Abiotic and Biotic Methane Dynamics in Relation to the Origin of Life

MEDDELANDEN

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STOCKHOLMS UNIVERSITETS INSTITUTION

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GEOLOGISKA VETENSKAPER

No. 348

Abiotic and biotic methane dynamics in relation to the origin of life

Nguyen Thanh Duc

Stockholm 2012

Department of Geological Sciences Stockholm University SE-106 91 Stockholm Sweden

A Dissertation for the degree of Doctor of Philosophy in Natural Science

Department of Geological Sciences Stockholm University SE-106 91 Stockholm Sweden

Abstract

Methane (CH4) plays an important role in regulating Earth’s climate. Its atmospheric concentrations are related to both biotic and abiotic processes. The biotic one can be formed

either by chemoautotrophic or heterotrophic pathways by methanogens. Abiotic CH4 formation can occur from several sequential reactions starting with H2 production by serpentinization of Fe-bearing minerals followed by Fischer-Tropsch Type reactions or thermogenic reactions from hydrocarbons. In the presence of suitable electron acceptors, microbial oxidation utilizes CH4 and contributes to regulating its emission. From the perspectives of astrobiology and Earth climate regulation, this thesis focuses on: (1)

Dynamics of CH4 formation and oxidation in lake sediments (Paper I), (2) Constructing an automatic flux chamber to facilitate its emission measurements (Paper II), (3)

dynamics of both abiotic and biotic CH4 formation processes related to olivine water interaction in temperature range 30 - 70°C (Paper III and IV).

Paper I showed that potential CH4 oxidation strongly correlated to in situ its formation rates across a wide variety of lake sediments. This means that the oxidation rates could be enhanced in environments having the high formation rates. Thereby, the oxidation would likely be able to keep up with potentially increasing the formation rates, as a result diffusive

CH4 release from freshwater sediments might not necessarily increase due to global warming. Paper II presented a new automated approach to assess temporal variability of its aquatic

fluxes. Paper III and IV together revealed that H2 can be formed via olivine-water interaction.

Abiotic CH4 formation was formed likely by Fischer-Tropsch Type reactions at low inorganic carbon concentration but by thermogenic processes at high inorganic carbon concentration.

Paper IV showed that biotic methanogenic metabolism could harvest H2 and produce CH4. The dynamics of these processes seemed strongly affected by carbonate chemistry.

List of publications

Included in this thesis:

I. Duc NT, Crill P & Bastviken D (2010) Implications of temperature and sediment characteristics on methane formation and oxidation in lake sediments. Biogeochemistry 100: 185-196

II. Duc NT, Silverstein S, Lundmark L, Reyier H, Crill P & Bastviken D, An automatic flux chamber for investigating gas flux at water – air interfaces. To be submitted to Limnology and Oceanography: Methods

III. Neubeck A, Duc NT, Bastviken D, Crill P & Holm N (2011) Formation of H2 and CH4 by weathering of olivine at temperatures between 30 and 70 degrees C. Geochemical Transactions 12: 6

IV. Duc NT, Neubeck A, Plathan J, Holm NG, Sjöberg B-M, Crill P & Bastviken D, The potential for abiotic and biotic methane formation fueled by olivine dissolution. Submitted to Nature Geoscience.

Not included in this thesis:

V. Nguyen Anh Mai, Nguyen Thanh Duc and Knut Irgum, Sizeable Macroporous Monolithic Polyamide Entities Prepared in Closed Molds by Thermally Mediated Dissolution and Phase Segregation, Chemistry of Material, 2008, 20 (19), pp 6244– 6247

VI. Duc NT, Silverstein S, Lundmark L, Wik M, Crill P & Bastviken D, Automated detection of bubble release from water columns. Manuscript

VII. Duc NT, Truc NM, Dung LT and Hung VT, Development of handy spectrometer and chemical kits for quick assessment of seawater quality, Journal of Marine science and Technology, 2010, 10 (1), pp 37-49. Vietnam Academy of Science and Technology. (in Vietnamese)

VIII. M. van Hardenbroek, A.F. Lotter, D. Bastviken, N.T. Duc, O. Heiri, Relationship between δ13C in invertebrate remains and methane flux in Swedish lakes. In Press in Freshwater Biology.

IX. Neubeck A., Duc NT, Hellevang H, Plathan J, Bastviken D, Bacsik Z., Crill P and Holm NG, The effect of dissolved inorganic carbon on low temperature olivine alteration and H2 formation. Manuscript.

Paper I was reprinted with permission of Springer.

My contribution to papers:

Paper I: Main responsibility for all parts of the study from sample collection to manuscript preparation.

Paper II: Design of the power control board, set up of the device by integrating all components, main responsibility for all tests, writing the C code to control the device, and for the manuscript preparation.

Paper III: A significant share of the planning and practical work including main responsibility for gas phase analyses as well as contributions to the manuscript preparation.

Paper IV: A significant share of the planning and practical work including main responsibility for gas phase analyses and for the manuscript preparation.

Stockholm, January 26th, 2012

Nguyen Thanh Duc

Contents

1 Introduction ...... 7

2 Background ...... 7

2.1 From Contemporary Earth ...... 7

2.1.1 Biotic CH4 formation and oxidation ...... 7

2.1.2 Environments and transport processes of importance for CH4 emissions ...... 8

2.2 Back to Early Earth ...... 9

2.2.1 Abiotic MF ...... 10

2.2.2 The emergence of life and CH4 metabolism ...... 14

2.2.3 Climate feedback ...... 14

3 Objectives ...... 15

3.1 Temperature effects on biotic CH4 dynamics in sediments...... 15

3.2 Constructing automatic floating flux chambers...... 16

3.3 Abiotic MF via water-olivine interaction ...... 16

3.4 Biotic MF via water-olivine interaction...... 17

4 Summary of the work...... 17

4.1 Implications of temperature and sediment characteristics on MF and MO in shallow lake sediments (Paper I)...... 17

4.2 An automatic flux chamber for investigating gas flux at water – air interfaces (Paper II)……...... 20

4.3 Formation of H2 and CH4 by weathering of olivine at temperatures between 30 and 70°C (Paper III) ...... 24

4.4 The potential for abiotic and biotic MF fueled by olivine dissolution (Paper IV) ...... 26

5 Discussion ...... 29

5.1 Biotic MF and MO from sediment ...... 29

5.2 CH4 emission measurement ...... 29

5.3 Abiotic MF from olivine-water interaction at temperatures between 30 and 70°C ...... 30

5.4 Biotic MF from olivine-water interaction at temperatures between 30 and 70°C ...... 35

5.5 H2 formation from olivine-water interaction at temperatures between 30 and 70°C ..... 37

6 Conclusions ...... 40

7 Acknowledgments...... 41

8 References ...... 44

1 Introduction

Methane (CH4) is an interesting molecule not only for studies climate change but also in astrobiological research. CH4, which is a greenhouse gas, contributes considerably to global warming due to its radiative properties of absorbing and emitting radiation within the thermal infrared range (IPCC 2007). From a general physical perspective, the Earth would start cooling soon after the accretion period ca. 4.65 Ga ago because accretion processes had left a hot planet in a cold interplanetary space (Pater & Lissauer 2001). The occurrence of life required the temperature of a substantial portion of Earth’s surface to be maintained between the freezing and the boiling points of water (Kasting et al. 1993). The energy required for this could have come from solar radiation, radioactive, relic heat from accretion, from Earth’s interior or water-rock interaction, such as serpentinization and the greenhouse effect slowing down the heat loss (Davies 1980; Kasting 1993; Smith 1981). With time, some of these heat sources including radioactive, relic from accretion and water-rock interaction, decreased while increasing energy flux from the Sun in combination with the greenhouse gas effect has kept the earth surface temperatures in the habitable range (Kasting 1993).

CH4 has been present in the atmosphere throughout Earth history because it is continually formed by both abiotic and biotic processes. Besides CH4 sources, there are also

CH4 sinks which contribute to the regulation of the CH4 concentration on Earth. Most of these processes are well defined but their dynamics are still being considered. Understanding

the dynamics of CH4 on Earth can support our understanding of both the present day and historic climate of the Earth and is the focus of this thesis.

2 Background

2.1 From Contemporary Earth

2.1.1 Biotic CH4 formation and oxidation

Biogenic production is the main source of contemporary atmospheric CH4. It can be formed either by chemoautotrophic or heterotrophic pathways by methanogens. These

archaeal microorganisms can use CO2, H2 or acetate, formate, methanol, and methylamines as substrates (Jeris and McCarty 1965; Smith and Mah 1966; Cappenberg and Prins 1974; Wolin 1976; Oremland 1988; Zinder 1993). In contemporary ecosystems, the overall biotic

CH4 formation (MF) depends not only on methanogenesis itself but also on the preceding

organic degradation steps producing substrates for methanogenesis. In anoxic environments, polymeric and complex organic matter is fermented by different groups of anaerobic microbes to smaller compounds (Fenchel & Finlay 1995). As a part of these anaerobic degradation processes, some microorganisms release exoenzymes to cut polymeric organic compounds to lower molecular weight compounds that can be imported into the cell. Once in the cell the organic matter is fermented to various fermentation products including short

chain fatty acids, alcohols, H2 and CO2. These low molecular weight organic products are

then used by other microorganisms that catalyze the terminal mineralization to CO2 or CH4.

H2 produced by some fermentation reactions can be used with CO2 by acetogens which

generate acetate as a respiration product or by methanogens resulting in CH4 production (Cicerone & Oremland 1988; King 1992; Valentine et al. 2000). Both organic and inorganic fermentation products can be the precursors for methanogenesis (Conrad 2005; Whalen 2005; Zeikus 1977). Biological methanogenesis relies on the activity of fermenting and other organisms, unless substrates are provided by abiotic processes such as serpentinization (described further below). The complete degradation of organic matter in anoxic environments is a result of tightly linked processes performed by a microbial community or consortia of several types of microorganisms (Oremland 1988).

While biotic MF in itself is performed by one specific group of microorganisms, the

CH4 emission to the atmosphere is ultimately regulated by not only microbial production but also microbial oxidation (Bartlett & Harriss 1993; Kankaala et al. 2003; King 1992; 2- Reeburgh 2003; Segers 1998). In the presence of suitable electron acceptors (e.g. O2, SO4 ),

CH4 is used as an energy and carbon source for CH4 oxidizing bacteria (King 1992). In

freshwater ecosystems, aerobic CH4 oxidation (MO) can occur. This process primarily

occurs at oxic-anoxic interfaces where both CH4 and O2 are abundant. Both biotic MF and MO are potentially regulated through microbial activity by several environmental factors that affect the oxidation state including temperature, organic substrate quality and supply, nutrient availability and oxygen concentration (Megonigal et al. 2005). Together, biotic MF and MO

will regulate net CH4 production.

2.1.2 Environments and transport processes of importance for CH4 emissions

Shallow freshwater sediments in wetlands and lakes contribute in the order of 75 % of

the total non-anthropogenic contemporary CH4 emissions (Bastviken et al. 2008; Reeburgh 2003; Walter & Heimann 2000). The environmental factors mentioned above can all be

variable in such environments. Hence, for the discussion of how CH4 emission is affected by

environmental change, the study of CH4 dynamics in shallow sediments appear particularly important.

CH4 produced in shallow lake sediments is much more likely to reach the atmosphere than CH4 produced in deeper sediments (Bastviken et al. 2008). CH4 is emitted to atmosphere from shallow sediments via at least three different pathways: ebullition, diffusive flux, and flux through aquatic vegetation (Bastviken et al. 2004). This makes CH4 emission evaluation difficult because different emissions pathways are regulated by different

environmental factors. During ebullition, CH4 is released as bubbles which rapidly pass through the sediment surface and water column, and thereby bypasses the MO zones. This flux may primarily be affected by MF rates and physical conditions such as sediment structure and weather conditions that affect lake mixing and hydrostatic pressure variability (Chanton et al. 1989; Chanton & Whiting 1995; Keller & Stallard 1994). During diffusive

flux, dissolved CH4, which enters water column from sediment, diffuses along the concentration gradient to the atmosphere. In water columns, this concentration gradient is maintained by advection and turbulence. Therefore, this flux is strongly affected by MO and

CH4 exchange rates across the sediment-water and water-atmosphere interfaces. During flux

through aquatic vegetation, vascular plants can function as conduits for CH4 transport to the

atmosphere. However vascular plants also transport O2 to the rhizosphere which may result in decreased methanogenesis and increased MO in the rhizosphere, and thereby decrease rates of ebullition or diffusive export (Dinsmore et al. 2009; Visser et al. 2000; Wießner et al. 2002).

2.2 Back to Early Earth

In contrast to apparently lifeless neighbor planets, contemporary Earth has an

atmosphere consisting of greenhouse gases including CO2 and CH4 as minor components

(Encrenaz 2007). Atmospheric CH4 and O2 levels have been inversely correlated when the Earth’s atmosphere evolved from a reduced to an oxidized state (Kump 2008; Zahnle et al.

2006). During the period with a reducing atmosphere, CO2 and CH4 are believed to have played an important role in regulating early Earth’s climate, counteracting low total solar radiation because of a faint young sun (Kasting 2005; Kiehl & Dickinson 1987; Kral et al. 1998; Pavlov et al. 2000). The first two billions years of Earth (4.5 – 2.5 Ga) was characterized by anoxic atmosphere and ocean (Canfield 2005; Holland 2006; Kasting & Ono

2006) and was likely to be favorable for MF processes. The anoxic early Earth was influenced by the dominance of reduced igneous rock on Earth’s surface after planetary accretion from solid materials in the solar nebula (Jones 2004; Ringwood 1966).

Olivine, a magnesium-iron silicate, was one of the most common minerals in the early Earth lithosphere (Hazen et al. 2008; Sleep et al. 2004). Within a few tens of Ma after planetary accretion, liquid water might have been present on Earth according to evidence from oxygen isotope ratios found in 4.4 Ga zircons from Jack Hills regions in Western Australia (Mojzsis et al. 2001; Wilde et al. 2001). Therefore, low-temperature (< 100 °C) aqueous alteration of olivine should have occurred at an early stage on Earth (Hazen et al. 2008; Valley et al. 2002). This aqueous alteration not only derived new secondary minerals in

the lithosphere including serpentine, chlorite, hydroxides and carbonate but also produced H2 that could potentially be an important reactant in abiotic organic synthesis processes (Hazen et al. 2008; Holm & Charlou 2001; McCollom & Seewald 2007).

H2 production has been observed in nature and in laboratory experiments at temperatures and pressures that range from 30 - 400°C and 1 - 500 bars (summarized in Table 1, Paper III and IV). This has been explained by the reactivity of iron-rich olivine to reduce

H2O by serpentinization according to the reaction

2+ 3+ − 2FeOlivine + 2H 2O(l) → 2Fe2ndMineral + H 2(aq) + 2OH (aq) (R1)

2+ 3+ Fe Olivine refers to the ferrous component of olivine and Fe 2ndMineral refers to ferric

component of Fe-bearing secondary mineral. In this reaction,H2 formation is proportional to olivine dissolution and the amount of Fe2+ that can be oxidized to Fe3+ (Hellevang 2008;

Marcaillou et al. 2011). As H2O is decomposed, it produces H2. The reactive sites have very

low redox potential which allows the abiotic reduction of CO2 to CH4 abiotically through processes described below, or biotically by methanogenic metabolism (Martin et al. 2008; McCollom 2007; McCollom & Seewald 2007; Proskurowski et al. 2008a; Proskurowski et al. 2008b)(Paper III and IV).

2.2.1 Abiotic MF

H2 can be used as the electron donor in the reduction of CO and CO2 to organic

compounds including CH4 as a result of Fischer-Tropsch type (FTT) reactions (Berndt et al. 1996; Holm & Charlou 2001; McCollom & Seewald 2001; 2007; Proskurowski et al. 2008b). In general, this process can be described schematically as follows: 10

CO2(aq) + 4H 2(aq) = CH 4(aq) + 2H 2O(l) (R2)

This reaction has been studied based on theory, laboratory experiment and field observations under conditions representative of deep sea hydrothermal systems (summarized in Table 1). Theoretical calculations show that R2 is thermodynamically favorable at temperatures lower than 450°C. Experiments under hydrothermal conditions showed that

only a small fraction (about 0.5 to 1 %) of dissolved CO2 was converted to CH4 while a large fraction of CO2 was converted to carbonaceous material (Berndt et al. 1996; Foustoukos &

Seyfried 2004; McCollom & Seewald 2003a). By injecting H2 into the reaction cell of these

experiments, they demonstrated that MF is proportional to H2 concentration (Foustoukos & Seyfried 2004). Catalysts such as Ni/Fe-alloys, Fe/Cr oxides and magnetite have been proven to increase the rate of reaction R2 (Foustoukos & Seyfried 2004; Horita & Berndt 1999). In general, experiments have shown that kinetics of reaction R2 depend upon the concentration of dissolved reactants, temperature, pressure and the availability of a catalyst (McCollom &

Seewald 2007). Beside H2 and CO2 as the initial reactants for abiotic MF, carbon monoxide (CO), which can be formed in the water-gas shift reaction (R3), may be involved in, and enhance MF by reaction R4 below (McCollom & Seewald 2007).

CO2 + H 2 = CO + H 2O (R3)

CO + 3H 2 = CH 4 + 2H 2O (R4)

In CO2 reduction processes under hydrothermal conditions, other single carbon compounds, such as formic acid, formaldehyde, methanol, can be formed (Seewald et al. 2006). Depending on the availability of catalysts such as Ni/Fe alloys, C1 compounds can be

accumulated or transformed to CH4 (McCollom & Seewald 2003a; Seewald et al. 2006). Longer hydrocarbon compounds from C2 to > C35, including amino acids, n-alkanes, n- alkenes, n-alkanols, n-alkanoic acids, n-alkyl formats, n-alkanals , and n-alkanones, can be formed under hydrothermal conditions (Andersson & Holm 2000; McCollom et al. 1999; McCollom & Seewald 2003b; 2007; Proskurowski et al. 2008b; Rushdi & Simoneit 2001).

These long hydrocarbon compounds can be decomposed and transformed to CH4 via pyrolytic (or thermolytic) processes (Welhan 1988). Abiotic MF coupled with serpentinization under hydrothermal conditions is well recognized at temperatures above 200°C (Table 1). Knowledge about the potential for abiotic MF under conditions which are

suitable for life is limited. While serpentinization generating H2 can occur at temperatures

less than 100 °C and at atmospheric pressure (about 1 bar) (Stevens & McKinley 2000), it is unclear whether abiotic MF via processes such as FTT or thermolysis can occur under such conditions. MF at low temperature natural sites such as Atlantic Lost City hydrothermal system and ophiolitic sites of Philippines, Oman, New Zealand, Turkey (summarized in Table 1) and on Mars have been observed but the MF mechanisms are still being investigated. In laboratory studies MF via FTTs at temperatures below 50 °C and atmospheric pressure has only been proven with specially synthesized catalysts (Jacquemin et al. 2010; Thampi et al. 1987). If MF can occur at low temperature with natural catalysts available as abundant minerals on the Earth, this would have large implications for how widespread abiotic MF could have been and to what extent it could have contributed to the

global CH4 budget and early Earth’s climate regulation.

Table 1. Summary of abiotic H2 and CH4 studies in relation to mineral-water interaction.

13 13 Category Material T P H2 CH4 δ CH4 δ CO2 Reference (°C) (bar) (mmol kg-1)* (mmol kg-1)* (‰ PDB) (‰ PDB) Theoretical Peridotite 350 500 3 ÷ 164.9 NA NA (Wetzel & Shock 2000) Theoretical harzburgite 50 ÷ 350 >0 ÷ 350 NA NA NA (McCollom & 400 Bach 2009) Hydrothermal Peridotite + 40 ÷ 249 ÷ 428 136 ÷ 285 NA (Kelley et al. system-Lost city gabbro 75 2001) Hydrothermal Peridotite + 40 ÷ <1 ÷ 15 1 ÷ 2 -13.6 ÷ - -8 ÷ -2 (Kelley et al. system-Lost city gabbro 90 8.8 2005) Continental 110 ÷ 8.4 ÷ 42.6 (mol 13.0 ÷ 55.3 - 7.0 ± -32 (Abrajano et al. ultramafic-hosted 125 %) (mol %) 0.4 1988) area - Zambales Ophiolite Continental 32 ÷ < 1 ppm ÷ 81 % <0.002 ÷ 2.2 % -34.5 -9.6 ÷ - (Sano et al. ultramafic-hosted 67 10.7 1993) area - Oman Ophiolite Columbia Basalt 22 1 0.06 0.002 ÷ 0.481 NA -20 ÷ + 20 (Stevens & River Basalt Group McKinley 1995) Hydrothermal Basalt 307 ÷ 0.08 ÷ 0.51 0.05 ÷ 0.12 -20.1 ± -4.08 ± (Proskurowski system- East 388 1.2 0.16 et al. 2008a) Pacific Rise Hydrothermal Basalt 18 ÷ 0.01 ÷ 0.16 0.06 ÷ 1.87 -30.2 ± - 4.55 ± (Proskurowski system- East 33 2.7 0.5 et al. 2008a) Pacific Rise Hydrothermal Peridotite + 365 16 2.5 -15.8 -3.15 (Charlou et al. system-Rainbow gabbro 2002) Experiment Olivine – 300 500 0 ÷ 158 0.084 NA NA (Berndt et al. Fo88 1996) Experiment Olivine 300 350 0 ÷ 17 0.0188 11% 99% (McCollom & 13 13 CH4 NaH CO3 Seewald 2001) Experiment Olivine 177 ÷ 350 16 ÷ 107 0.0067 ÷ 0.037 1 % 99% (McCollom & 13 13 250 CH4 NaH CO3 Seewald 2003a) Experiment Peridotite a 200 500 0 ÷ 76.7 NA NA NA (Seyfried Jr et al. 2007) Experiment Olivine-Fo89 400 500 1 ÷ 2.3 NA NA NA (Allen & Seyfried 2003) Experiment Harzburgite b 300 500 0.1 ÷ 0.33 0.066 NA NA (Janecky & Seyfried Jr 1986) Experiment Basalt c 300÷ 400 0.06 ÷ 0.87 1 ÷ 1.6 NA NA (Seewald & 400 Seyfried Jr 1990) Experiment Basalt 30 ÷ 1 0 ÷ 56 NA NA NA (Stevens & 60 (nmol/g_sample) McKinley 2000) Experiment forsterite 30 1 2 ÷ 14 NA NA NA (Stevens & (nmol/g_sample) McKinley 2000) Experiment Olivine_Fo91 30÷70 1 0.1 ÷ 1.28 0.048 ÷ 0.208 NA NA (Paper III) (nmol/g_sample) (nmol/g_sample) Experiment Olivine_Fo91 30÷70 1 0.1e-4 ÷ 1.2e-4 5.5e-11 ÷ 4.1e-7 -51.08 ÷ - 90% (Paper IV) 13 42.09 NaH CO3 Note: *: the dissolved gas concentration per kg solution. Unit of underline data was original from the publications. (a): 62 vol % olivine, 26 vol% orthopyroxene, 10 vol % clinopyroxene and 2 vol% spinel. (b): 75 wt% olivine, 25 wt% orthopyroxene. (c): It has a hyalophitic texture composed of microphenocrysts of plagioclase and clinopyroxene with approximately 5% large (0.5 mm) euhedral plagioclase phenocrysts, 3% olivine, 3% glass and 10% opaques.

2.2.2 The emergence of life and CH4 metabolism

There are still questions about how and when life appeared on Earth. Life occurred on Earth and changed the biogeochemical cycles with its metabolism. Theoretical studies have suggested that the geochemistry at hydrothermal systems might have initiated the chemistry of life (Corliss et al. 1981; Holm 1992; Martin et al. 2008). Water-rock interactions at these sites provided high flux of thermal energy and essential substrates favorable for microorganisms (Kelley et al. 2005). So this process can have been important in the context of the origin of life, not only because of its contribution to production of hydrocarbons on

early Earth but also its possibility to generate energy sources that could sustain life (e.g. H2) (Russell et al. 2010; Stevens & McKinley 1995).

Atmospheric CH4 is efficiently oxidized via direct photolysis and reaction with

hydroxyl radicals produced from H2O photolysis. It is likely that there was a higher UV flux

onto the early Earth atmosphere due to the lack of an O3 layer (Haqq-Misra et al. 2008). In order for CH4 to be abundant in early Earth atmosphere, abiotic MF alone may not have been

sufficient (Kasting & Ono 2006). It has been suggested that biotic methanogenesis via CO2

reduction with H2 (Zinder 1993) appeared early (about 3.46 - 3.9 Ga) in Earth’s history 13 (Kasting 2005). This is supported by isotopic signature of reduced C (δ CH4 < -56 ‰) in

CH4-bearing fluid inclusions (approx 3.5-Gyr-old hydrothermal precipitates from Pilbara craton, Australia) and isotopically light carbon (δ13C value from -21‰ to -49 ‰) in the carbonaceous inclusions within grains of apatite from the oldest known sediment sequences; ~3.8-Gyr-old banded iron formation from the Isua supracrustal belt, West Greenland and a similar formation from the nearby Akilia island (Mojzsis et al. 1996; Rosing 1999; Ueno et 13 12 13 al. 2006). The mass of C is 8.36% greater than that of C. In nature, abiotic CH4 (δ CH4 13 range from roughly -50 ‰ to -20 ‰) is generally enriched in C compared to biotic CH4 13 (δ CH4 range from roughly -110 ‰ to -50 ‰) (Whiticar 1999). This stable isotopic dissimilarity can be generally explained by biologically preferred 12C, which causes the 13 13 depletion of C in biological products. However, this is not exclusive because the δ CH4 signature is related to several factors including precursor compounds, MO, and the type and magnitude of the isotopic kinetic effect and temperature.

2.2.3 Climate feedback

Model results of Kasting (2005) implied that the high atmospheric CH4 concentration of the early Earth eventually decreased, likely as a consequence of the development of an

oxic atmosphere (Kasting 2005). The decrease of atmospheric CH4 is coincident with the

onset of global glaciations (Pavlov et al. 2003). The Sun was still faint (80% of So) but solar luminosity is calculated to have a steady linear increase (Gough 1981; Haqq-Misra et al. 2008) . Many factors have been involved in the regulation of Earth climate over time. The

balance between CH4 sources and sinks is a part of this regulation and several abrupt climate

shifts have been attributed to changes in atmospheric CH4 concentrations (Shaw 2008). Feedbacks on MF and MO are likely to have occurred and, in turn, have implications for the

global climate. Increased knowledge about CH4 production and oxidation under various environmental conditions can improve our understanding of Earth climate regulation both presently and in the past.

3 Objectives

This thesis will cover several aspects of CH4 biogeochemistry addressing processes believed to be important under both contemporary and early Earth conditions. An underlying assumption regardless of how each study is introduced is that studies focusing on contemporary conditions can contribute to knowledge about the past and the future.

3.1 Temperature effects on biotic CH4 dynamics in sediments.

The most abundant forms of carbon on early Earth were probably inorganic: CO2 in the atmosphere, dissolved inorganic carbon in the hydrosphere, and carbonate in the lithosphere (Sundquist & Visser 2003). Over Earth history, organic carbon became more abundant and the atmosphere became oxygen-rich (Holland et al. 1986; Sundquist & Visser

2003). It still not clear whether anaerobic MO would have contributed to CH4 dynamics

under the early reducing atmosphere, but there is no doubt that the CH4 dynamics on contemporary Earth is regulated by both MF and MO. Lake sediment was a focus because it

is one of the most important environments for CH4 cycling on modern Earth. Freshwater

sediments and peat represent the most important natural source of atmospheric CH4 (Bastviken et al. 2004; Crill et al. 1991; Walter et al. 2007). A question of importance for climate feedbacks is to what extent the MF and MO balance is affected by global warming. Increased temperatures are likely to result in increased MF (Segers 1998) while it is less clear

whether this would lead to increased CH4 emissions. Can MO counterbalance an enhanced MF rate and thereby limit emission changes? The aim of the study presented in Paper I was to measure and compare rates of MF and MO at different temperatures in the same sediments in the context of their interaction in order to understand potential feedbacks.

3.2 Constructing automatic floating flux chambers.

Knowledge about CH4 emissions to the atmosphere is necessary to understand how the transport of CH4 to the atmosphere is regulated. The measurement of CH4 flux is challenging because spatial and temporal variability can be large. Particularly ebullition of

CH4, a major flux pathway from inland waters/wetlands, is highly variable in space and occurs episodically with large fluxes over short time periods (Bastviken et al. 2004; 2008).

Therefore, CH4 emission data with high temporal resolution would be very beneficial for emission assessments. Manual chamber measurements of flux are commonly used in aquatic environments. They are labor intensive if used to assess spatial and temporal variability but they represent a robust and conceptually straightforward measurement technique of a well- defined area.

This study aims at developing portable stand-alone, inexpensive units for automated measurements of gas emissions from aquatic environments. A design challenge is to combine the robustness and flexibility of the flux chamber technique with the ability to obtain measurements at high temporal resolution and less labor. The goal is also to keep the cost down to allow the simultaneous use of many such units to cover as much area as possible.

An automatic flux chamber (AFC) for automatic sampling and monitoring CH4 flux has been developed and is described in Paper II.

3.3 Abiotic MF via water-olivine interaction

This particular study focuses on abiotic MF driven by olivine-water interactions (see above). Abiotic MF will only occur if the thermodynamic conditions are favorable (Shock 1990; 1992). So far, numerous laboratory experiments which simulate hydrothermal systems (i.e. high temperature from 100°C to 500°C and high pressure 100-500 bar) showed the

formation of H2 and CH4 (Table 1). Such studies have implications about early Earth in extreme conditions. However, these types of conditions, which happened either on short time scales (Ma) relative to Earth’s history (Ga) (Hazen et al. 2008; Rosing et al. 2006; Zahnle et al. 2007) or locally (e.g. hydrothermal vent in deep ocean) were not representative for the habitable zone on Earth. We also need to understand more about abiotic processes of serpentinization and FTT reactions at temperatures below 100°C and at lower pressures because, even if production rates are low, they could be significant given long time periods and large areas having such conditions. This study therefore asks the questions: How

extensive can the abiotic production of CH4 by water-olivine interactions be at temperatures

between 30 and 70 ºC and at 1 atm pressure? This study aimed to test if there was detectable MF from olivine and water interactions at 30 to 70 ºC and, if so, over what time scales (Paper III).

3.4 Biotic MF via water-olivine interaction.

As mentioned above, life could have emerged early and, if so, it has existed on Earth

for billions of years. H2 from water-rock interactions could have provided important energy sources for the first life forms on Earth. Chemolithoautotrophy has been suggested as a main metabolic pathway of ancient microorganisms (Canfield et al. 2006; Dobretsov et al. 2006; Kasting 2005; P. Kharecha 2005). In this follow-up study to Paper III, the biotic MF was

considered as a potentially important CH4 source in addition to abiotic CH4 source on the early Earth. In this study we tested the potential of olivine-water interactions to fuel methanogenic metabolism over extended time periods (Paper IV).

4 Summary of the work

4.1 Implications of temperature and sediment characteristics on MF and MO in shallow lake sediments (Paper I).

Study site

Potential CH4 dynamics was studied by incubating sediments from the littoral zones of eight lakes in central Sweden. These lakes were chosen according to differences in nutrient levels and quantity and quality of organic matter (Table 1, Paper I). The catchments of these lakes were also quite different and included coniferous forest, podsol soils over granitic bedrock and agricultural loam or clay soils. Therefore, the sediment types vary from lake to lake. Their characteristics were determined by water content, percentage of organic carbon and C:N ratio (Table 2, Paper I).

Methodology

Potential rates of MF and MO were determined by incubating sediment slurries and monitoring headspace CH4 concentration over time. Sediment slurries were made from the uppermost 5 cm of the sediment which were retrieved from 6.4 cm i.d. cores collected at 1-2 m water depth. Incubations were carried out in the dark at 4 temperature levels: 4oC, 10oC, 20oC and 30oC. For both MF and MO, there were three replicate slurry bottles per temperature.

Results and Discussion

MF was sensitive to temperature with higher rates at higher temperatures in all sediments (Figure 1). The temperature response was exponential in most cases with the most drastic increase between 20oC and 30oC (Figure 2 in Paper I). The lakes that had low o potential MF at 20ºC produced much more CH4 when the temperature was increased to 30 C. A two-way ANOVA showed that MF rates were significantly affected by temperature and C:N ratio (p ≤ 0.001 for both variables; Table 4 in Paper I). At the in situ temperature, different lake sediments had different potential MF rates. These differences could be explained by several factors including nutrient content, organic matter quality, O2 penetration depth, abundance and activity of relevant microbial populations.

MO was less affected by temperature (Table 6 in Paper I), but was correlated with MF at the in situ temperature ( 20 °C) and the sediment total organic matter content (Figure 2; Correlation analyses, Table 4 in Paper I). A two-way ANOVA was run with temperature as a factor and MF at 20 ºC, sediment water content, organic matter content and C:N ratio as covariates. This ANOVA showed significant effects from MF at 20 ºC (p < 0.005), indicating that MO rates were primarily correlated to in situ MF rates rather than sediment characteristics or temperature. This supports the idea that MO is primarily substrate regulated and depends on CH4 and O2 levels.

Figure 1: Potential MF (gray bars) and CH4 oxidation (white bars) rates at different temperatures in the studied sediments. The graphs were ordered by total phosphorous concentration in the water to some extent reflecting lake productivity (see Table 1). Error bars denote ± 1 SD (n=3). Note that the y axis is logarithmic and that scales differ between the lakes.

The observation that MO depends on MF implies that MO rates could adjust to an

increasing MF given sufficient time for the CH4 oxidizing microbial populations to respond to elevated CH4 concentrations. Our result that MO was correlated to MF across highly different sediment types representative for sediments in the temperate and boreal climate

zones (Figure 4 in Paper I) is supported by findings that MO kinetics depend on CH4 concentration (Bender & Conrad 1995; King 1997; King et al. 1990; Knief & Dunfield 2005; Segers 1998).

Conclusions

Altogether our results and the literature evidence imply that MO will adapt to rising

CH4 concentrations. Therefore, we cannot just assume that CH4 emissions will increase following warming. Instead MO may respond rapidly enough to counteract large diffusive

emission increases from sediments in a warmer climate under stable conditions. Future CH4 emissions may not be easily predictable from environmental conditions and robust direct emission measurements are needed as discussed below (Paper II).

4.2 An automatic flux chamber for investigating gas flux at water – air interfaces (Paper II).

The AFC was developed based on a common, general-purpose datalogger board built around an inexpensive 16-bit flash microcontroller (Microchip PIC24). The board, designed by Dr. Samuel Silverstein at the Stockholm University Department of Physics.

Description of automatic sampling device: This device was developed from classical techniques for flux measurements based on floating chambers. The automatic sampling device was designed to collect samples at desired

time intervals while at the same time monitoring in situ CH4 concentration in the chamber headspace. There are additional inputs, so selected environmental variables can be simultaneously monitored. This AFC was configured to control up to 22 peripheral devices including a pump and electronic valves, the last two of these 22 channels can be switched to analog signal receiver (Figure 2). The left four channels were set as analogue receiver to monitor battery level, atmospheric pressure, water and air temperature. Depending on sample collection intervals the device can be deployed for up to 20 days without a solar panel, recording information such as air and water temperatures and atmospheric pressure.

Figure 2: PIC24 data-logger board (top right), a power delivery board (top left), the gas pump for pumping samples from chamber to sample containers (bottom left) and solenoid valves controlled by the PIC24 (bottom right).

The AFC was designed to automatically collect samples. The power control board composed of 22 parallel 4 A power transistors, in assisting of 1000 μF capacitor (Figure 3), which used as general-purpose switch to control electronic pump and valves. The power transistor was used as a complementary amplifier for 18 mA current sending from datalogger/controller board. The digital signal from datalogger/controller board was activated according to program which was written in C code. In addition, this AFC can read basic setting parameters, such as the number of samples, how much gas to sample (sampling duration) and when sample should be collect, etc, from SD card.

Figure 3: Principle schematic of power controller board: connector -1,-2 for peripheral device; digital-1 for receiving digital signal from the data logger/system controller board. When the base (pin 3) of transistor is triggered by the digital signal, collector (pin 2) connects to emitter (pin 1), the circuit closes. This will start the peripheral device which connects to connector -1, -2. The capacitor will help the start up process as a power buffer. In this study, the peripheral devices are the electromechanical solenoid devices such as pump, valve. Therefore, a diode is connected to prevent the reversed current to protect this circuit.

This AFC was equipped with a CH4 sensor and a 10 bit high speed AD converter (conversion speed: 5.105 samples per second) of datalogger, could solve the requirement

about temporal resolution in measuring CH4 emission. Comparing with long time deployment and manual sampling method that have been used before, this AFC will provide detailed information about CH4 emission including diffusion and ebullition flux that accumulates in the floating chamber. Altogether, this AFC is fully programmable and can be adapted to the temporal resolution requirements of different systems.

Performance

In the field, the AFC has two main states comprising a full measurement cycle: accumulation state (A) and sampling state (S) including subroutines to flush the chamber with external air to restart the accumulation state. During the accumulation state, the AFC floats on water surface and traps all emitted gases (Figure 4). The embedded CH4 sensor will

measure in situ CH4 mixing ratio when the chamber is closed. In the sampling state, the AFC will collect the accumulated gas from the chamber headspace, flush then restart a new accumulation period. During sampling state, four sub-processes are run, including (Figure 4a) rinsing tubes and valves, (Figure 4b) gas sampling, (Figure 4c) venting the chamber by inflating the balloon and opening the chamber and (Figure 4d) closing chamber by balloon deflation and restarts a new accumulation period. These sub-processes are described in detail in Paper II.

Figure 4: Simplified illustration of the automatic flux chamber (AFC) sampling cycle including (a) rinsing of tubes and valves, (b) gas sampling, (c) venting and (d) balloon deflation to close the chamber for a new accumulation period. See text for details. Note that the CH4 sensor is placed inside the flux chamber, the power supply units and the electronics were excluded in this figure and each sample vial also has an exhaust needle through which the pre-loaded brine can be pushed out by the gas sample (not shown in figure). The figure is not to scale.

Field deployment test:

The AFC was tested in Tämnaren, a shallow lake near Uppsala, Sweden. In situ

chamber headspace CH4 concentrations, measured by the CH4 sensor, and weather information such as atmospheric pressure and air and water temperature were recorded during 12 hours of deployment with two hour accumulation periods. Discrete gas samples were collected at the end of every accumulation period (Figure 5). The gas samples were analyzed

for CH4 concentration by GC-FID. Fluxes were calculated using the difference in CH4 concentration between air and last measurement point over two hours accumulation time. The

calculated fluxes from data recorded from the on board CH4 sensor signal were indistinguishable from results calculated using concentrations in automatically filled vials. -2 -1 CH4 fluxes were in the range of 0.021 to 0.043 mmol m h . The maximum linear CH4

fluxes from the CH4 sensor during time periods shorter than 2 hours were in the same range

as simultaneous 30 min manual CH4 flux measurements using identical nearby chambers. The results of short term manual fluxes were in the range of 0.023 to 0.134 mmol m-2 h-1.

25 air temperature

water temperature

20 CH vial 4 CH sensor 4

15 P air V Level (for units see legend) battery 10

5 15:00 18:00 21:00 00:00 03:00 06:00 09:00 time

Figure 5: Data from a deployed AFC unit at the lake Tämnaren over time including CH4 concentration measured by CH4 sensor in ppmv (black line), CH4 concentration in discrete sample bottles in ppmv(black circle), battery level of AFC in V (dashed black line), atmospheric pressure in psi (dashed gray line), water temperature in °C (black scatter) and air temperature in °C (gray scatter).

The developed AFC provides an alternative to estimate CH4 flux at air-water interface which can be used for long term measurements with less labor effort than conventional FCs.

4.3 Formation of H2 and CH4 by weathering of olivine at temperatures between 30 and 70°C (Paper III)

Methodology

Natural olivine sand (Forsterite 91, Fo91) from North Cape Minerals in Åheim, Norway was incubated in Milli-Q water and 2.2 mM bicarbonate buffer at three different temperatures; 30, 50 and 70°C. The incubated bottles were evacuated and flushed with CH4 free N2 repeatedly; the pressure in the bottle was equilibrated to atmospheric pressure and then autoclaved. (For full method details see Paper III)

Results and discussion

At all temperature levels, CH4 concentration increased linearly over time (Figure 6). The ANOVA test showed that MF rates were clearly difference between temperatures (p<0.0005, F=41.97) but no difference was found between buffer and water (p=0.42, F=0.71).

H2 also accumulated in the bottles but without significant sensitivity to temperature.

Figure 6: Accumulation of CH4 (in nmol) in the headspace of the incubation bottles as a function of time. a) shows the concentration of CH4 in the bottles containing 2.2 mM bicarbonate buffer a and b) shows the concentration of CH4 in the bottles containing only Milli-Q water . Numbers next to the regression lines denote the accumulation rates in mol/m2/s. All values are after control values are subtracted. Error bars denote ± 1 SD (n=3).

The possibility of release of potential CH4 already present in the olivine had to be addressed. Microscopic analyses of thin sections did not reveal any gas inclusions so CH4 release from the olivine crystal structure or from small fluid inclusions seem very unlikely.

The trapped CH4 in the crystal structure cannot be measured. Further, XPS analyses showed that the amount of available surface carbon was not large enough to form CH4 to the extent that was measured. Therefore, although we cannot exclude the contribution of thermogenic

MF from the available surface carbon contributed, it could not explain all CH4 production. The remaining and most likely its source is abiotic MF via FTT reactions.

All the components which are involved in FTT reactions (i.e. H2, CO2, CO, CH4, H2O and necessary catalysts), was detected in incubation bottles. FTT reactions are supported by catalysts such as native metals or oxides of Fe, Ni or Cr, which are common constituents of natural olivines. XRD and microscopy analyses of the olivine used in these experiments

revealed the presence of magnetite (Fe3O4) and chromite (FeCr2O4), both of which are known to have a catalytic effect on the FTT reaction.

4.4 The potential for abiotic and biotic MF fueled by olivine dissolution (Paper IV)

Methodology

The same incubation method as in Paper III was applied, but this study used 20 mM 13 - H CO3 buffer as incubation solution. This provided possibilities to trace formation of CH4 from 13-C labeled the bicarbonate (for full method details see Supplementary information of Paper IV). In parallel with an abiotic treatment, a mixture of 10 strains of hydrogenotrophic methanogens was added into incubation bottles. (For full method details see Supplementary information of Paper IV).

- In this experiment, HCO3 concentration was prepared to simulate early Earth’s ocean. -1.4 From palaeosol evidence, pCO2 was about 10 atm in the CH4-rich early atmosphere (Rye et al. 1995). Seawater pH could have been in the range of 6.5 to 7.5 (Kesler & Ohmoto 2006); - therefore, with the experimental temperature from 30 to 70°C, the HCO3 concentration of 20 mM was chosen.

Results and discussion

In the abiotic treatments, the CH4 concentration increased 10-fold from 0.004 to 0.04 -1 nmol golivine for the first three days at 30 and 50°C but not at 70°C. However, in the later phase of the incubations significant MF occurred at 70°C with an average abiotic formation -5 -1 -1 13 rate of 5.6 10 nmol golivine d . The abiotic δ CH4 signatures were in the range of -51.08 to

-42.09 ‰ (relative to the PDB standard) and were not correlated with either CH4 concentration or incubation temperature (Figure 7).

The accumulated CH4 concentration was low; therefore the link of reduction 13 - 13 dissolved H CO3 to CH4 could not clearly establish. The δ CH4 was in the range of

commonly found for thermogenic CH4 (Schoell 1988). This means that we cannot exclude

that some of the CH4 was formed from trace level hydrocarbons possibly present in the olivine.

6900

5900

4900

3900 (‰ PDB) 4 2900 CH 13 δ 1900

900

-100 0 2 4 6 8 10 12 14 16 CH concentration (ppm) 4

13 Figure 7: δ CH4 versus CH4 concentration including the abiotic treatments at 30°C (), 50°C (), and 70°C(○), and the biotic treatments at the same temperatures, respectively (♦, , and ●).

In the biotic treatments, the increasing of CH4 concentration (Figure 8) and strong 13 relationship between CH4 concentrations and δ CH4 (Figure 7) showed that olivine-water incubations supported biotic methanogenesis at 70°C for at least 182 days and also at 30°C 13 for about 30 days. There was no clear change in CH4 concentrations or the CH4 enrichment at 50°C. This study provides experimental evidence for the potential of olivine-water interactions to sustain biotic processes over extended time periods and that biotic methanogenesis apparently can overcome constraints limiting abiotic MF processes.

Figure 8: H2 concentrations (left) and CH4 concentrations (right) over time at different temperatures in the abiotic (), and the biotic () treatments. Each point represents the mean of five separate incubation flasks. The x-axis showing incubation day was logarithm transformed to better illustrate the data in the early phase of the experiment. Note the different scales at the y axis. See text for details.

5 Discussion

5.1 Biotic MF and MO from sediment

The result that potential MO rates were well correlated to potential MF rates at the in

situ temperature is consistent with observations of CH4 concentration dependence on biotic MO kinetics in different types of environments including soils, sediments and water columns (Bender & Conrad 1995; King 1997; Knief & Dunfield 2005; Moore & Dalva 1997; Segers 1998; Sundh et al. 2005). Together with results that biotic MF was enhanced by rising temperature, this implies that elevated temperatures will enhance MF rates which may cause

increased CH4 release until MO increases as well, as a response to higher CH4 levels. As

microbial processes contributed up to 69% of global atmospheric CH4 budget (Conrad 2009), the balance between MF and MO could be an explanation for the report of decreasing -1 -1 atmospheric CH4 growth rate from 12 ppb year in the 1980s down to 4 ppb year since 1999 (Bousquet et al. 2006). In further, the biotic MO kinetics, seemingly being a function of

the availability CH4 and electron acceptors could have contributed to regulate atmospheric

CH4 dynamics during global warming or cooling that has been occurring many times along Earth’s history. However, it is still necessary to study the sensitivity of MO to the changes in MF as well as the effects of transient periods of elevated temperatures on MF and MO rates.

5.2 CH4 emission measurement

The CH4 flux from a field test of this AFC is shown in Figure 9. The AFC

successfully collected gas samples and recorded in situ headspace CH4 concentrations as well as weather information of air and water temperature, and atmospheric pressure. The

integrated average flux over the two hour accumulation periods based on CH4 sensor data on one hand and on the automatically collected samples on the other hand corresponded well (Table 1, Paper II). Separate manually sampled chambers were used in parallel and nearby the AFCs but having shorter accumulation periods (ca 30 minutes). Maximum fluxes from

CH4 sensor data during short periods (about 20 – 30 minutes), corresponded well with simultaneous manual chamber measurements and were for a few occasions substantially higher than the 2 hour average (Table 1, Paper II). This illustrated a substantial temporal variability in fluxes which can be resolved using the AFC information, but not with the commonly used short term manual chamber measurements. Because wind conditions and thereby water turbulence and gas exchange across the water-atmosphere interface may be highly variable locally and over time scales of minutes to hours, not only ebullition but also

diffusive CH4 flux could potentially show significant temporal variability. This could lead to bias when 30-minute measurements are used only once per system.

Figure 9: Calculated CH4 flux over time of day with data from (♦) CH4 sensor highest linear flux detected during the 2h accumulation period, () Average CH4 sensor flux over 2h, () Average automatic sampling flux over 2h, () CH4 concentration of manual sampling of 30 minutes accumulation.

More field tests are required to confirm the performance of this AFC under various conditions and settings. The results, so far, showing the capability of automatically collecting

sample and measuring in situ CH4 emission, weather information including temperature and

pressure, indicates that this AFC can contribute to studies of CH4 emissions with advantage of well defined footprint, high temporal resolution, less labor requirement.

5.3 Abiotic MF from olivine-water interaction at temperatures between 30 and 70°C

In a low redox potential environment where H2 can be formed, inorganic carbon could - thermodynamically be reduced to CH4. With pH range from 8 to 9, HCO3 is the dominating dissolved inorganic carbonate form in the incubated solution. Hence, reaction R2 can be rewritten as:

− + -1 -1 HCO3(aq) + H (aq) + 4H 2(aq) = CH 4(aq) + 3H 2O(l) G° = -229.3 (kJ mol l ) (R2a)

This reaction is thermodynamically favorable at standard conditions. The olivine incubation with Milli-Q water and 2.2 mM bicarbonate buffer in Paper III lacked measurements of pH, CO2 and H2 over time. Hence, its free energy over time cannot be calculated. However, reaction R2a has enthalpy ΔH° of -273.76 kJ mol-1 that is strong exothermic, so R2a is more preferable at low temperature. This provides support for the 13 - accumulation of CH4 over time in the experiment. Due to the lack of tracer H CO3 , we

cannot define the origin of CH4 in low carbonate buffer experiments.

In the follow up experiment with 20 mM bicarbonate buffer (Paper IV), applying the same thermodynamic calculation as above (R2a), Figure 10 showed that ΔG’ (in the range from -82.02 to -21.07 kJ mol-1 l-1) was enough for abiotic MF naturally happening. However, 13 the δ CH4 signatures were in the range from -51.08 to -42.09 ‰ (relative to the PDB

standard) and were not correlated with either CH4 concentration or incubation temperature

(Figure 7). As the amount of accumulated CH4 was too low, we could not establish a link to 13 - the reduction of dissolved H CO3 .

From the chemical point of view, despite of favorable thermodynamics, it is difficult to achieve 8 electrons reducing CO2 to CH4 that could create large kinetic barrier and inhibit reaction R2a. There is no doubt about the crucial role of catalysts that can enhance reaction

rate of H2 reducing CO2 to CH4 (Jacquemin et al. 2010; Lunde & Kester 1973; McCollom & Seewald 2003a; Seewald et al. 2006; Thampi et al. 1987). So far, MF at low temperature only

occurred with special catalysts which can act as chemisorption for H2 and CO2 (Jacquemin et al. 2010; Thampi et al. 1987). In our experiment, if abiotic MF via reaction R2a could happen, it would preferably take place on olivine surface where H2 (explained below in part

5.5) and catalysts were available. CH4 accumulated in the experiments, it is not completely clear whether reaction R2a could happen or if CH4 is formed via other pathways. With the

available natural catalysts magnetite (Fe3O4) and chromite (FeCr2O4) on olivine surface (Neubeck et. al. 2011) and in temperature range of 30 - 70°C, intermediate products such as

HCHO or CH3COOH might potentially be formed instead of CH4 (Foustoukos & Seyfried 2004; McCollom & Seewald 2003a; b). Further studies outside the scope of this thesis would be necessary to determine detailed reaction pathways.

Figure 10: Calculated free energy of rxn R2a in experimental conditions.

- Comparing abiotic MF rates at different HCO3 concentrations from the results of Paper III and IV, these rates seem to be influenced by concentrations of inorganic carbon - with increasing MF rates at HCO3 concentrations from 0 to 2.2 mM but substantially lower - rates at 20 mM HCO3 at all temperature levels (Figure 11). These differences were in agreement with Jones et. al. (2010) who studied the effect of carbonate on MF at 200°C and - 300 bar (Jones et al. 2010). The decreasing of potential MF rate at high HCO3 concentration could be caused by the formation of carbonate minerals coating catalytic sites. This indicated the important role of carbonates in MF process. Therefore, it is of interest to discuss how the - HCO3 concentration can affect carbonate precipitation during olivine dissolution at different temperatures. Olivine dissolution can raise solution pH by the reaction of the solution with - brucite or dissolved silicate (McCollom & Bach 2009; Moody 1976b) that results in HCO3 2- - dissociated into CO3 . Under specific conditions of temperature and pH, the HCO3 can 2- either be dissociated into CO3 which lead to form carbonate precipitation or be converted to

CO2(gas) and escape out of solution (Figure 12). In paper IV study, experimental results of - blank bottles with only HCO3 solution (no olivine) showed pH values were in the range of 8.29 to 8.59 and independent of the temperature. The rising pH from initial pH7 was a result 2- -1 of glass wall released SiO4 (Si concentration ranged 0.13 – 5.10 μmol mL ). In the sample - bottles containing olivine and HCO3 solution, pH values were in the range of 8.14 – 9.09 and well correlated to temperature with Pearson correlation: r2 = -0.790 and p = 0.000 (Figure 13). This indicated that pH was clearly driven by the olivine dissolution. And this result was in agreement with thermodynamic models study that was done by McCollom and Bach

(2009). As an oxide acid gas, measured headspace CO2 concentrations were well correlated with pH at all incubated temperature with Pearson correlation: r2 = -0.603 and p = 0.001.The

CO2 concentrations were well correlated with incubation time but showed different net

concentrations at different incubated temperatures. Figure 13 showed CO2 consumption (negative net concentration) with pH > 8.8 in 30 and 50°C. This indicated high potential of carbonate formation at these temperatures. At 70°C, on the other hand, positive net CO2 accumulation with pH < 8.8 indicated low carbonate precipitation.

2- Using the measured headspace CO2 concentration and the pH of the solution, CO3 concentration was calculated using temperature corrected Henry’s law and acid dissociation 2- constants (Plummer & Busenberg 1982). The results indicated that CO3 concentrations were inversely proportional to temperature while concentration of cations (such as Ca2+, Ba2+, Al3+: needed for carbonate precipitation) from olivine dissolution was directly proportional to temperature (Brantley et al. 2008; Oelkers 2001; Wogelius & Walther 1992) (Figure 14 2- 2- illustrated ICP data and CO3 concentrations versus temperature at 315 days). At 30°C, CO3 concentration was high but e.g. Ca2+ concentration was low and vice versa at 70°C. According to our calculations based on data in Figure 14, the potential for carbonate precipitation had an optimum at 50°C and was enhanced during the last period of 30°C incubation when cation concentration had become high enough. Therefore, the formation of - carbonate precipitation was constrained by temperature and HCO3 , cation concentrations. In turn, the precipitation could affect MF kinetics by blocking catalyst sites. This can be an explanation for the difference in abiotic MF between Papers III and IV.

- Figure 11: Potential CH4 accumulation rates. Note: white bar represents for 0 mM HCO3 - experiments, light gray bar for 2.2 mM and dark gray bar for 20 mM HCO3 experiments.

100

80 CO2(g) 70

60 - CO-- HCO3 3

T (°C) 50

30 CO2(aq)

10 0 2 4 6 8 10 12 14 pH

Figure 12: Diagram showing predominance of carbonate species versus temperature and pH

9.2

9.1

8.9

8.8

8.7 pH 8.6

8.5

8.4

8.3

8.2 -300 -200 -100 0 100 200 300 -1 Net CO2 concentration (nmol g_olivine )

Figure 13: The net CO2 concentration versus pH at different incubated temperature. (♦) 30°C, () 50°C, () 70°C. Note: net CO2 was the subtraction of CO2 concentration from sample bottle to control bottle.

10 20 30 8 16 24 30 60 90 400 600 800

temp 45

Al (nmol/mL) 20

10 24

Ba (nmol/mL) 16

8 90

60 Ca (nmol/mL) 30

[CO32-] (nmol/mL)

3+ 2+ 2+ 2- Figure 14: Matrix plot of temp, ICP data of Al , Ba , Ca and CO3 concentrations at 315 incubated days.

5.4 Biotic MF from olivine-water interaction at temperatures between 30 and 70°C

Results indicated that olivine water interactions supported methanogenic activity throughout the whole experiment at 70 °C and initially at 30°C. The methane formation stopped in the later stage in the 30°C treatment and never started at 50°C. This pattern was

consistent between replicates. The correlation between CH4 concentration (Figure 8) and 13 relative abundance of δ CH4 (Figure 7) clearly show that CH4 was formed from H2 and - HCO3 , in turn showing that the methanogens successfully competed for H2 compared with alternative abiotic processes. Experimental results showed dissolved H2 concentration in the range of 15.6 nM to 22.5 nM from day 133 (when methanogens was amended) in the abiotic experiments; therefore, it should be, in theory, enough to sustain hydrogenotrophic methanogens (Chapelle et al. 2002; Kral et al. 1998; Stevens & McKinley 1995). However, to

maintain the methanogenic activity, H2 had to be produced continuously by olivine dissolution processes. As describe above, carbonate precipitation was likely different at

various temperatures. In line with the discussion above, this could have constrained H2 formation, in turn, affecting methanogenic activity differently at different temperature. Biotic methanogenesis was most striking at 70°C and that could be a result of low carbonate

precipitation and higher potential for H2 formation. The methanogenic activity ceased towards the end of the biotic incubation period at 30°C and could did not start at 50°C in line with the potential for carbonate precipitation at 30 and 50°C as described above. Paper IV

empirically shows the potential of olivine-water interactions to support methanogenic activity under the studied conditions.

In summary, olivine-water interaction in the temperature range of 30 - 70 °C showed - limited capacity to support abiotic MF at HCO3 levels to be expected in nature but could clearly support biotic MF (Figure 15). Therefore, under habitable condition on the young

Earth, the abiotic CH4 flux from reduction of CO2 could have been kinetically inhibited. The thermogenic CH4 flux could have been an abiotic CH4 source based on hydrocarbon materials distributed from hydrothermal systems, but otherwise, the work in this thesis indicate that abiotic MF would have been restricted to specific areas of the seafloor with high temperature - and suitable catalytic potential, or areas with lower temperature but also lower HCO3 levels on early Earth (McCollom & Seewald 2007; Shock 1990). Rather than sustaining abiotic MF, olivine-water interaction at favorable habitable temperatures may have generated enough H2 to sustain biotic MF, especially at 70°C. This is consistent with the suggestion of high temperature theories for origin of life (Holm 1992; Massimo 2000; Wächtershäuser 2006) as temperatures (around 70°C) could eliminate carbonate precipitation, favor reactive surfaces on minerals to produce H2, and at the same time allow cells to survive. The presence of active methanogenic metabolism may have acted as a natural catalyst, selectively producing CH4 by reducing CO2 with H2, and enhancing MF rate. Therefore, methanogenic metabolism once evolved likely contributed to the green-house effect and the regulation of the early Earth’s climate. Further, the results of this thesis indicated that early Earth seafloor at temperatures in the range studied here were probably favorable for the development of many life forms because the availability of H2 and the widespread ability to use H2 as an energy source among microorganisms (Enrich-Prast et al. 2009). The results of Paper IV opens up interesting and important possibilities for future studies of the relative contribution of abiotic and biotic MF to atmospheric methane concentrations early on in the earth history, and also about anaerobic - 2- MO with nitrite (NO2 ) and sulfate (SO4 ) that could occur early in the reducing atmosphere and contribute to regulate early Earth climate (Ettwig et al. 2010; Valentine & Reeburgh 2000).

Figure 15: CH4 accumulation rates at different temperatures in the experiment in Paper 4. For the 50 and 70 °C treatment rates were calculated over the period from the introduction of methanogens to the end of the experiment (day 133 to 315). For the 30°C treatment the period with significant detectable accumulation (day 133-164; see Figure 8) was used. White bars with dashed border lines denote the abiotic treatment, and gray bars show the biotic treatment.

5.5 H2 formation from olivine-water interaction at temperatures between 30 and 70°C

Paper III and IV showed that H2 was formed and played an important role for MF. The

variation of H2 concentration over experimental time and its potential to support biotic MF is

discussed here. The H2 formation via Fe-bearing mineral-water interaction has been observed in both nature and experiments under various conditions (summarized in Table 1) but the detailed reaction pathways and their stoichiometry are presently not well-constrained. One pathway has been explained as released Fe2+ from fayalite component in olivine being

oxidized predominately to magnetite (Fe3O4) and leading to production of H2 from H2O (Hellevang 2008; Moody 1976a) as follows:

+ 2+ Fayalite(s) + 4H (aq) = 2Fe(aq) + SiO2(aq) + 2H 2O (R5)

2+ + -1 -1 3Fe(aq) + 4H 2O(l) → Fe3O4(s) + H 2(aq) + 6H (aq) G° = 187.55 (kJ mol l ) (R6)

In our experiment, mössbauer spectroscopy analyses did not detect serpentine, Fe3O4 3+ and Fe2O3 in any formed secondary minerals, but it revealed the amount of Fe on olivine

increased about 1% compare to initial olivine (Neubeck et. al. in prep). This result was in

agreement with suggestions of McCollom & Bach (2009) and Klein et. al. (2009) that Fe3O4 is only formed when temperature was higher than120°C. Therefore, reaction pathways with 2+ oxidized dissolved Fe to Fe3O4 or Fe2O3; and related production of H2 may not be the most

likely mechanism to explain the H2 produced at temperatures from 30 to 70°C.

As an attempt to understand the H2 formation, a potential mechanism is suggested based on experimental data. With the experiment conditions, if released Fe2+ was oxidized,

Fe(OH)3 would be a main product at pH >8 (reaction showed in R7) instead of Fe3O4 (Figure 16 shows the dominant compounds that might be formed in the solution according to Eh-pH values).

2+ + -1 -1 2Fe(aq) + 6H 2O(l) → 2Fe(OH )3(s) + H 2(aq) + 4H (aq) G° = 205.35 (kJ mol l ) (R7)

- Figure 16: Eh-pH diagram of ferrous iron with the present of 20 mM HCO3 at 30, 50 and 70°C.

2+ + The free energy of reaction R7, as calculated from the activity of Fe , H2(aq), H and -1 -1 unity activity for H2O(l) and Fe(OH)3(s), showed positive values from 8.77 to 21.7 kJ mol L (Figure 17) which is not favorable. If the activity of Fe2+ was also assumed to be unity, thermodynamic value of reaction R7 became negative (-75.8 - -45 kJ mol-1 l-1) (Figure 17). This indicates that R7 could be thermodynamically favorable as long as Fe2+ was not completely hydrated/dissolved. The completely dissolved Fe2+ would be incorporated with 2- CO3 to form siderite which has been revealed by Raman analysis (Neubeck et. al. in prep). This is in agreement with siderite formation at high dissolved inorganic carbon (Jones et al. 2+ 2010). When Fe on an olivine surface is exposed to a H2O molecule, it possibly could - - 3+ donate one electron to split H2O to an H radical and OH . The OH could associate with Fe

while two H radicals could form a H2 molecule. Therefore, we hypothesize that reaction R7 might occur naturally on olivine surface when solid phase Fe2+ is instantaneously exposed to

H2O upon to the dissolution of olivine. This is in agreement with suggestion about H2 formation on fresh reactive mineral surface by Parkes et. al. (2011) and Kita et. al. (1982).

Concomitantly with H2 formation, Fe(OH)3 can be either directly precipitated or combined 2- 2- with other dominant anions such as SO4 or CO3 and precipitated on the olivine surface.

These precipitates might block or diminish H2O diffusion to reactive sites that could hamper 3+ H2 formation. Further solid phase analysis is required to confirm about Fe minerals.

Figure 17: Calculated free energy of reaction R7 in experimental conditions at 30°C (♦), 50°C (), 70°C(●), and the free energy with assumption of unity activity of Fe2+ at the same temperature respectively (, , and ○).

In our experiment, H2 could, in principle, originate from the rubber septa and

hydrocarbon dissociation. The subtraction of H2 from sample to control resulted in positive value (Figure 8). It indicated that rubber septum can be ignored as a H2 source. Other studies

observed that hydrocarbon dissociation with Fe-bearing mineral assemblage produced H2 under hydrothermal conditions (Seewald 1994; Seewald 2001). However, the lack of TOC data and hydrocarbon species analysis did not allow us to investigate this possibility. If it is the case, it will open an interesting future study.

With amended methanogens, the active methanogens seem to have stimulated H2

formation which was observed in experiment at 30 and 70°C over long incubation times. H2 - was also consumed by methanogens to produce CH4, which also resulted in HCO3 consumption. Possibly, this could have reduced carbonate precipitations and indirectly enhanced H2 formation. At 50°C, H2 formation was hampered probably due to strong carbonate formation and cover of olivine surface as explained above.

6 Conclusions

From this thesis, it can be briefly concluded that:

- Biotic MF rates are highly sensitive to temperature while MO rates are not.

- Biotic MO rate is proportional to in situ MF rate across ecosystems showing that MO is primarily substrate dependent and may catch up with changed MF rates.

- CH4 emission may not simply increase by rising temperature because MO rates can increase and balance increased MF.

- Automatic flux chambers have the potential to be a powerful tool for retrieving high temporal resolution flux information, with less field work effort compared to commonly used approaches.

- CH4 can be abiotically formed from water-olivine interactions at 30-70 °C at bicarbonate concentrations of 2.2 mM or less. However at bicarbonate levels of 22 mM abiotic MF could not be detected. Hence, the kinetics of olivine-water interaction seems strongly affected by carbonate chemistry.

- Olivine-water interactions can support microbial MF for extended time periods (months) under low nutrient conditions. In fact, olivine-water interaction at 70°C most strongly supported the methanogenic metabolism which is interesting given high temperature theories for origin of life.

- H2 can be generated by olivine-water interaction in the temperature range 30 - 70°C.

- Active methanogenic metabolism seems to stimulate H2 production by olivine-water interactions.

7 Acknowledgments

Coming from Vietnam, I would like to thank Stockholm University for providing me opportunity to become a PhD student at the Department of Geological Sciences. I would like to take this opportunity to express my deep gratitude to my supervisors, my colleagues, my family and my friends who has supported me in making these studies enjoyable and possible. This thesis could not be like this if I don’t have:

My greatest supervisors:

David Bastviken. Thanks for giving me chances, as well as, brilliant advices to discover new science world. You always encourage and support me in new scientific challenges and train me become an independent scientist. Your beautiful life with connecting music and science, your happy family are my idol.

Patrick Crill. I greatly appreciate and love your fantastic guidance. I really admire your broad knowledge and experience. Freedom in doing research, you have inspired my work.

Nils Holm. I would like to acknowledge your gentle, valuable advice and support. I very adore your amazing broad network.

My teachers:

Lars Lundmark. There is no word that can express my gratitude. Come to Sweden, open knowledge in analytical instrument, keep working in instrumentation, enjoy Swedish nature and culture, there are too many things to remember what you have done for me and Analytical Chemistry Department in Vietnam.

Samuel Silverstein. We just have a short working time but it gives me a great chance to learn about PIC and programming. You help me to make my dream come true. Thank you very much.

I also would like to express my gratitude to all teachers in Vietnam for providing me with good foundation so that I can continue study in Sweden. Your dedication to be a teacher in hard conditions of Vietnam is my great life motivation.

My colleagues

Anna Neubeck. I cannot recall how many discussions, email and phone calls that have been carried out with you. We are dealing with a difficult project. This is never ending, it always open new questions. I have learned a lot while working with you. Thank you very much and wishing you all the best. Don’t worry, just relax!

Henrik Reyier. The automatic flux chamber cannot work in the field without you. Your great effort has worked out. I owned you a great thank for your valuable work.

Martin Wik. Thank you very much for your effort to setup and test the bubble counter in Abisko. Even there is still some problem, but we are dealing with problem and we call it a business of question.

Thank you very much to all the coauthors. Your knowledge contribution as well as English correction build up my confidence and help me bring forward all the publications. I learn a lot from our collaboration.

I also would like to thank all people in Astrobiology Graduate School at Stockholm University, especially Wolf Geppert. It is great memory and experience to know and travel with all of you. All the trips, from all brain storm meetings to visiting all Astrobiology Institutes, are wonderful and I can learn a lot. Hope we can stay in touch.

I want to send my great thank to all people who assisted me with administrative tasks at the department. I also want to send special thank to Arne and Anders at the workshop for your valuable help with computer and hard stuff. I would like to thank Klara, Heike, Malin for your help with laboratory work and help me to clean up my mess. Here in IGV, I have opportunity to work with, learn from and know many nice people. This is a great working environment. Thanks to all people at IGV for enjoyable time together, valuable advice and discussions.

My friends

I own a big thank to all of my friends in Stockholm and Umeå. I am very happy to meet you in Sweden. In Stockholm, Nguyen Huy Ha, not just a friend, you are also my C program teacher. Dr Tran Chi Thanh, not just a big brother, you are also my tennis coach. Viet Anh, my beer friend, Hai, Nga, be Lycka, anh Thuy, chi Ngoc, Giang, be Viet, Linh, Viet, anh Minh, chi Nga, chi Huong anh Leis. Thanks for sharing enjoyable times. In Umeå,

anh Son, chi Chung, Phuoc, Nhat, Chau, Lan, anh Huy, chi Thu, Lorrent, Mathias, Erik, Fredrik, thanks for warm feeling whenever I was in Umeå.

Family

I could not reach to this education level without my family who provide me unlimited and unconditional support. My parents who have dedicated their life for me and my sister I don’t know how to properly express my gratitude. I would have stopped my studying when I was in primary school if my mother did not take me back to school from Nintendo game shop. My father has introduced me to science world. Picture of my father working with numbers, writing and reading motivate me until now.

Tran Thi Thuy, my wife, how supportive you are during my studying time. This thesis is meaningless without you.

Nguyen Thanh Nam, it is too much for a 1 year old boy to understand what I write here. I am very happy to see my happy boy growing up every day.

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