degradability, i.e. how fast and to what extent permafrost carbon can be mineralized to CO2 and CH4, are scarce and limited to studies from northern Kolyma and the Lena Delta regions in northeastern Siberia [3, 31–33] and from continuous and discontinuous permafrost in Alaska [33, 34]. Meaningful predictions on the production and liberation of GHG from thawing permafrost need a quantitative understanding on how fast permafrost organic matter (OM) is decomposed into CO2 and CH4 after thaw under in situ ­conditions. The microbial decomposition rates of OM in permafrost soils depend on a variety of environmental parameters such as temperature [35, 36], nutrients [37], the availability of fresh organic matter [35, 38], the redox status [39], and the microbial community composition [40, 41]. Several studies quantified the effect of rising temperatures and thawing permafrost on CO2 and CH4 fluxes [42–45]. However, deriving permafrost OM decomposition rates from surface fluxes is challenging since these always integrate GHG fluxes from the thawed permafrost and the overlying active layer, which generally may not be partitioned. Few studies used radiocarbon analysis of released CO2 and CH4 and quantified the contribution of permafrost carbon to overall fluxes

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 1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback

using the lower δ14C signatures of permafrost OM [17, 46]. But a wider application of such studies is hampered by their high costs and the limitations of sample preparation in remote areas such as northern Siberia. The chemical characterization of permafrost organic matter composition indicates its rapid decomposability after thaw [47, 48]. Radiocarbon analysis of different OM fractions showed that thawed OM is only weakly protected by organo-mineral interactions [49] and that most OM in the active layer and the permafrost is available for microbial decomposition [50]. However, it is still unclear how far chemical analyses of organic matter relate to its decomposability. Therefore, incubation experiments have been conducted that enable the direct quantification of GHG production from permafrost OM under defined conditions. Although laboratory incubations are highly artificial since they isolate the incubated soil material from its natural environmental conditions, they enable studying the impact of selected environmental parameters on OM decomposition. The first incubation studies estimated a decomposition of more than 90% of thawing permafrost organic matter into CO2 over four decades by extrapolating aerobic carbon decomposition rates of Yedoma permafrost samples [32]. This is substantially more than the 15% over 100 years simulated with a calibrated carbon decomposition model [31]. The lower estimates do not result from lower carbon decomposition rates at the end of permafrost incubations feeding into the model, which were surprisingly similar in the Lena River Delta permafrost (0.1–2.8 µg C g−1 dw d−1) and the Kolyma Yedoma (0.3–1.6 µg C g−1 dw d−1) [31, 32]. Rather, these estimates derive from considering different carbon pools in a model representing the non-linear carbon decomposition. Due to the high contribution of the stable carbon pool with a turnover time of more than a century, carbon degradation decreases significantly in the time frame considered. Furthermore, the projections of CO2 production for the next 100 years consider that the active layer soils in permafrost landscapes are completely frozen during most of the year. Freezing substantially reduces, or even ceases, the turnover of organic matter [51] and is the primary driver for the high organic matter accumulation in permafrost. In a meta-analysis, a three pool carbon decomposition model was calibrated with data from aerobic incubations of soil samples from 23 circum-arctic sampling sites [52]. The samples were grouped into three classes: organic (>20% OC), shallow mineral (1  m depth). The OM in all three groups was dominated by the passive pool, which had a carbon turnover time of more than 2500  years and a contribution of more than 90% in the shallow and deep mineral material. The calibrated three-pool model simulated an aerobic mineralization of 22% and 3% of initial OC in the shallow and deep mineral soils, respectively, over an incubation period of 50 years. Estimating the long-term mineralization of organic matter in thawing ­permafrost requires knowledge on how much carbon will thaw and under which conditions, e.g. in water-saturated anaerobic or in well-drained aerobic soils. In particular, anoxic conditions, widespread in permafrost landscapes, substantially reduce

1.2 Carbon decomposition and GHG release from thawing permafrost 

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CO2 production and the mineralization rate of soil organic matter as shown previously [33, 53, 54]. Long-term permafrost incubations indicate a significantly smaller labile carbon pool and a 16-fold higher turnover time of the stable carbon pool in the absence of oxygen [41]. Anaerobic conditions from permafrost reduced the projected CO2 release in 100 years by a factor of 8 (15.1% initial C mineralization aerobically vs. 1.8% anaerobically) (Fig. 1.2). In the arctic Siberian Lena River Delta, a land cover classification identified more than 40% of the land area as a substantial source of CH4, which indicates water saturated, anaerobic conditions at deep in the active layer [55]. A deepening of the active layer due to permafrost thaw in these landscapes would result in mainly anaerobic degradation of carbon in the former permafrost and, hence, a substantially lower CO2 formation than estimated for aerobic conditions. Low initial methane production rates (1 m) mineral permafrost material, which was simulated for 50 incubation years at a permanent temperature of 5°C [52]. The estimate in the study of Schädel et al. [52] is likely due to the calibration of a three-pool OC decomposition model with a very high turnover time of the inert carbon pool in this particular model. However, Elberling et al. [65] calibrated a similar model with incubation data from near-surface permafrost samples from three sampling sites and simulated a loss of 58 to 77% of initial OC in 50 incubation years at two dry sites and still 13% at one wet site [65]. A 90-year simulation until 2100 revealed substantially lower organic matter decomposition under anoxic conditions, with a CO2 formation of 17 ± 9.3 g CO2-C kgC−1 and a CH4 production of 22 ± 13 g CH4-C kgC−1, which represents 1.7% and 2.2% of initial permafrost OC, respectively [41]. Total OC decomposition under anoxic conditions was, on average, three times lower than under oxic conditions. When comparing the GHG production under oxic and anoxic conditions based on CO2-C equivalents (CO2-Ce), i.e. by considering a global warming potential (GWP) of 28 for CH4 (weight corrected), the CO2-Ce production under anoxic conditions sums up to 241 ± 138 g CO2-Ce kg of initial permafrost OC, which is, on average, 2.4 ± 1.2 times higher than under oxic conditions [41]. Hence, these observation-calibrated long-term estimates of anaerobic CO2 and CH4 production contradict recent studies reporting a minor importance of methane production and a lower GHG production after permafrost thaw under water-saturated, anoxic conditions [33, 59, 63, 64]. Using a GWP of 28 for CH4 [21] is likely a conservative estimate of its radiative forcing in comparison to CO2 since recent metrics based on more realistic assumptions of CO2 and CH4 ecosystem fluxes indicate a higher sustained-flux GWP of 45 [20]. Decadal-scale (60  yr) in situ CH4 production rates from thawing permafrost surrounding arctic thermokarst lakes were estimated to be about 0.50 g CH4-C kgC−1 yr−1 [66]. However, in situ CH4 fluxes from thawing permafrost in two boreal peatlands (0.02–0.04  g  CH4‑C  kgC−1  yr−1) [46] were below the range of the long-term estimates (0.09–0.46 g CH4-C kgC−1 yr−1) from laboratory incubations [41]. These low in situ CH4 fluxes might be due to a low decomposability of the woody remains and Sphagnum peat

1.3 Microbial consumption of methane in permafrost environments 

 7

in the studied peatlands [46] and to CH4 oxidation during CH4 transport from the anoxic peat into the atmosphere [29, 67]. In summary, the large variability in predictions on OC turnover indicates a high heterogeneity of permafrost organic matter decomposability. Depending on the future atmospheric GHG concentration pathways (RCP4.5 or RCP8.5), a total amount of 67 ± 70 Pg and 243 ± 127 Pg of current permafrost OC was estimated to thaw and become available for microbial decomposition [6, 41]. These numbers are lower than recently reported results [68], most probably since these simulations used updated thaw projections [69]. In this scenario, less than one third the amount of OC will thaw under water-saturated relative to non-saturated conditions and the GHG production in terms of carbon will be substantially higher in non-­ saturated soils (2.6 ± 3.2 Pg–9.5 ± 7.0 Pg OC into CO2) than in saturated soils (0.4 ± 0.6 Pg–1.4 ± 1.4 Pg OC into CO2 and CH4). These numbers account only for additional OC mineralization from thawing permafrost and exclude the surface active layer, which thaws every summer, receives fresh organic matter from the surface vegetation, and generally exhibits higher carbon decomposition rates than deeper permafrost layers [46, 52]. Hence, these estimates of permafrost carbon release are substantially lower than the range recently reported from spatially explicit approaches (21–174 Pg carbon) considering the whole soil column including the current active layer [7, 69]. However, due to the equal contribution of CO2 and CH4 production to total GHG formation and due to the higher GWP of CH4, water-saturated soils are predicted to contribute almost equally to the overall GHG production from thawing permafrost if expressed as CO2-C equivalents at a pan-Arctic scale (2.4 ± 3.7 Pg–8.9 ± 8.8 Pg CO2-C equivalents) [41]. The large uncertainties in those estimates reflect the simplicity of the calculations, which show on the one hand the importance of anoxic decomposition pathways for GHG production in permafrost regions and on the other hand the need for further spatially explicit approaches.

1.3 Microbial consumption of methane in permafrost environments GHG production in thawing permafrost cannot be directly transcribed into GHG emissions to the atmosphere since a variable fraction of CH4 produced in thawing permafrost will be oxidized by microorganisms to CO2 when passing oxic soil or sediment layers [29, 67]. Methane oxidizing bacteria and archaea, so-called methanotrophs, may oxidize more than 90% of the produced CH4 before it reaches the atmosphere [29, 67, 70, 71]. In this context, the vegetation composition plays a crucial role since many vascular wetland plants channel CH4 from its production zone into the atmosphere, thereby circumventing its oxidation [29]. Also, depending on other environmental conditions than vegetation, the extent of methane oxidation may vary largely. Aerobic methanotrophs belong to different groups in the α-Proteobacteria (type II) and the γ-Proteobacteria (type I) and use CH4 as their sole energy and carbon

8 

 1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback

source. In environmental samples, they can be identified by their specific unsaturated phospholipid fatty acids (PLFAs) [72, 73]. In particular, the PLFAs 16:1ω7c/8c and 18:1ω7c/8c were shown to be produced specifically by methanotrophic bacteria [72, 74]. Besides their signature fatty acids, type I and type II methanotrophs also exhibit an overall different PLFA pattern, with 16:1 PLFAs predominating in type I and 18:1 PLFAs in type II [72, 74, 75], with only a few exceptions among the acido-, thermo-, and halophilic methanotrophs [76–78]. Combining biomarker studies with stable isotope labeling enables the identification of active microbial groups in environmental samples, and this approach has been repeatedly applied to study CH4 oxidizing bacteria [79–81]. The fraction of CH4 that is oxidized by methanotrophs before being emitted into the atmosphere strongly depends on the transport pathways of CH4. The main pathways are ebullition, molecular diffusion, and plant-mediated transport [82]. Ebullition is the release of gas bubbles through the water into the atmosphere. Due to the low solubility of CH4, the formation of gas bubbles is widespread in environments with active methanogenesis [83]. Since gas bubbles rapidly rise through the water column, CH4 oxidation has a minor impact on ebullitive fluxes [84]. Diffusive CH4 transport follows the concentration gradient between the production zone and the atmosphere. Since CH4 diffusion through water is about 104 times slower than through air, diffusive transport of dissolved CH4 through the water phase leaves elevated contact time between dissolved CH4 and methanotrophic bacteria [85]. Plant-mediated CH4 transport is conducted by many emergent macrophytes through air-filled aerenchyma inside their roots, stems, and leaves [86]. Oxygen is transported into the roots, and CO2 and CH4 diffuse through the air phase into the atmosphere [87]. Vascular plants further impact microbial processes in the rhizosphere, e.g. by supplying CH4 oxidizers with molecular oxygen leaking from the roots [67, 88] or by releasing root exudates and border cells fueling microbial activity including CH4 production [89, 90]. Hence, vascular plants have a positive effect on the processes of methane production (release of root exudates), methane oxidation (release of oxygen into the anoxic bulk soil), and methane emission (rapid transport through aerenchyma). Besides vascular plants, also submerged mosses were reported to interact with methanotrophic bacteria likely by providing oxygen for microbial methane oxidation [91, 92] (see also section 1.5). But while numerous studies highlight the importance of vascular plants in the CH4 cycle of wetlands, the quantitative impact of mosses on CH4 oxidation and CH4 emission is not well understood [80, 93]. Methane oxidation in soils is further strongly affected by environmental parameters such as water table height or temperature.

1.4 Temperature response of methane oxidizers Microbial methane production and oxidation may respond differently to a change in temperature [94]. Furthermore, the effect of rising temperatures on microbial activities

1.4 Temperature response of methane oxidizers 

 9

600

80

A

CH4 oxidation (nmol gdw–1 h–1)

CH4 oxidation (nmol gdw–1 h–1)

may be modified by the limitation of substrate, which causes the apparent temperature sensitivity of the microbial community under field conditions being smaller than expected [95]. The strong seasonality in the Arctic means that temperatures in the active layer may fluctuate between about −30°C in winter and +10°C in summer [96]. Our knowledge of how methanotrophs cope with these extreme temperatures is very limited. Which organisms are active at low in situ temperatures and how these communities will respond to the rising temperatures predicted for the future is unclear. Rising temperatures in the Arctic tundra will not only cause permafrost thaw at the bottom of the active layer but also directly impact the active microbial communities, causing increasing metabolic rates but potentially also changes in the active community composition [73, 97]. The temperature optima of methane production and oxidation are generally above in situ temperatures, which causes an increase in activities with rising temperatures [81, 94]. Maximum CH4-oxidizing activities in northern Siberian soils of the Lena Delta (Fig. 1.3) were found at temperatures far above in situ temperatures of about 5°C during the summer thaw period [81]. The temperature of maximum CH4 oxidation rates (Topt) in the active layer and in the permafrost of Samoylov was at around 28°C, while Topt was substantially lower (22°C) on Mamontov Klyk. All samples showed still activity at 0°C, although this represented only 5.3 to 7.6% of the activity at the temperature optimum in the Samoylov active layer soils, but 31.8% and 39.6% in Mamontov Klyk and the Samoylov permafrost, respectively [81]. Previous studies on the temperature response of methane oxidation in different ecosystems between the subarctic and the tropics have indicated only a weak correlation between Topt and environmental in situ temperatures. The Topt for methane oxidation increases from about 20°C to 25°C in boreal bogs and up to 36°C in subtropical paddy soil [94, 98–101]. Methane oxidation in water-saturated subarctic and arctic soils and sediments have consistently shown a clear increase in CH4 oxidation with rising temperatures and a Topt

500

B

60

400

50

40

300

30

200

20

100 0

70

0

10

20 30 40 Temperature (°C)

50

10 0

0

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20 30 40 Temperature (°C)

50

Fig. 1.3: Temperature response of methane oxidation in north Siberian surface soil of Mamontovy Klyk (A) and Samoylov (B). Error bars indicate the standard deviations of quadruplicate incubations. Data from [81].

10 

 1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback

above 20°C [94, 97, 102, 103]. Only substrate limitation, i.e. low CH4 concentrations, might prevent increasing rates at rising temperatures in Arctic lakes [102, 104]. The only indication of a low Topt for methane oxidation was reported by Liebner et al. in a study of methane oxidation in a polygonal center on Samoylov using 14C-labeled CH4 [105]. The CH4-oxidizing community in the upper part of the water-saturated profile, where cell numbers of methanotrophs were highest, showed either temperature optima above 21°C or did not show a clear response to temperatures between 0 and 27°C. Only in one sample close to the permafrost table, where cell numbers of methanotrophs were lowest, maximum CH4 oxidation rates were found at 4°C. If such a low Topt is more widespread in Siberian tundra soils and characteristic for the in situ CH4-oxidizing communities needs further investigations.

1.5 Interaction of methane oxidizers and mosses A typical landscape of the Arctic zone is the polygonal tundra (Fig. 1.4 left) [106], with a characteristic surface micro-topography of elevated rims and depressed polygon centers (low-center polygons) [107]. While the soils of the elevated rims are non-saturated in their upper horizons, the water table position in the depressed polygon centers may be at or above the soil surface [108, 109], which causes an extensive abundance of water-saturated, anoxic soils in the tundra. The vascular vegetation of these water-saturated and inundated soils is dominated by sedges and grasses, such as Carex spp., Eriophorum spp., Arctophila fulva, Arctagrostis latifolia, and Dupontia fisheri [29, 109–112], while the moss stratum of the wet and non-acidic tundra in northern Siberia and Alaska is dominated by Bryopsida species like Scorpidium scorpioides, Meesia triquetra, and Drepanocladus revolvens [108–111, 113]. These moss species also grow submerged as dense layers at the bottom of Arctic lakes [114] and are widespread in the polygonal centers on Samoylov [29, 80] (Fig. 1.4 right).

Fig. 1.4: Polygonal tundra landscape in the Siberian Lena Delta with polygonal ponds occurring in the centers of many polygons (left). Submerged brown mosses are widely occurring in these ponds (right).

1.5 Interaction of methane oxidizers and mosses 

 11

A close interaction between methane-oxidizing bacteria and Sphagnum mosses has been shown for acidic peatlands [92, 115, 116] in which Sphagnum mosses may benefit from CO2 provided by the microbial oxidation of CH4 [91]. However, whether a similar interaction is also established in neutral fen-type soils of the polygonal tundra and what effect this might have on the CH4 fluxes to the atmosphere were unclear, and so was the potential benefit of such an interaction for the methane-oxidizing bacteria. Since maximum CH4 oxidation rates were measured in the living moss layers of inundated soils on Samoylov, it was hypothesized that submerged mosses and CH4 oxidizers form a symbiosis similar to the one demonstrated for Sphagnum mosses [92]. In such a symbiosis, the oxygen produced by the brown mosses via photosynthesis would be used for microbial oxidation of CH4 to CO2, which in turn could be used as carbon substrate for moss photosynthesis. In this case, CH4-carbon should be incorporated into the moss biomass and the activity of CH4 oxidation should depend on moss ­photosynthesis. To test this, the carbon stable isotope signatures of submerged and non-submerged mosses from a variety of sampling sites on the Siberian Island Samoylov were measured [80]. Samples were collected from different parts of the aquatic mosses Scorpidium scorpioides and Meesia triquetra from three different polygonal centers. At the same polygons, the terrestrial mosses Hylocomium splendens, Aulacomnium turgidum, A. palustre, and Drepanocladus cossonii were sampled in a transect from the water-saturated centers to the top of the dryer rims. Since dissolved CH4 has a substantially lower δ13C signature than dissolved or atmospheric CO2, moss biomass should be depleted in 13C if they fix substantial amounts of CO2 from CH4 oxidation. The δ13C signatures of submerged S. scorpioides and M. triquetra biomass ranged between −29.6 and −33.9‰ VPDB (mean −31.5 ± 1.3‰ VPDB) and were significantly lower than the carbon stable isotope signatures of the terrestrial mosses that ranged between −24.9 and −28.6‰ VPDB. The carbon stable isotope signatures of leaves and stems of S. scorpioides and M. triquetra reaching above the water surface into the free air (−25.6 to −27.0‰ VPDB) were similar to those of terrestrial mosses but significantly higher than those of the submerged growing biomass of the same species. Hence, the relatively low carbon stable isotope signatures in submerged growing mosses are not species-specific but related to submerged growth. It has been demonstrated that the water table is the key determinant for CH4 oxidation also for Sphagnum-associated CH4 oxidation and that highest rates were found when the water table is close to the moss surface [115, 116]. Since aquatic mosses are not able to use bicarbonate, they require dissolved CO2 [117–119]. Because CO2 diffuses 10,000 times slower through water than through air, the growth of aquatic mosses may be limited by the dissolved CO2 supply [120]. Under these conditions, CO2 deriving from subaquatic CH4 oxidation may provide additional carbon substrate for moss photosynthesis. The significantly lower δ13C values of the submerged moss biomass in comparison to emersed living and terrestrial mosses indicate that this is indeed the case under in situ conditions. To evaluate if the decreased carbon stable isotope signature in the brown mosses originates from the fixation of CO2 from CH4 oxidation,

12 

 1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback

potential CH4 oxidation rates were measured in mosses from five different polygonal centers with the water table either above the soil surface containing submerged living mosses (Meesia triquetra, Scorpidium scorpioides) or with a water table below the surface and emersed living mosses (Drepanocladus cossonii) [80]. The methane oxidation rates of the submerged mosses ranged between 162 ± 7.5 and 390 ± 24.4  nmol  g−1  d−1 and were hence similar as in brown moss layers of two neighboring polygonal centers but substantially higher than in mineral soils on Samoylov (1.9–7.0 nmol g−1 h−1) [105, 121] and most of previously reported CH4 oxidation rates in northern water-saturated soils and sediments [28, 100, 104]. These high CH4 oxidation rates were accompanied by elevated CH4 concentrations at the bottom of the moss layer. In contrast, at least 10-fold lower potential CH4 oxidation rates (15.1 ± 3.7 nmol g−1 d−1) were measured in the mosses (D. cossonii) of the polygonal center with a water table below the surface, which showed also low CH4 concentrations (0.22 µM). The high CH4 oxidation rates in the submerged mosses and the relatively low carbon stable isotope signatures are a clear indication that a close interaction between the CH4 oxidizing bacteria and the mosses enables the transfer of CO2 from CH4 oxidation into the mosses. While the benefit of the interaction with CH4 oxidizers is clearly established for the mosses, which are supplied with CO2 from CH4 oxidation, the benefit for the CH4-oxidizing bacteria is less clear, although increased oxygen supply from photosynthesis and stabilization of bacterial cells in the oxic–anoxic interface has been suggested [91, 115]. The proof for a positive effect of moss photosynthesis on the activity of CH4 oxidizers was established by quantifying CH4 fluxes from the surface of the water-saturated soils with an opaque and a transparent chamber [80]. Since light penetrates through the transparent chamber, summerged living mosses may photosynthesize and produce oxygen. In contrast, no light reaches the surface unter the opaque chambers, photosynthesis is inhibited and no oxygen is produced. In the case that CH4 oxidation benefits from oxygen production through moss photosynthesis, CH4 fluxes should be higher under the opaque chamber when oxygen production is prevented. Plots without vascular plants were selected to prevent plant-mediated CH4 transport. Methane fluxes into the transparent chamber were as low as −1.7 ± 11.3 mg CH4 m−2 d−1 (mean ± SD, n = 8), while with the opaque chamber, CH4 fluxes were significantly higher and always positive, with a mean CH4 flux of 21.6 ± 9.1 mg CH4 m−2 d−1 [80]. At the similar flooded polygonal centers without vascular plants on Samoylov, diffusive CH4 fluxes ranged between 0.64 ± 0.73 mg CH4 m−2 d−1 and 3.1 ± 1.7 mg CH4 m−2 d−1 [29], which is well in the range of the fluxes into the transparent chamber but much less than CH4 fluxes in the absence of light. Hence, obviously, CH4 oxidation is higher under light conditions, which is most likely caused by the oxygen production through photosynthetic mosses. In summary, the interaction between photosynthetic mosses and bacteria, in general (Fig. 1.5), and with methane-oxidizing bacteria, in particular, is a mutualistic association with a benefit for both partners. A large fraction of the CO2 demand

1.6 Methane oxidation in submarine permafrost 

 13

Fig. 1.5: Association between mosses and bacteria. The pictures show fluorescence micrographs of Sphagnum cells colonized by bacteria.

of the submerged living mosses is provided by microbial CH4 oxidation. In turn, methane-oxidizing bacteria are supplied with oxygen from moss photosynthesis below the water table, which supports substantially higher CH4 oxidation rates than in water-saturated mineral soils without a moss cover. The combined activity of the moss-CH4 oxidizer association substantially reduces CH4 fluxes from Arctic polygonal tundra into the atmosphere. Even uptake of atmospheric CH4 was observed, despite high CH4 concentrations in the water-saturated soils of the polygonal centers.

1.6 Methane oxidation in submarine permafrost Subsea permafrost developed after low-lying terrestrial permafrost areas were flooded by sea water during the Holocene marine transgression and by coastal retreat since the Last Glacial Maximum. The main distribution of subsea permafrost is in the Laptev and the East Siberian Sea, where the shelf extends hundreds of kilometers offshore [122, 123]. After flooding, submarine permafrost is exposed to marine waters with temperatures ranging between about +2 and −1.5°C [124], causing a thawing of submarine permafrost with a current rate between 5 and 14 cm yr−1 [125, 126]. Like terrestrial permafrost, subsea permafrost contains organic matter that may be decomposed to CO2 and CH4 but also CH4 that may be release after thaw. Furthermore, subsea permafrost was supposed to form a gas-impermeable cover above thermogenic CH4 and CH4 hydrates in deeper sediment layers of the Arctic Ocean [127, 128]. Accelerated rates of submarine permafrost thaw were suggested to cause accelerated CH4 release into the atmosphere [124, 129, 130]. However, a large range of CH4 fluxes above subsea permafrost was reported from 3–4 mg m−2 d−1 when using direct observations [131] to 100–630  mg m−2 d−1 when upscaling CH4 release from estimated gas bubble fluxes [129]. These largely differing fluxes, which were derived from measurements in the same ocean area, underline the high uncertainties on CH4 fluxes from submarine

14 

 1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback

permafrost. These uncertainties are mainly caused by a very limited number of direct observations, in particular those covering the properties of the subsea permafrost itself. In April 2005, an international Arctic expedition drilled one core (C2) into the seabed about 12  km offshore in the western Laptev Sea at Cape Mamontov Klyk (73.71N, 117.17E) [132]. The core was drilled below a water depth of 6 m to a total depth of 71 m below the sea floor (mbsf). Frozen sediments, which were deposited under terrestrial conditions, were encountered between 34.3 and 57.4 mbsf [133, 134]. At deeper depths, sediments were unfrozen again. The temperature in the bore hole ranged between −1.0 and −1.7°C. A second core (BK2) was drilled in 2012 about 0.75 km off the eastern coast of Buor Khaya Bay in the central Laptev Sea (71.47N, 132.08E), at a water depth of 4.2 m [135]. The temperature in the bore hole ranged between −1.6 and −1.7°C. Ice-bonded permafrost was encountered in a depth of 24.8  mbsf. Both cores were analyzed concerning their stratigraphy, salinity, and ion concentrations [125, 134]. Furthermore, CH4 concentrations, carbon stable isotope signatures, and OC contents were measured to evaluate the turnover of organic matter and CH4 in these subsea permafrost sediments. Both cores showed substantial CH4 concentrations in the ice-bonded permafrost layers either in specific layers (Mamontov Klyk) or throughout, as can be seen in Fig. 1.6. At the same time, the isotopic signatures of CH4-carbon suggest that CH4 is oxidized in particular at the submarine permafrost thaw boundary. In detail, the core of Mamontov Klyk was affected by the intrusion of sea water above about 40 mbsf, which could be shown by elevated salinity [134] and sulfate concentrations of up to 5 mM [136]. This would be supportive for the process of anaerobic CH4 oxidation. However, if the process is relevant at these sediment depths could not be finally resolved. CH4 stable isotope signatures in the frozen section of core C2 were partially higher than could be explained by methane production alone, which forms CH4 with δ13C values between about −50 and −110‰ [137]. Consistently, δ13C values of CH4 produced in Arctic lakes, ponds, and the anoxic bottom of water-saturated permafrost soils were between about −50 and −80‰ [29, 138–142]. Hence, δ 13C values of CH4 above −40‰, as found between 50 and 52 mbsf in the frozen core C2, are likely affected by microbial CH4 oxidation. Whether this CH4 was oxidized prior to permafrost formation or currently in situ at temperatures below 0°C remains unclear. The activity of methanogenic archaea has been demonstrated repeatedly in permafrost at sub-zero temperatures, similar to those in the frozen core [58, 143], which makes in situ CH4 production from the elevated organic matter concentrations at 52 mbsf and CH4 oxidation with sulfate as electron acceptor in the frozen sediment between 50 and 52 mbsf feasible. Based on isotopic fractionation factors for anaerobic CH4 oxidation [144] and CH4 diffusion [71], 72 to 86% of CH4 diffusing upward were calculated to be oxidized [136]. The Buor Khaya core could be split into three sections based on CH4, sulfate concentrations, and the carbon stable isotope signatures of CH4 [125]. In the frozen part below 25 mbsf, the highest CH4 concentrations, lowest stable carbon

1.6 Methane oxidation in submarine permafrost 

 15

Fig. 1.6: Methane concentrations (black lines) and its carbon stable isotopic signatures (red lines) in two submarine permafrost cores from the Siberian Arctic Shelf. The profiles suggest that CH4 oxidation occurs in the ice-bonded permafrost (Mamontov Klyk core) and at the permafrost thaw boundary (Buor Khaya core). The graph was kindly provided by Paul P. Overduin (Alfred Wegener Institute for Polar and Marine Research, Potsdam, Germany).

isotope signatures, and lowest sulfate concentrations could be found. This frozen core section comprised sediments formed under terrestrial conditions [125]. At the boundary between frozen and unfrozen sediments, CH4 concentrations dropped by one order of magnitude, δ13C values of CH4 increased by 30‰, and sulfate concentrations raised from 0.2 to 1.6  mM. The changing CH4 and sulfate concentrations indicate the anaerobic oxidation of CH4, which is strongly supported by the steep increase in CH4 δ13C values. In the following 12  m of unfrozen sediment, CH4 concentrations, its carbon stable isotope signatures, and sulfate concentrations stayed relatively constant. CH4 concentrations in the frozen part of the Buor

16 

 1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback

Khaya core were substantially higher than in the frozen section of the Mamontov Klyk core, but the δ13C signatures of CH4 were higher and in the range typical for microbial methane production. Furthermore, the concentrations of potential electron acceptors for anaerobic methane oxidation (nitrate, sulfate, iron, and manganese) were all below 0.5 mM in the frozen section [136]. Hence, in contrast to the Mamontov Klyk core, the geochemical data do not indicate anaerobic CH4 oxidation in the frozen part of the Buor Khaya core. Furthermore, no anaerobic methane-oxidizing archaea were detected in this section of the core, while they were present in high abundance at the interface between frozen and unfrozen core sections (24.0–24.7 mbsf) [136]. Using the isotope approach to quantify CH4 oxidation, between 79 and 100% of CH4 released from permafrost thaw is oxidized at the thaw front [136]. Combining the release of CH4 through submarine permafrost thaw with these fractions of CH4 oxidized would yield potential anaerobic CH4 oxidation rates of 1.7–2.1 nmol cm−3 d−1 ± 0.9–1.2 nmol cm−3 d−1 [126, 137], which are similar to anaerobic CH4 oxidation rates in ocean margins but higher than in the subsurface or marine anoxic water columns [145]. All together, those first in situ studies on the fate of CH4 in submarine permafrost provide evidence that CH4 oxidation commonly occurs when subaquatic permafrost thaws under marine conditions. Recent work suggests that similar processes occur in thermokarst lakes with connection to the sea, so-called thermokarst lagoons (Fig. 1.7). All these studies give no evidence for an elevated CH4 release from the thawing of submarine permafrost or thermokarst lagoon sediments themselves as previously suggested [130] but rather indicate an efficient oxidation of the CH4 released from permafrost thaw in the sea bottom. The observed CH4 fluxes from the Siberian shelf seas [131] rather seem to be related to different processes such as ebullition of old CH4 from below the permafrost or from gas hydrates through discontinuities in the permafrost [146, 147]. In particular ebullition fluxes are unaffected by oxidation processes [148] and might contribute a substantial fraction to overall CH4 fluxes from the Arctic shelf oceans [129, 147].

1.7 Summary Permafrost regions store between 1100 and 1500 Gt of OC. They cover only 15% of the exposed land surface of the northern hemisphere but represent approximately 50% of the global soil carbon storage and contain about twice as much carbon as is currently in the atmosphere. About 10–20% of near-surface permafrost was lost due to surface temperature increase between 1960 and 2000, and between 10% and 65% of near-surface permafrost is predicted to disappear until 2100. The amplification of surface warming due to the mobilization of carbon from thawing permafrost and the associated release of the GHGs CO2 and CH4 is called the PCF. Accounting for

1.7 Summary 

 17

Fig. 1.7: Drilling into a thermokarst lagoon on the Bykovsky Peninsula in the Siberian Arctic (top). Profiles of CH4, sulfate, and δ13C-CH4 indicating methane oxidation with sulfate at the marinefreshwater transition at ~2.3 m sediment depth of the same lagoon (bottom).

additional GHG emissions from this mobilized SOC could reduce permissible anthropogenic fossil fuel emission budgets by up to 25% to limit warming to 2°C above pre-industrial levels by 2100. Predictions on OC turnover in thawing permafrost show large variabilities, estimating between more than 90% of organic matter turning into CO2 over four decades and only 15% or even less of organic matter being mobilized over 100 years. This indicates a high heterogeneity of permafrost organic matter decomposability and stresses

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 1 Microbial carbon dioxide and methane turnover and the permafrost carbon feedback

the large uncertainties associated with PCF projections. However, it was shown that thawing organic matter is only weakly protected and that most of it is available for microbial decomposition. Also, there are strong indications that the relevance of CH4 within the PCF is larger than long thought. GHG production in thawing permafrost cannot be directly transcribed into GHG emissions to the atmosphere since a variable fraction of CH4 produced in thawing permafrost will be oxidized by microorganisms to CO2 when passing oxic soil or sediment layers. Methane-oxidizing bacteria and archaea may oxidize more than 90% of the produced CH4 before it reaches the atmosphere but the extent of methane oxidation varies largely due to the complex interplay between hydrology, vegetation, and microbial adaptations to temperature and optimum CH4 concentrations. Mostly, maximum CH4-oxidizing activities were found at temperatures far above in situ temperatures in the active layer. Not only roots of vascular plants and aerobic soil layers but also submerged mosses have been identified as hot spot for microbial methane oxidation in Arctic tundra landscapes. The association between mosses and methane-oxidizing bacteria is a symbiosis in which oxygen is released by photosynthesis of brown mosses and Sphagnum, which then is likely used for microbial oxidation of CH4 to CO2. The latter, in turn, is used as carbon substrate for moss photosynthesis. Finally, the first in situ studies on the fate of CH4 in submarine permafrost and in coastal (thermokarst) lagoons give no evidence for an elevated CH4 release from the permafrost itself, as previously suggested. Rather, those studies point toward efficient oxidation of the CH4 released from permafrost thaw in the sea bottom. The reported, often very high, CH4 fluxes from the Siberian shelf areas rather seem to be related to different processes such as ebullition of old CH4 from below the permafrost but might also be a substantial overestimation [131, 149].

References [1] Hugelius G, Strauss J, Zubrzycki S, et al. Estimated stocks of circumpolar permafrost carbon with quantified uncertainty ranges and identified data gaps. Biogeosciences. 2014;11: 6573–6593. [2] Schirrmeister L, Grosse G, Wetterich S, et al. Fossil organic matter characteristics in permafrost deposits of the northeast Siberian Arctic. J Geophys Res. 2011;116:G00M2. [3] Zimov SA, Davydov SP, Zimova GM, et al. Permafrost carbon: stock and decomposability of a globally significant carbon pool. Geophys Res Lett. 2006;33:L20502. [4] Peng S, Ciais P, Krinner G, et al. Simulated high-latitude soil thermal dynamics during the past 4 decades. Cryosphere. 2016;10:179–92. [5] Beer C, Zimov N, Olofsson J, Porada P, Zimov S. Protection of permafrost soils from thawing by increasing herbivore density. Sci Rep. 2020;10:4170. [6] Koven CD, Schuur EAG, Schädel C, et al. A simplified, data-constrained approach to estimate the permafrost carbon-climate feedback. Philos Trans R Soc A. 2015;373:20140423. [7] Schuur EAG, McGuire AD, Schädel C, et al. Climate change and the permafrost carbon feedback. Nature. 2015;520:171–9.

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Rachel Mackelprang, Nesliah Tas and Mark Waldrop

2 Functional response of microbial communities to permafrost thaw 2.1 Introduction Permafrost soils underlie one quarter of the land surface in the northern hemisphere and contain vast stores of carbon (C) and nitrogen (N) [1]. Because permafrost is frozen by definition and has been for hundreds, tens of thousands, or even millions of years, it has not been thought of as a particularly hospitable environment for microbial activity. Over the past several decades, this understanding has shifted, and we now know that microbial communities in permafrost soils are active, in part because the temperature of permafrost is not too extreme. We also know that in the coming century, models predict that large areas of permafrost will thaw. Permafrost thaw imposes large changes on the carbon cycle and will potentially releasing 12 to 112 Pg of carbon back to the atmosphere by 2100 [2, 3], and much of this could be ancient permafrost carbon [4]. Increased Arctic temperatures have already initiated large carbon losses from thawing permafrost [5–7]. This is coincident with observations of increased greenhouse gas fluxes and lateral dissolved organic carbon (DOC) fluxes [8, 9] that outweigh potential increases in plant productivity [10, 11]. Much of this increase is due to the phase change from ice to water, warmer soils, and the presence of labile organics and nitrogen in the newly thawed environment [12, 13] flowing through the metabolic engine that is the soil microbial community. Although we have some understanding of the soil and ecosystem response to permafrost thaw, less is known about the functional response of the microbial community. Functional response can be defined as rates of respiration, methanogenesis, and methane oxidation, temperature and moisture responses, ATP production, and growth post-thaw. At the molecular level, much more detailed functional responses can be elucidated, including changes in functional gene abundance, transcript abundance, or protein abundance, as well as changes in the abundance of rRNA gene sequences that describe shifts in archaeal, bacterial, or fungal community composition. We review the literature for functional responses in both in lab and in situ conditions. At the molecular level, there is a body of work that evaluates community structural (i.e. 16S rRNA gene and ITS2 amplicon sequencing) response to thaw, and other studies that track functional response to surface soil warming. However, there are few that evaluate functional response of permafrost microbes to thaw. Other chapters cover amplicon-based studies and surface soil studies but do not capture the transition that occurs when permafrost microorganisms that are long-adapted to life at sub-zero temperatures cross the freezing point threshold. We therefore focus the molecular-based component of this chapter on the smaller body of work comparing functional-level https://doi.org/10.1515/9783110497083-002

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differences between intact and thawed permafrost. Understanding functional shifts in response to thaw will help elucidate the mechanisms controlling decomposition and greenhouse gas fluxes in this impending global ecosystem transformation.

2.2 Activity in permafrost Permafrost is a challenging environment for microbial communities. As temperatures drop below freezing, frozen water constricts activity to water films where diffusion of substrates, nutrients, and waste products is limited [14]. Fig. 2.1 shows an image

Fig. 2.1: Scanning electron microscopy (SEM) of permafrost samples from Interior Alaska. The embedded image shows fractures within the permafrost where the mix of organics, minerals, and microbes are concentrated as permafrost forms. The dark material is ice. The inset is a plane view of a fracture fill that has peeled away here showing “patterned ground” of ice in miniature. Microbial activity likely persists along these fractures in areas of presumably high salinity and sediment concentrated during ice formation. (Photo courtesy of Marjorie Schulz and Kristen Manies.)

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of permafrost soil spanned by cracks that have squeezed out their internal constituents. These constituents are likely a mixture of organics, clay minerals, and microbial cells. These cracks may also contain higher concentrations of salts than the surrounding ice, lowering the freezing point of water but increasing osmotic stress for microbial communities. Microbial communities in permafrost face additional challenges, including low thermal energy [15], high viscosity [16], protein denaturation, and loss of cell membrane fluidity impacting the transport of nutrients [17]. Microbial metabolic activity does not cease when soil temperatures drop below 0°C. Evidence for microbial activity in frozen permafrost soils comes through incubation experiments, aspartic acid (Asp) racemization measurements [18], growth assays, microscopy, and immunological approaches [19]. For incubation experiments, samples are placed into airtight jars and changes in headspace CO2 are monitored over time. Although most studies show microbial activity from 0 to −8°C or so, some studies have observed microbial activity down to as low as −39°C [14, 20–23]. D-Asp/LAsp ratios from intact cells in ancient permafrost (up to 1.1 Ma) suggests that they are viable and metabolically active [18]. To measure microbial activity at these very cold temperatures, 14C-labeled substrates had to be used [23, 24]. Methanotrophic activity, likely a lower detectable rate than respiration, has also been detected below freezing using 14CH4 in 1.8–3 million years old permafrost samples from Siberia [25]. Methane fluxes from permafrost, however, seem to show higher spatial variation, more complex dynamics, and surprisingly low fluxes compared to CO2 from the same soils [26, 27]. Several incubation studies show that there is a lag in CH4 production from frozen soils, that there is low or negligible CH4 from permafrost, and methane production from permafrost seems lower in more ancient permafrost deposits [26, 28–31]. This could indicate a very low abundance of methanogens in permafrost soils, low availability of substrates such as acetate or CO2 for methanogenesis, physical stress, or that other microbes are outcompeting methanogens for available substrates. However, the temperature response of CH4 may be higher than for CO2 [28, 32], indicating that the low CH4 flux in permafrost soils may have more to do with limited substrate availability than low biomass numbers. Respiration rates from frozen permafrost have been shown to be higher than frozen active layer soils incubated at the same temperature [29, 33, 34], indicating that permafrost microbial communities may have developed unique adaptations that have allowed them to be more active in frozen conditions. One such adaptation may be the higher production of ATP in microorganisms in frozen environments compared to warmer environments [35]. Under frozen conditions, microbes increase ATP production, possibly to maintain reaction rates as temperatures drop. This pattern could also occur if permafrost contains higher concentrations of water-soluble organic carbon (WSOC). In many cases, permafrost has higher concentrations of WSOC than seasonally thawed soils [9, 33], and WSOC is correlated to respiration rates from frozen soils [33]. Several internal molecular mechanisms allow microbial communities in permafrost to be more active than their active layer counterparts under frozen conditions and are described below but increased lability of organic matter within permafrost may add to this putative adaptive response.

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In the field, studies of microbial activity in frozen soils are mostly limited to the comparison of surface soils in summer and winter, not permafrost. But studies show that microbial activity continues through wintertime often under the insulating blanket of snow that keep surface soils just below freezing [36–40]. Dissimilatory iron reduction and methanogenesis are shown to be active part of permafrost microbial metabolism [41], but to our knowledge, there is as yet no in situ measurement of microbial activity in permafrost soils. Enrichment cultures and isolates from different permafrost habitats have given us additional knowledge about the types of microorganisms in permafrost soils and their strategies for survival [42]. The majority of permafrost isolates are from bacterial phyla, Firmicutes, Actinobacteria, Bacteroidetes, and Proteobacteria [43–45]. Many of these isolates can form spores [24, 46] or have high-GC content that may reduce DNA damage due to cold, high salt, and high radiation in permafrost [47]. For example, six bacterial isolates of the genus Carnobacterium (Firmicutes) obtained from ancient Siberian permafrost (6000–8000 years) grow at 0°C under low pressure (7 mbar) and CO2-enriched anoxic conditions [48]. Studies of permafrost isolates showed that cells lower their metabolic rates to maintain survival during low temperatures. [3H]leucine and [3H]thymidine incorporation studies done with the Siberian permafrost isolate, Psychrobacter cryohalolentis, showed that cells were able to synthesize DNA and proteins (1 to 16 proteins per day) at −15°C [35]. In another study, permafrost isolate Arthrobacter sp. 9-2 was shown to incorporate 14 C-ethanol during growth at temperatures as low as −17°C [34]. Similar incorporation rates were also detected for some other permafrost isolates and enriched co-cultures [49, 50], suggesting that even with slow metabolic rates it is possible to compensate for the DNA damage and protein denaturation that occurs under low temperatures [35].

2.3 Response to thaw Changes in the rate of CO2 flux as permafrost soils warm and thaw is an important relationship to understand and model in global change research [51]. The Q10 value of microbial respiration, defined as the change in rate per 10°C change in temperature, above freezing often is around 2.0 [52]. But the Q10 of frozen soils can be much larger, due to extracellular barriers to diffusion and osmotic stress in the frozen state [14, 29, 37]. Some researchers have found simple exponential increases in activity as soils cross 0°C mark [34, 39, 53], whereas others have found strong step changes in activity [14]. Given that permafrost soils often will remain just below the freezing point of water for long periods of time before undergoing phase change, and high Q10 values just below freezing indicate rapid changes with minimal warming, it is important that we understand the rates and controls on microbial activity at this critical point. Functional response to thaw can be examined in greater detail by tracking changes in the abundance and types of genes, transcripts, and proteins (Fig. 2.2).

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Fig. 2.2: Active microbial processes observed in intact and thawed permafrost. Metabolic models represent wide variety of functions detected in pure culture, metagenome, metatranscriptome, and metaproteome studies [41, 54–64]. Cell morphology is arbitrarily displayed. Pathways—PPP: pentose phosphate pathway; TCA: citric acid cycle; CHE: bacterial chemotaxis.

Genes show the potential of microbial populations to carry out certain functions, while transcripts show genes that are actively being expressed, and proteins show transcripts that have been translated and should be the most direct observation of the microbial cell carrying out a process. Each of these types of observations are measurements at different levels of resolution. Changes in functional gene abundance can occur over hours, days, months, and years and may reflect the long-term adjustment of the microbial community to the new environment. Proteins and transcripts can change over minutes to hours and may dominate the initial thaw response because they can change more rapidly relative to the time required for cell division and changes in community structure. Despite the importance of measuring change at all three levels of resolution – genes, transcripts, and proteins – such studies are rare.

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Permafrost is characterized by low biomass and activity, making RNA and protein extraction difficult. To our knowledge, only two studies have successfully obtained RNA and/or protein from intact permafrost, and then only in relatively young permafrost samples [41, 56]. Laboratory-based incubation experiments show that the functional thaw response measured at the level of genes [54], transcripts [56], and enzyme activity [65] is robust and rapid – on the order of days to weeks – despite the low incubation temperatures. The phase change of water, which can cause orders of magnitude-level changes in activity and metabolic processes, likely explains the quick response. Suggesting a greater reliance on respiratory process for ATP production after thaw, Mackelprang et al. [54] found that genes encoding components of the respiratory electron transport chain and enzymes that link the breakdown of organic carbon to the tricarboxylic acid cycle increased in relative abundance. Indicating a general increase in microbial activity and growth, Coolen and Orsi [56] found that transcripts involved in translation, biogenesis, and ribosomal structure were more abundant in thawed permafrost compared with frozen samples after short-term (11 days) incubation. In the field, even slight temperature increases (~1.1°C) can trigger shifts in gene relative abundance [66]. Genes changing the most indicate that soil redox state and energy acquisition drives adaptation to thawed conditions. After 5 years of warming, methanogenesis, sulfate reduction, and dissimilatory nitrate reduction genes increased in relative abundance while cytochrome c oxidase genes decreased at the active layer/permafrost transition zone [66]. Changes in respiratory functions where microbes break down organic matter and transfer electrons to an inorganic terminal electron acceptor (TEA) generate large amounts of CO2 and are particularly relevant to the climate change equation. The genomic potential to use a variety of TEAs has been found in pristine permafrost and includes genes involved in denitrification [41, 54, 58, 67–69], sulfur/sulfate reduction [57, 58, 68], and iron reduction [41, 58]. Together, these data suggest that the functional reservoir in permafrost will enable communities to rapidly take advantage of higher redox conditions after thaw. Permafrost thaw that results in soil saturation creates anaerobic conditions, which necessitates alternative TEAs or fermentation for energy production. Nitrate and nitrite may be particularly important in thawed permafrost. Permafrost nitrogen concentration can be high [70] and thaw increases nitrogen availability [12, 71–73]. Nitrogen-cycling organisms are abundant in intact permafrost [19], and laboratory and field studies show that nitrogen metabolism genes increase in relative abundance after thaw [54, 66]. In a short-term laboratory incubation experiment, Mackelprang et al. [54] observed increases in genes for multiple components of complete denitrification pathway where nitrate is reduced, eventually generating N2. Specifically, the nitrate reductase (nar), nitric oxide reductase (nor), and nitrous oxide reductase (nos) suite of genes increased relative to frozen controls. They also observed increases in respiratory nitrite ammonification (nrfAC) genes. At the

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permafrost-active layer boundary, Johnston et al. [66] observed increases in dissimilatory nitrate reduction to ammonia genes (but not in norB or nosZ genes) after 5 years of warming. Tas et al. [55] found that nitrate reduction genes (nar) were more abundant in wildfire-thawed upland permafrost compared with unthawed, intact permafrost. However, the other genes in the denitrification pathway decreased after thaw. This suggests that nitrate may not be completely reduced to N2 and may instead produce N2O, a potent greenhouse gas. Comparing intact permafrost with a thermokarst bog along a thaw gradient, Hultman et al. [41] found that the relative abundance of nitrate reductase (narG) genes and transcripts were greater in permafrost. However, denitrification and nitrate reduction rates were greater in bog samples. This discrepancy is likely due to the use of relative abundance, where the high abundance in permafrost indicates the importance of denitrification relative to other processes rather than overall process rates. One of the primary factors contributing to the functional response to thaw is the availability and quality of carbon substrates [9, 74], and it appears that the abundance of genes and transcripts is altered based on carbon chemistry. Coolen and Orsi [56] found that transcripts for an ABC type sugar transporter, 6-phosphogluconate dehydrogenase (from the pentose phosphate pathway), a sugar isomerase, and pyruvate formate lyase activating enzyme (fermentation) were more abundant in thawed permafrost. They also found an increase in transcripts from a gene encoding aminopeptidase C (involved in extracellular protein degradation), suggesting that labile proteins may be C and N sources after thaw. As evidence for rapid degradation of plant-derived cellulose, Coolen et al. [65] measured extracellular enzyme activities and found that beta-glucosidase activity was highest in thawed permafrost compared to pristine soils. Leucine aminopeptidase was initially active post-thaw but decreased, suggesting that labile proteins or polypeptides were initially available but used rapidly. Mackelprang et al. [54] also found evidence for increased use of labile carbon, although the specific substrates targeted by changing genes differed between samples. In contrast to permafrost thaw in lowlands, thaw in uplands can result in warmer and drier soils due to the loss of the insulating cover of moss and the loss of permafrost that limits the downward movement of water. In an upland soil 7 years after a fire, Tas et al. [55] found no overall increases in carbon-processing genes. Few genes (specifically those involved in galactose metabolism) were less abundant in thawed soils. Further analysis of these data showed that while there were few differences at the community level, individual members may adapt to disturbance conditions by altering their genomes [57]. Through assembling metagenomic sequence data and binning contigs into draft genomes, they were able to compare the genomes of two individual community members in pristine and fire-impacted soils. Adaptation occurred by altering 2–3% of their genomes, resulting in the loss of genes involved in carbohydrate transport and metabolism, amino acid transport and metabolism, transcriptional regulation, and nutrient transport. The reduction in carbon metabolism-related gene might reflect the lower carbon content soils post-thaw.

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Much of the variation in CH4 fluxes from permafrost landscape is due to the moisture regime and plant community present post-thaw [75]. In lowlands, carbon-rich anoxic methane-producing bogs and fens can form following thaw, ground subsidence, and water inundation [76, 77], whereas thaw occurring in uplands can result in drier soils and net CH4 uptake due to CH4 oxidation. Compared to intact permafrost and the active layer, newly forming thermokarst bog and fen features contain high levels of methanogenesis-related genes, transcripts, and proteins that indicate capacities for hydrogenotrophic and acetoclastic methanogenesis, as well as the ability to grow on formate, methanol, and methylamine [41, 78, 79]. It is clear that both hydrogenotrophic and acetoclastic methanogens play integral roles in CH4 emissions in thawed soils [80], but that the specific type depends on soil chemistry, environmental conditions, and the paleoenvironment during permafrost formation [81]. Some evidence suggests that thaw causes a shift from hydrogenotrophic toward acetoclastic methanogenesis [66, 78]. However, other studies have shown a decline in the ratio of acetoclastic to hydrogenotrophic methanogenesis across gradients of permafrost degradation in bogs and palsa peatlands [82, 83]. Holm et al. [81] found that acetoclastic methanogens dominated Eemian permafrost (formed under higher temperatures and precipitation) during thaw, whereas other Pleistocene and Holocene permafrost samples were dominated by hydrogenotrophic methanogens. Using porewater isotopes of dissolved CO2 and CH4, Neumann et al. [83] modeled microbial respiration, methanogenesis, methane oxidation, and acetogenesis at the edge and center of a thermokarst bog. At the edge of the bog, where permafrost thaw was taking place, microbial respiration, methanogenesis (both acetoclastic and hydrogenotrophic), acetogenesis, and methane oxidation all were higher than the older part of the bog, indicating that freshly released organics and/or N from permafrost soils could dramatically fuel microbial activity in situ. Five years of experimental warming at the natural thaw-front increased the relative abundance of methanogenesis genes, particularly those involved in methane production from acetate [66]. In a lab experiment of tussock tundra permafrost, Coolen and Orsi [56] found that transcripts involved in acetoclastic methanogenesis increased after 11 days of thaw. They also found that acetogenesis transcripts (but not transcripts for acetogenic fermentation) were expressed after thaw, suggesting that acetogenic bacteria are active and producing acetate post-thaw. Methanotrophs play a critical role in net CH4 flux from permafrost ecosystems. Methane that is not oxidized by methanotrophs can be released into the atmosphere. Thus, methanotrophs may mitigate CH4 emissions from thawing permafrost. In the field, genes for aerobic methanotrophy are present in the active layer of permafrost affected soils [41, 84, 85], and both aerobic and anaerobic methane oxidation occurs in thermokarst wetlands where subsurface methane concentrations are high [83, 86]. In submarine permafrost, anaerobic methane oxidizers likely mitigate methane release when thawed [87]. In uplands, where methane production is very low, genes for methanogenesis and methane oxidation are detectable but low, and they may be

2.4 Future directions 

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reduced after thaw [55]. In the lab, Mackelprang et al. [54] found that thaw yielded a burst of trapped CH4, which decreased within 7 days as methane monooxygenase (pmoA) – a key genetic indicator of methanotrophy – increased. This was despite incubation with anaerobic headspace, suggesting that oxygen necessary for methanotrophy originated from permafrost water or aerobic microsites in the soil. The methyl coenzyme-M reductase alpha subunit (mcrA) gene, which catalyzes the last step in methanogenesis, did not change in abundance. In contrast, Coolen and Orsi [56] found that methanogenesis (but not methane oxidation) transcripts increased after just 11 days of thaw. We expect that future work will enable us to better utilize genomic data to directly link gene and transcript abundance to processes that control net methane fluxes from soils.

2.4 Future directions Our understanding of the functional consequences of permafrost thaw on microbial communities is clearly still in its infancy. Although several studies have investigated changes in the composition of soil microbial communities following thaw, very few have investigated microbial communities at the level of functional omics. Several gaps in knowledge revolve around an incomplete understanding of the response and role of other domains of life to thaw, most especially fungi and viruses. We also do not know the extent to which, or under what circumstances, changes in community composition or functional gene abundance guarantees a change in biogeochemical function. The common result we observe is that microbial communities are responding to the unique physical, chemical, and substrate-accessibility conditions present within their microenvironment. Importantly, most work is conducted as a lab assay, and microbial communities in the field may not respond as they do in incubation studies because microorganisms can immigrate into newly thawed environments, the external supply of organic material or TEAs may change, and plant interactions occur whereas they are decoupled from these interactions in the lab. And although most effort is focused on how microbial responses affect CO2 and CH4 flux, we still have little understanding of the role that microbes play in affecting plant responses postthaw through mechanisms such as nutrient mineralization, symbioses, and pathogenic interactions. At the field scale, permafrost thaw occurs in many different environments in many different forms [76], thus limiting our ability to generalize results until many more sites have been examined. Each geographic location has a very different ecological and disturbance history affecting the types of microorganisms that may be entrained in permafrost and thereby affecting how they may respond post-thaw [88]. We have primarily focused on soil or wetland microbial communities in this article and have avoided lake, sediment, marine, and riverine microbial community responses to thaw. Most permafrost thaw experiments in the field are space for time substitution

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experiments. Chronosequences or other gradient studies offer the ability to look at processes over large temporal and spatial scales [5, 7, 89] but can be limited by incomplete knowledge of the history of the different sites. Field soil warming manipulations can be useful in understanding specific microbial responses to warming, but they are typically limited to surface soils [90–94]. Only one study currently exists that tracks genetic changes at the active layer-permafrost transition [66]. This may in part be due to technical constraints of heating deep soils, but techniques such as snow manipulations [40, 95] or water manipulations [96] result in subtle yet important permafrost manipulations that help us to understand system change [66]. It is important that microbial ecologists collaborate with field experimentalists to best understand microbial linkages to changing biogeochemical processes. Understanding how microbial communities respond to thaw and coupling thaw response to greenhouse gas emissions and/or plant community dynamics is important for understanding system responses. It will require a systems-level approach to investigate microbial community functional processes in the thawing permafrost over multiple spatial and temporal scales.

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[31] Walz J, Knoblauch C, Böhme L, Pfeiffer E-M. Regulation of soil organic matter decomposition in permafrost-affected Siberian tundra soils – impact of oxygen availability, freezing and thawing, temperature, and labile organic matter. Soil Biol Biochem. 2017;110:34–43. [32] Yvon-Durocher G, Allen AP, Bastviken D, Conrad R. Methane fluxes show consistent temperature dependence across microbial to ecosystem scales. Nature. 2014;507:488–91. [33] Michaelson GJ, Ping CL. Soil organic carbon and CO2 respiration at subzero temperature in soils of Arctic Alaska. J Geophys Res Biogeosci. 2003;108:1231. [34] Panikov NS, Flanagan PW, Oechel WC, Mastepanov MA, Christensen TR. Microbial activity in soils frozen to below −39°C. Soil Biol Biochem. 2006;38:785–94. [35] Amato P, Christner BC. Energy metabolism response to low-temperature and frozen conditions in Psychrobacter cryohalolentis. Appl Environ Microbiol. 2009;75:711–8. [36] Fahnestock JT, Jones MH, Brooks PD, Welker JM. Significant CO2 emissions from tundra soils during winter: implications for annual carbon budgets of arctic communities. Glob Biogeochem Cycles. 1999;13:775–9. [37] Öquist MG, Sparrman T, Klemedtsson L, et al. Water availability controls microbial temperature responses in frozen soil CO2 production. Glob Chang Biol. 2009;15:2715–22. [38] Drotz SH, Sparrman T, Nilsson MB, Schleucher J, Öquist MG. Both catabolic and anabolic heterotrophic microbial activity proceed in frozen soils. Proc Natl Acad Sci USA. 2010;107:21046–51. [39] Monson RK, Lipson DL, Burns SP, et al. Winter forest soil respiration controlled by climate and microbial community composition. Nature. 2006;439:711–4. [40] Natali SM, Schuur EAG, Webb EE, Pries CEH, Crummer KG. Permafrost degradation stimulates carbon loss from experimentally warmed tundra. Ecology. 2014;95:602–8. [41] Hultman J, Waldrop MP, Mackelprang R, et al. Multi-omics of permafrost, active layer and thermokarst bog soil microbiomes. Nature. 2015;521:208–12. [42] Gilichinsky D, Vishnivetskaya T, Petrova M, Spirina E, Mamykin V, Rivkina E. Bacteria in the Permafrost. In: Margesin R, ed. Psychrophiles: from biodiversity to biotechnology. Berlin, Heidelberg, Germany: Springer; 2008. p. 83–102. [43] Steven B, Niederberger TD, Whyte LG. Bacterial and archaeal diversity in permafrost. In: Margesin R, editor. Permafrost soils. Berlin, Heidelberg: Germany, Springer; 2009. p. 59–72. [44] Wilhelm RC, Radtke KJ, Mykytczuk NCS, Greer CW, Whyte LG. Life at the wedge: the activity and diversity of Arctic ice wedge microbial communities. Astrobiology. 2012;12:347–60. [45] Jansson JK, Tas N. The microbial ecology of permafrost. Nat Rev Microbiol. 2014;12:414–25. [46] Niederberger TD, Steven B, Charvet S, Barbier B, Whyte LG. Virgibacillus arcticus sp. nov., a moderately halophilic, endospore-forming bacterium from permafrost in the Canadian High Arctic. Int J Syst Evol Microbiol. 2009;59:2219–25. [47] Johnson SS, Hebsgaard MB, Christensen TR, et al. Ancient bacteria show evidence of DNA repair. Proc Natl Acad Sci USA. 2007;104:14401–5. [48] Nicholson WL, Krivushin K, Gilichinsky D, Schuerger AC. Growth of Carnobacterium spp. from permafrost under low pressure, temperature, and anoxic atmosphere has implications for Earth microbes on Mars. Proc Natl Acad Sci USA. 2013;110:666–71. [49] Christner BC. Incorporation of DNA and protein precursors into macromolecules by bacteria at −15 degrees C. Appl Environ Microbiol. 2002;68:6435–8. [50] Junge K, Eicken H, Swanson BD, Deming JW. Bacterial incorporation of leucine into protein down to −20°C with evidence for potential activity in sub-eutectic saline ice formations. Cryobiology. 2006;52:417–29. [51] Schaefer K, Jafarov E. A parameterization of respiration in frozen soils based on substrate availability. Biogeosci Discuss. 2015;12:12027–59. [52] Bracho R, Natali S, Pegoraro E, et al. Temperature sensitivity of organic matter decomposition of permafrost-region soils during laboratory incubations. Soil Biol Biochem. 2016;97:1–14.

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Rhiannon Mondav

3 Genomic inventory of permafrost microorganisms 3.1 Introduction The genomic inventory of a permafrost soil can be obtained through genomic characterization of the whole community (metagenomics) and/or an organism isolated from permafrost through culture or single-cell based methods. While there is often debate among the scientific community over whether genomic, transcriptomic, proteomic, or culture-based methods are best, the most promising current work combines knowledge from multiple culture-dependent and culture-independent approaches in comparative analyses of different permafrost and active-layer soils and nearby or related ecosystems. Because high-throughput shotgun methods of genomics and proteomics are relatively recent (and not always financially accessible to research groups), most work done on permafrost communities has utilized phylogenetic surveys based on small sub-unit (SSU) rRNA and functional gene (e.g. mcrA, pmoA) surveys to characterize cryotic soil communities. In less than a decade, our knowledge of permafrost and cryosol communities has shifted from piece-meal to the beginnings of a holistic understanding.

3.2 A very recent history The majority of permafrost metagenomic work has centered around three geographic regions: Alaska and Canada in northern North America [1–3], near coastal regions of Fennoscandia [4–6] and Russia in northern Eurasia, and the coastal regions of Antarctica [7]. Areas with significant permafrost that have yet to be well characterized metagenomically (at least in the English language literature) are alpine regions, plateaus, and the vast expanses of low-land permafrost in eastern Russia, some recent publications from the Tibetan plateau [8] and Siberia [9, 10] notwithstanding. Information on the genetic potential and metabolism of active-layer and permafrost microbes has been approached from both culture-dependent and culture-independent sequencing of genomes. Whole-community incubations with microcosms and ­culturing of isolates link process rates and metabolic functionality in vitro [11–13]. The struggle to isolate and culture individual bacteria and archaea from permafrost can take many years of laboratory work and is often not possible to complete within the short-term projects held by early career researchers. Short-term studies can isolate and identify ­sub-groups of microbes, often obligate aerobic heterotrophs that can grow on simple sugars and solid media. For this reason, they tend to be close phylogenetic relatives of already characterized isolates [12, 14]. Specialized media and conditions and a long generation time are necessary considerations in isolating fastidious microbes and those https://doi.org/10.1515/9783110497083-003

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from deep or very cold permafrost [15–19]. Known issues with culture-based studies are that only a small percentage of microbes can be grown axenically due to complex auxotrophies and (potentially) intra-species communication molecules and that microbial behavior in isolation is different from how a microbe functions and evolves within its community and habitat. The work of characterizing isolates is still, however, essential as it provides experimental evidence linking genomic potential to enzyme function and the metabolism of microbes. To complete the picture on how microbes behave in their environment and how a whole community functions is best examined with in situ measurements. One of the first and most significant methods developed for in situ community genomic functional potential was the GeoChip microarray. The GeoChip has been extensively deployed in circumpolar terrestrial environments, including Siberia, the Antarctic, and most recently Alaska [20–23]. GeoChip-based research allows for robust statistical comparisons of functional potential across multiple research projects. Several projects have dealt with the de novo exploration of in situ community genomics through shotgun sequencing of permafrost community DNA [3, 24]. Metabolic pathways present in the community can be predicted from metagenomic datasets and compared to other environments. Genomes of individual populations of microbial species can also be recovered from metagenomes and their metabolism predicted [1, 2, 6]. Metagenome-based metabolic predictions of uncultured microbes can inform design of growth conditions to culture and characterize a species isolate. Predicted metabolism can be tied to process rates such as done with the methanogen Candidatus Methanoflorens or the unknown “upland soil cluster α” methanotrophs when geochemical measurements such as methane (CH4) flux, or δ13CCH4 are taken and other “omic” techniques such as metaproteomics or metatranscriptomics are also employed [25, 26]. Because metagenomes include active and non-active genes of living, dormant, and dead genomes, there has been criticism that it does not accurately represent the currently active or living community. While this is true, the genomic material of non-active genes and non-active cells forms part of the genetic library and seed bank available in that environment and so is worth considering, even when not currently in use. The genomic information present in an environment is a library of the potential microbiological processes that have, could, and do occur. A genetic record of past processes and organisms can be stored and cryogenically protected in cryotic soils. Uptake and incorporation of exogenous DNA can give “new life” to genes from genomes that have been lost from the active or living population. And because the genetic repertoire of any organism, or community, is not actively transcribed or translated in its entirety at any given point in time, examining only the active processes provides in fact an accurate snapshot of the metabolic processes occurring at the time of sampling, but not the full genomic potential available in an environment. To focus on only the active members and active processes of a community, a few studies have analyzed the actively transcribed genes of either in situ [4] or cultured

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communities through shotgun sequencing of cDNA (metatranscriptomics). Metatranscriptomics provides information on the active microbes and so theoretically has higher information density than metagenomics. However, it also has higher associated expenses for both sampling and sequencing preparation, has a lower data content return due to the high abundance of rRNA transcripts, and so is not commonly used. Another weakness is that metatranscriptomics does not indicate the in situ abundance of gene products and it rarely correlates to either process rates or gene-product abundance [27]. Metaproteomics (shotgun sequencing of peptide fragments) can, however, link genomic potential to in situ activity and is most successful when paired with metagenomic sequencing of the same community. Recently, a few large research consortia have been able to investigate permafrost communities using a combination of different meta‘omic techniques (multi-omics) to construct a more complete picture of the genomic potential linked to in situ activity of permafrost microbes. Mondav et al. showed that Ca. Methanoflorens stordalenmirensis (formerly RCII) was an active hydrogenotrophic and formatotrophic methanogen using CH4 flux and stable carbon (C) isotope measurements and multi-omics [6]. Hultman et al. were able to describe and compare whole community metabolic pathways and activities while tying specific processes such as iron reduction and methanogenesis to individual species using multi-omics [1].

3.3 (Meta)metagenomics of permafrost Studies done so far on the metabolic potential encoded in permafrost genomes have noted the presence of important pathways in nutrient cycling, mostly C and nitrogen (N) and also iron (Fe). Adaptive and reactive pathways in individual genomes have also been noted and compared. A statistical analysis of how cryotic soil metagenomes compare to other habitats, however, has not been done. Major questions that are asked of permafrost metagenomes relate to element cycling in relation to climate change, habitat shift, and cold-associated genes and pathways. To address this lack, a meta-analysis of diverse metagenomes on Integrated Microbial Genomes and Microbiomes (IMG/M), including permafrost, active layer, and ice habitats, is here presented. A list of the 311 metagenomes analyzed is available by joining the group “MLC_Chpt_1.9” on IMG/M ER [28]. Differences in the percentage of a meta/genome allocated to a functional category, in this case cluster of orthologous groups (COGs), can indicate environmental selection. Fig. 3.1 shows individual environmental samples colored according to their depletion or enrichment in COG categories compared to the overall mean of the 311 selected environmental metagenomes. Permafrost samples were enriched in seven COG categories: chromatin structure and dynamics (B), carbohydrate transport and metabolism (G), signal transduction (T), transcription (K), defense (V), energy production and conversion (C), and amino

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Fig. 3.1: Differences between estimated COG category abundance in 311 environmental metagenomes by habitat. A total of 350 environmental metagenomes were selected from as wide a range of habitats and locations as possible then reduced to contain only those with >10,000 annotated genes and more than ten representatives per habitat. A table of estimated abundance of COGs was downloaded from IMG, and the percentage that each COG category contributed to a metagenome was calculated and then displayed as a z-score heat map. Metagenomes with ≥1 standard deviation (SD) enriched COG category are colored green and those depleted by ≥1 SD are colored pink. Superscript * next to COG letter denotes a category where permafrost is both enriched compared to mean and significantly enriched compared to at least three other habitats, while superscript # denotes depletion (Kruskal-Wallace, KWmc post hoc test, p