The role of meteoric water recharge in stimulating biogenic methane generation: A case study from the Tempoku Coal Field, Japan

Accepted Manuscript The role of meteoric water recharge in stimulating biogenic methane generation: A case study from the Tempoku Coal Field, Japan

Shuji Tamamura, Takuma Murakami, Noritaka Aramaki, Akio Ueno, Satoshi Tamazawa, Alam A.K.M. Badrul, Shofa Rijalul Haq, Toshifumi Igarashi, Hideo Aoyama, Shinji Yamaguchi, Katsuhiko Kaneko PII: DOI: Reference:

S0166-5162(18)30466-X https://doi.org/10.1016/j.coal.2018.12.002 COGEL 3131

To appear in:

International Journal of Coal Geology

Received date: Revised date: Accepted date:

15 May 2018 2 December 2018 2 December 2018

Please cite this article as: Shuji Tamamura, Takuma Murakami, Noritaka Aramaki, Akio Ueno, Satoshi Tamazawa, Alam A.K.M. Badrul, Shofa Rijalul Haq, Toshifumi Igarashi, Hideo Aoyama, Shinji Yamaguchi, Katsuhiko Kaneko , The role of meteoric water recharge in stimulating biogenic methane generation: A case study from the Tempoku Coal Field, Japan. Cogel (2018), https://doi.org/10.1016/j.coal.2018.12.002

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ACCEPTED MANUSCRIPT The role of meteoric water recharge in stimulating biogenic methane generation: A case study from the Tempoku Coal Field, Japan Shuji Tamamura a,*, Takuma Murakami a , Noritaka Aramaki

a,†

, Akio Ueno a , Satoshi Tamazawa a ,

Alam AKM Badrul a,‡, Shofa Rijalul Haq b, Toshifumi Igarashi b, Hideo Aoyama c, Shinji Yamaguchi c, Katsuhiko Kaneko a

a

Horonobe Research Institute for the Subsurface Environment, Northern Advancement Center for Science

& Technology, 5-3 Sakae-machi, Horonobe-cho, Teshio-gun, Hokkaido, 098-3221, Japan Division of Sustainable Resources Engineering, Faculty of Engineering, Hokkaido University, North 13

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b

West 8, Kita-ku, Sapporo, Hokkaido, 060-8628, Japan

Mitsubishi Materials Corporation, 1-3-2, Otemachi, Chiyoda-ku, Tokyo, 100-8117, Japan.

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c

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* Corresponding author. Tel: +81-1632-9-4112; Fax: +81-1632-9-4113

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E-mail: [email protected] Present address †

National Institute of Technology, Kagawa College, 355 Chokushi-cho, Takamatsu, Kagawa, 761-8058,

‡

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Japan.

Petroleum and Mining Engineering Department, Military Institute of Science and Technology, Dhaka,

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1216, Bangladesh.

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Shuji Tamamura: [email protected]

Takuma Murakami: [email protected] Noritaka Aramaki: [email protected]

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Akio Ueno: [email protected]

Satoshi Tamazawa: [email protected] Alam AKM Badrul: [email protected] Shofa Rijalul Haq: [email protected] Toshifumi Igarashi: [email protected] Hideo Aoyama: [email protected] Shinji Yamaguchi: [email protected] Katsuhiko Kaneko: [email protected]

ACCEPTED MANUSCRIPT Abstract While meteoric water recharge is known to stimulate biogenic methane formation in shale and coal seams, the underlying mechanisms are currently unresolved. To this end, we conducted fieldwork in the Tempoku Coal Field, Japan. Pore water in core samples and well water in boreholes were analyzed for dissolved components and isotopic compositions. The hydraulic gradient was determined using values of borehole hydraulic head. Cl− concentrations (38 mg L−1 to 16,600 mg L−1 ), δ 18 O(H2O) values (−10.5‰ to −2.4‰), and δD(H2 O) values (−71.6‰ to −17.8‰) increased with depth (<200 m), indicating meteoric water

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recharge. The positive correlation between δ13 C(CH4 ) values (–74.9‰ to –42.7‰) and δ13 C(CO2 ) values (– 27.5‰ to +13.3‰), as well as between δD(H2 O) and δD(CH4 ) values (–264‰ to –200‰), in the core

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samples and groundwater, indicted in situ methanogenesis in the zone of mixing and migration of meteoric water and saline groundwater. Some of the pore water samples contained biogenic acetate, propionate, and

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succinate at remarkably high concentrations (~200 mg L−1 ), implying: (i) the thermodynamic inhibition of fermentation at fermentation sites, and (ii) spatial separation between the fermentation sites and

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methanogenesis sites. Planar fracture modeling indicates that at distances greater than a millimeter between fermentation and methanogenesis sites, the advec tive transport of fermentation products dominates rather than diffusive transport of those. Hence, meteoric water recharge would stimulate biogenic methane

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formation by inducing advective transport of the fermentation products, thus (i) relaxing the thermodynamic inhibition of fermentation at the site of the fermentation, and (ii) enhancing the rate of transport of the

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fermentation products to the site of methanogenesis. Keywords:

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Advection.

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Coalbed methane (CBM); Pore water; Organic acids; Meteoric water recharge; Biogenic methane;

ACCEPTED MANUSCRIPT 1. Introduction Energy requirements continue to increase despite conventional hydrocarbon reserves becoming depleted; leading to the development of unconventional hydrocarbon reservoirs such as shale gas and coalbed methane (CBM) (Dong et al., 2016; Vedachalam et al., 2015). These gases originate from the thermal (>100 °C) decomposition of buried sedimentary organic matter (SOM) or from the activity of methanogenic archaea at shallower depths (<3000 m) and cooler temperatures (< 80 °C) (Schoell, 1983; Strąpoć et al., 2011). Since the late 1980s, studies related to the intensive mining of these deposits have

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revealed a more significant contribution of biogenic methane than previously expected (Faiz and Hendry, 2006; Martini et al., 1996; Scott et al., 1994; Strąpoć et al., 2011). Various works have indicated that

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biogenic methane is formed after geological uplift of the basin to a shallower, cooler (< 80 °C) position (e.g., Martini et al., 1996; Scott et al., 1994). This later biogenic methane is called secondary biogenic methane,

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while biogenic methane that formed at the beginning of SOM diagenesis (peat stage) is called primary biogenic methane (Scott et al., 1994).

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Secondary biogenic methane often forms where meteoric water and saline groundwater mix (Flores et al., 2008; Martini et al., 1996; McIntosh et al., 2002; Rice et al., 2008; Scott et al., 1994). In the mixing zone, the deuterium enrichment of groundwater (i.e., δD(H2 O)) generally increases with depth,

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because δD(H2 O) of saline groundwater are often higher than those of meteoric water (e.g., Martini et al., 1996). Biogenic methane generated in groundwater takes at least some of its hydrogen atoms from the ambient water (Balabane et al., 1987; Burke, 1993; Schoell, 1980; Valentine et al., 2004; Waldron et al.,

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1999), resulting in a positive correlation between the deuterium enrichment in methane (i.e., δD(CH4 )) and δD(H2 O). Such a correlation has been observed in many shale gas and CBM deposits in the mixing zone of

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the meteoric water and saline groundwater (Golding et al., 2013; Vinson et al., 2017), indicating the importance of meteoric water recharge in stimulating in situ methanogenesis (Flores et al., 2008; Scott et al., 1994). However, the underlying mechanisms associated with this stimulus are poorly understood.

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Scott (1999) proposed the concept of microbially enhanced coalbed methane (MECoM or MECBM) to augment CBM production by imitating and enhancing the natural processes of secondary biogenic methane generation through the introduction of nutrients (bio-stimulation) and/or microbes (bioaugmentation) into CBM wells. Laboratory batch experiments investigating MECBM produced biogenic methane (> 1 cm3 g−1 coal) relatively quickly (< 1 year) from coal as the sole carbon source (e.g., Fallgren et al., 2013; Green et al., 2008; Harris et al., 2008; Jones et al., 2008, 2010; Papendick et al., 2011; Ulrich and Bower, 2008; Wawrik et al., 2012). A pilot-scale field test of MECBM, however, did not produce biogenic methane as much as these experimental expectations (Ritter et al., 2015), indicating an incomplete understanding of the mechanisms of in situ biogenic methane generation (Colosimo et al., 2016; Davis and Gerlach, 2018). To expand this area of research, a field study was undertaken in the Tempoku Coal Field, Japan. Both pore water in core samples and well water from boreholes were analyzed for major ions, organic acids, and stable isotopes. Dissolved gases from the well water samples and sorbed gases from core samples were

ACCEPTED MANUSCRIPT also analyzed for their composition and stable isotopes. The groundwater hydraulic gradient in the sampling location was determined using values of borehole hydraulic head. Coupling these results with planar fracture modeling leads us to suggest a potential role of meteoric water recharge in secondary biogenic methane generation in shale and coal seams. 2. Materials and methods 2.1. Study site

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Hokkaido, the northernmost of Japan’s main islands (Fig. 1a), is located near the Asian continent in the northwestern Pacific Ocean. The basement of northern Hokkaido is a Cretaceous accretionary prism

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(the Gakubuchi Group) that developed as part of an ancient island arc–trench system. The Gakubuchi Group is unconformably overlain by the Oligocene Magaribuchi Formation of mostly marine origin, which in turn

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is unconformably overlain by the non- marine Soya Formation of the middle Miocene. The Soya Formation (340–400 m thick) comprises alternating layers of mudstone and sandstone, and contains at least 10 layers of

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inter-bedded lignite (1–2 m thick) that form the largest lignite deposit in Japan (the Tempoku Coal Field: recoverable resources, 109 ton) (Fig. 1a; Suzuki, 2010). This formation is unconformably overlain by the middle Miocene Onishibetsu Formation (50–200 m thick) that is dominated by marine sandstone. The

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Onishibetsu Formation is unconformably overlain by the middle Miocene Masuhoro Formation (350– 1800 m thick), which comprises marine turbidite sediments. Fig. 1b summarizes the stratigraphy of these formations.

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Along the geological boundary between the Soya and Onishibetsu formations, three vertical boreholes, 25-1 (63.0 m), 25-2 (167.6 m), and 25-3 (203.7 m) were drilled in Sarufutsu village from 2013 to

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2015 (Fig. 1c) on a hillside beside a stream (Fig. 1d). The wells were screened to ensure that the well water was derived solely from the No. 4 lignite seam (Fig. 2). The vitrinite reflectance (Ro ), carbon, and ash contents of the lignite are typically 0.30–0.40, 65%–70% (dry, ash free basis), and 5%–15%, respectively.

2.2. Core samples

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More information for these lignite samples are provided by Aramaki et al., (2017).

After the careful removal of dirt from the surface, each core sample (4.5 cm in diameter) collected from the boreholes (25-1, 25-2, and 25-3) was immediately wrapped with clingfilm, vacuum packed, and stored in a freezer (−20 °C) in the laboratory until pore water extraction. The porosity, pore size distribution, total organic carbon (TOC) content, and sulfur content were determined for all samples. 2.3. Pore water Pore water from the sediment core samples was extracted under compression at 70–140 MPa in a cylindrical stainless mold equipped with a water conduit (Fig. S1; Supplementary data). Just before compression, the core was manually pulverized (< 2 mm) in a mortar as swiftly as possible (<4 min) to allow it to fit in the mold (Fig. S1). This procedure, however, did not recover pore water in the lignite core.

ACCEPTED MANUSCRIPT Pore water was automatically filtered (< 0.5 μm) during the compression (Fig. S1) and analyzed for the concentrations of major ions (e.g., Cl−, organic acids). δ18 O(H2 O) and δD(H2 O) of the pore water were also determined. Sample temperatures remained at room temperature throughout the compression phase, indicating that potential thermal alteration of pore water during this compression was negligible. The concentrations of components in the pore water may have been underestimated owing to dilution by relatively pure drilling fluid. If the contamination had been significant, the Cl− concentration, for example, would have been highly scattered depending on the extent of contamination. However, the Cl− concentration and other values (e.g., δD(H2 O)) showed consistent variations with depth (see the following

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results), implying limited contamination by drilling fluid.

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2.4. Well water

Water samples from the 25-1 and 25-3 boreholes were collected by a bailer submerged near the

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depth of the No. 4 lignite seam. Water from the 25-2 borehole was sampled anaerobically maintaining an argon atmosphere inside the well (Ueno et al., in press). Regardless of the sampling method, the well water

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was extensively pumped (more than 2.5 times the well volume) before sampling to ensure pristine groundwater was collected. The pH of the well water in the boreholes was measured on site by a portable pH meter (HM-20P, DKK-TOA) immediately after sample collection. The well water was filtered (< 0.45 μm),

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and chemical analyses were conducted in the laboratory within one day of sampling. The samples were also analyzed for δ18 O(H2O), δD(H2 O), alkalinity, and tritium concentration. In the following descriptions,

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“groundwater” refers to both “pore water” and “well water”, which are distinct from each other. 2.5. Gas

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A glass vial (30 mL or 120 mL) was filled to about 80% volume with a pulverized core sample (< 2 mm), and each was tightly sealed with a rubber septum and aluminum cap. The air in the vial was substituted by argon through the septum before the vial was heated at 80 °C for 5 days to accelerate the

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desorption of gas from the core sample to the headspace of the vial. Any potential for microbial activity to alter the headspace gas composition was unlikely at such high temperature (Wilhelms et al., 2001), yet the temperature was low enough to prevent thermogenic gas generation from SOM during such a short period of time (Cramer, 2004; Pepper and Corvi, 1995; Quigley and Mackenzie, 1988). After heating, the vial was cooled at room temperature and the headspace gas was used for compositional and isotopic analyses. When the well water was brought up to ground level, tiny bubbles emerged, indicating that gases were being released from the well water under atmospheric pressure. To estimate in situ dissolved gas concentrations, a glass vial (120 mL) in the presence of 2.5 mL of concentrated Hg(II)Cl2 (6.8 g L−1 ) for sterilization was filled with the well water and sealed with a rubber septum and aluminum cap. Then, 15 mL of the water in the vial was substituted with argon at atmospheric pressure, following the method of Kampbell et al. (1989). From the headspace gas composition, in situ dissolved gas concentrations were estimated by theoretically correcting the effect of effervesce on the dissolved gas concentrations during well water sampling (i.e., the atmospheric sampling method; Tamamura et al., 2018).

ACCEPTED MANUSCRIPT 2.6. Bioenergetics The Gibbs energy (ΔGr (T)) for hydrogenotrophic methanogenesis (4H2 + CO 2 → CH4 + 2H2 O) in the well water is calculated from

∆𝐺𝑟 (𝑇 ) = ∆𝐺𝑟0 (𝑇) + 𝑅𝑇ln (

(CH4 (aq))

) , (1) 4 (H2 (aq)) (CO2 (aq))

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where ΔGr0 (T) [J mol−1 ] is the standard Gibbs energy of the reaction, R [8.314 J K –1 mol−1 ] is the gas constant, T [K] is absolute temperature, and (CH4 (aq)), (H2 (aq)), and (CO 2 (aq)) are in situ dissolved CH4 , H2 , and CO 2 concentrations [mol L−1 ], respectively. Similarly, ΔG r (T) for acetoclastic methanogenesis

γHCO3 − (HCO3−)(CH4 (aq)) + 𝑅𝑇ln ( ) , (2) γCH3COO− (CH3 COO − )

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∆𝐺𝑟 (𝑇) =

∆𝐺𝑟0 (𝑇)

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(CH3 COO − + H2 O → CH4 + HCO 3 −) in the well water is calculated from

where (CH3 COO −) and (HCO 3 −) are in situ CH3 COO − and HCO 3 − concentrations [mol L−1 ], respectively,

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and γCH3COO− and γHCO3− are activity coefficients for (CH3 COO −) and (HCO 3 −), respectively. Table S1 (Supplementary data) lists the values for ΔGr0 (T), concentrations, and activity coefficients relevant to these

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equations. 2.7. Analysis

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Total carbon (TC) and inorganic carbon (IC) contents in the core sample were determined using a TOC analyzer (TOC-VCSH, Shimadzu) equipped with a SSM-5000A solid sample combustion unit. Their relative standard deviations (RSDs) for the measurement were < 0.7% and < 1.1%, respectively. The TOC

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content of a core sample was derived by subtracting its IC content from the TC content (TOC = TC − IC). Sulfur content in the core sample was determined by energy dispersive X-ray fluorescence spectrometry (SPECTRO XEPOS, Ametek), with an RSD of < 0.1%. The porosity and pore size distribution of the core samples were determined by a mercury porosimeter (AutoPore V 9600, Micromeritics). Here, the contribution from larger pores (> 120 μm in diameter) was neglected as an artifact, because they could be formed during sample shaping before undertaking the measurements. The porosities of some core samples were also measured by helium gas expansion porosimeter (UPore 300, Core Laboratories) for comparison. Alkalinity was measured by titration of the water sample with 0.01 mol L−1 H2 SO4 until the pH became 4.30. The dissolved organic carbon concentration in the well water samples was measured by the TOC-VCSH, with an RSD of < 0.3%. The tritium concentration in the well water samples was measured by a low-background liquid scintillation counter (Quantulus 1220, PerkinElmer) after electrolytic enrichment; the RSD of this measurement was < 2.3%. Concentrations of cations (Na+, NH4 +, K +, Mg2+, and Ca2+) and

ACCEPTED MANUSCRIPT anions (Cl−, Br−, SO 4 2−, formate, acetate, propionate, oxalate, malonate, and succinate) were determined by ion chromatography (ICS-1000, Dionex, for the cations; 761 Compact IC, Metrohm, for the anions) with RSDs of < 0.7% for the cations and < 3.5% for the anions. Values for δ18 O(H2 O) and δD(H2 O) were determined by cavity ring down spectroscopy (LWIA DLT-100, Los Gatos Research) with RSDs of < 0.1‰ and < 0.5‰, respectively. The activity coefficients γCH3COO− and γHCO3− in Eq. (2) were calculated from the ionic concentrations in the well water (Table S2 ; Supplementary data) using the extended Debye–Hückel equation (Drever, 1997). Gas concentrations of N 2 , O 2 , CH4 , and CO 2 were determined by a gas chromatograph (GC-14B,

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Shimadzu) equipped with a thermal conductivity detector with an RSD of < 1.7%. The H2 gas concentration was determined by a gas chromatograph (TRA-1, Round Science) equipped with a reduction gas detector;

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the RSD was < 0.5%. δ13C(CH4 ) and δ 13C(CO2 ) were determined by gas chromatography–combustion– isotope ratio mass spectrometry (GC–C–IRMS; Trace GC Ultra GC Isolink Delta V Advantage, Thermo

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Fisher Scientific) with a measurement error of < 0.2‰. δD(CH4 ) was determined by a continuous flow (CF)–IRMS system (Trace GC and Delta V, Thermo Fisher Scientific) with a measurement error of < 1.0‰.

𝑅x − 𝑅std ) × 1000, (3) 𝑅std

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δ= (

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The isotopic composition is expressed in delta (δ) notation:

where Rx and Rstd are defined as the ratio of the heavy to the light isotope in the sample and standard,

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respectively. δ13 C measurements are relative to the Pee Dee Belemnite, and δD and δ18O measurements are

3. Results

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relative to the Standard Mean Ocean Water (Sharp, 2007).

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3.1. Porosity and pore size distribution

Both mercury porosimetry and helium gas expansion porosimetry gave comparable porosities (15%–25%) for the sediment core samples, but the porosity of the lignite core sample measured by the former (7%–21%) tended to be lower than that by the latter (22.3%) (Table 1). Considering that the mercury porosimeter neglected the contributions from larger pores (>120 μm), while the helium gas expansion porosimeter did not, the discrepancy may imply the presence of such pores in the original lignite. Fig. 3 shows representative pore-size distributions of lignite and sediment samples with the smallest (0.0067 μm) and largest (2.13 μm) median pore sizes. Lignite pores were predominantly > 1 μm and < 0.1 μm in size. Pores in the sediment with the smallest median pore size were mainly in the range 0.001–0.1 μm, while the sample with the largest median pore size had pores mainly in the range 0.1– 10.0 μm. The other sediments showed roughly bimodal pore-size distributions, with a smaller peak at 0.001– 0.1 μm and a larger peak at 0.1–10.0 μm.

ACCEPTED MANUSCRIPT 3.2. Geochemical characteristics of water and gas components Cl− concentrations in groundwater (Fig. 4), δ18 O(H2O) values (Fig. 5a), and δD(H2 O) values (Fig. 5b) generally increase with depth from 38 mg L−1 to 16,600 mg L−1 , −10.5‰ to −2.4‰, and −71.6‰ to −17.8‰, respectively. The correlation coefficients of the Cl− concentrations with the δ18O(H2 O) and δD(H2 O) values are 0.88 and 0.89, respectively, and the correlation coefficient between the δ18 O(H2 O) and δD(H2 O) values is 0.99 (Fig. 6). All vial headspace gas compositions showed CH4 and CO 2 as the major gases (Fig. 5c,d; Table S3; Supplementary Data). The CH4 concentrations in the presence of lignite (3.2%–10.1%) and well water

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(21.6%–44.0%) were generally more than an order of magnitude higher than those in the presence of sediment (0.003%–0.91%). The CO 2 concentrations in the headspace were highly variable (from below the detection limit (D.L.) to 15.9%), with no clear correlation with sample type. Headspace δ13 C(CH4 ) (Fig. 5e),

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δ13 C(CO 2 ) (Fig. 5f), and δD(CH4 ) values (Table S3) generally increased with depth across all samples, from

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−74.9‰ to −42.7‰, −27.5‰ to +13.3‰, and −264.0‰ to −200.0‰, respectively. Formate (D.L.–67.0 mg L−1 ), acetate (D.L.–212.0 mg L−1 ), propionate (D.L.–183.0 mg L−1 ),

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oxalate (D.L.–13.9 mg L−1 ), malonate (D.L.–5.7 mg L−1 ), and succinate (D.L.–194.0 mg L−1 ) concentrations in pore waters (Fig. 7; Table S4; Supplementary Data) were highly variable, with no discernible correlation with depth (Fig. 7) or core lithology (Table 1). The concentrations of the majority of organic acids were

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below the detection limit in all samples from well waters; only acetate concentrations could be measured (<2.5 mg L−1 ; Fig. 7; Table S4).

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4. Discussion

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4.1. Hydrology

The non- marine origin of the Soya Formation and the marine origin of the Onishibetsu Formation were confirmed from the TOC (mg g−1 ) and sulfur (mg g−1 ) content ratio (C/S) of the core samples (Fig. S2;

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Supplementary data). Regardless of this geological distinction, the Cl− concentration (Fig. 4) in groundwater tends to increase with depth through the Onishibetsu Formation to the Soya Formation. The positive trend in Cl− concentration with depth indicates two-component mixing between meteoric water (low Cl− concentration) and saline groundwater (high Cl− concentration). Considering the absence of halite around the sampling location, the saline groundwater likely originates from ancient marine water, migrating to the non- marine Soya Formation in the sampling location. The marked increase in Cl− concentration from the Onishibetsu Formation to the Soya Formation (Fig. 4) suggests at least a partial stratigraphic control on groundwater movement. The linear correlation between δ18 O(H2O) and δD(H2 O) values (Fig. 6) also shows that the groundwater is a mixture of the two endmembers: (isotopically depleted) meteoric water and (isotopically enriched) saline groundwater. The isotopic composition of fresh well water from the 25-1 borehole, however, does not match with the “present” local meteoric water line (Fig. 6). A similar discrepancy in isotopic composition between fresh groundwater and present meteoric water was measured in fresh groundwater at

ACCEPTED MANUSCRIPT the bottom of an alluvial formation in Horonobe town (Ikawa et al., 2014), near our sampling location (Sarufutsu village). The age of this Horonobe fresh groundwater is presumed to be 12 Ka –42 Ka (i.e., the last glacial age), based on the hydrogeochemical information collected at Horonobe town (Ikawa et al., 2014). Hence, the fresh groundwater at our sampling location may include such old meteoric water. Nuclear weapons testing between 1950 and 1970 has produced significant amounts of tritium in the atmosphere, which subsequently ‘washed out’ as precipitation (Clark and Fritz, 1997). The presence of tritium in groundwater above 0.8 tritium units (TU, 1 TU = 3 H/1 H = 10–18 ) can indicate relatively modern recharge from such precipitation (Clark and Fritz, 1997). Both the fresh (25-1, 0.07 ± 0.02 TU) and saline

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well water (25-2 and 25-3, < 0.03 TU) had tritium present at less than 0.8 TU, implying no recent recharge. This is consistent with the above argument that the fresh groundwater in the sampling location would

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include “old” meteoric water.

Using a geometric method as described in Schwartz and Zhang (2003), the hydraulic heads of

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boreholes 25-1 (EL 54.60 m), 25-2 (EL 47.16 m), and 25-3 (EL 44.79 m) give a hydraulic gradient of 0.034 towards the northwest. This corresponds to the direction towards the stream from the mountain ridge (Fig.

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1d) and indicates a topographic influence on the hydrology. Even if saline well waters from 25-2 and 25-3 are derived from an aquifer different to that supplying fresh water to 25-1, the hydraulic gradient of the saline groundwater will exceed 0.012 (calculated by dividing the difference in head of the 25-2 and 25-3

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wells (2.37 m) by the horizontal distance between the two wells (195 m)). These relatively high hydraulic gradients, together with the geochemistry presented above, suggest that fresh meteoric water, old meteoric hydraulic gradient.

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4.2. Organic acids in groundwater

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water, and saline groundwater would flow at shallow, moderate, and deep depths, respectively, along the

Groundwater at temperatures above 80 °C in oil fields often contains organic acids at concentrations as high as 10,000 mg L−1 (Carothers and Kharaka, 1978; Fisher, 1987; Means and Hubbard,

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1987; Surdam and MacGowan, 1987). Such acids are plausibly formed in situ by abiotic thermal degradation of kerogen (Franks et al., 2001). Although biogenic (i.e., fermented) organic acids in pore water occur typically at lower concentrations than those of the “thermogenic” organic acids, they can be present at moderately high concentrations in pore water, as reported for Cretaceous sediment of the Atlantic coastal plain (~320 mg L−1 ; Chapelle, 2001; Chapelle and Bradley, 1996), Miocene sediments of Blake Ridge (~900 mg L−1 ; Wellsbury et al., 1997), and a North Sea tidal flat (~70 mg L−1 ; Heuer et al., 2006). Lower concentrations (< 10 mg L−1 ) of biogenic organic acids have also been reported in both modern sediment (Barcelona et al., 1980; Ebenå et al., 2007; Hines et al., 1994; Knab et al., 2008; Michelson et al, 1989; Novelli et al., 1988; Xiao et al., 2009) and ancient sediment (Heuer et al., 2006, 2009; McMahon and Chapelle, 1991; Shipboard Scientific Party, 2003). The subsurface temperatures at the sampling depths (< 12 °C) in this study exclude the possibility of “thermogenic” organic acids in the water samples. This suggests that the maximum concentrations of the organic acids measured in our pore water samples (Fig. 7) are comparable to the maximum concentrations of

ACCEPTED MANUSCRIPT biogenic organic acids reported in previous studies (described above). Conversely, organic acid concentrations in well waters were either undetectable or were very low (acetate) (Fig. 7). The well water was naturally supplied from lignite seams to the borehole along the hydraulic gradient, and thus likely percolated from larger pores of higher permeability in the lignite. In contrast, pore water in the core samples was extracted under compression, and so was likely derived from smaller pores of lower permeability in the sediments. In Cretaceous sediment of the Atlantic Coastal Plain, Chapelle and Bradley (1996) and McMahon and Chapelle (1991) found that organic acids were present at relatively high concentrations in smaller pores (clayey sediments), rather than in larger pores (sandy sediments) of the strata. The positive correlation

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between organic acid concentrations and pore size reported in these studies is consistent with the findings of

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the present study. 4.3. Source of methane

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Methane with δ13 C(CH4 ) < −60‰ is simply classified as biogenic (Bernard et al., 1976; Schoell, 1980; Whiticar et al., 1986), but methane with higher δ 13 C(CH4 ) values has several possible origins: (i)

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inclusion of thermogenic methane (Bernard et al., 1976; Schoell, 1980; Whiticar et al., 1986 ); (ii) methanogenesis in a partially closed environment (Brown, 2011; Golding et al., 2013; Strąpoć et al., 2011); (iii) methanogenesis in the presence of significant amounts of substrates (Penning et al., 2005; Valentine et

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al., 2004); (iv) a predominance of acetoclastic methanogenesis rather than hydrogenotrophic methanogenesis (Goevert and Conrad, 2009; Krzycki et al., 1987; Penning et al., 2006; Valentine et al., 2004); and (v) microbial oxidation of methane in subsurface environments (Coleman et al., 1981; Templeton et al., 2006;

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Whiticar and Faber, 1986). In the present study, about 60% of δ 13C(CH4 ) values were > −60‰ (Fig. 5e; Table S3), indicating the need for a careful examination of the methane’s source.

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Possibility (iii) can be excluded because the substrates are not so abundant in subsurface environments (Hoehler et al., 1998; Jackson and McInerney, 2002). Possibility (v) requires that if δ13 C(CH4 ) values increase by the microbial oxidation of methane, δD(CH4 ) values should show an even greater

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increase, owing to the larger isotopic fractionation for D/1 H than for

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C/12 C during the microbial oxidation

of methane (Coleman et al., 1981; Holler et al., 2009; Kinnaman et al., 2007; Rasigraf et al., 2012). The lack of such a correlated increase of δ 13 C(CH4 ) with δD(CH4 ) (Table S3) excludes this possibility. The isotopic fractionation factor of 13 C between CH4 (gas) and CO 2 (gas) (αCO2(g)–CH4(g)),

𝛼CO2 (g)−CH4 (g)

1000 + 𝛿 13 C(CO2 ) = , ( 4) 1000 + 𝛿 13 C(CH4 )

is dependent on the type of methanogenesis, being 1.05–1.09 for hydrogenotrophic methanogenesis and 1.04–1.06 for acetoclastic methanogenesis in subsurface environments (Whiticar et al., 1986). For thermogenic gas, αCO2(g)–CH4(g) would be < 1.04 from Eq. (4) using common ranges of δ13 C(CO2 ) (−25‰ to −15‰; Andresen et al., 1995; Boreham et al., 1998; Gaschnitz et al., 2001; Scott et al., 1994) and δ13 C(CH4 ) (< –50‰; Bernard et al., 1976; Schoell, 1980; Whiticar et al., 1986) for thermogenic carbon dioxide and

ACCEPTED MANUSCRIPT thermogenic methane, respectively. When δ 13 C(CH4 ) increases by methanogenesis in a partially closed environment (possibility (ii)), δ13C(CO2 ) also increases by a mass balance requirement, thereby maintaining the constancy of αCO2(g)–CH4(g) in Eq. (4) (Blair, 1998; Brown, 2011). Such a co- increase was observed here (Fig. 5e,f), with the calculated αCO2(g)–CH4(g) from Eq. (4) (average ± standard deviation = 1.060 ± 0.011) falling in the range of hydrogenotrophic methanogenesis. Hence, the sampled methane with δ13 C(CH4 ) > −60‰ would imply hydrogenotrophic methanogenesis in a partially closed environment (possibility (ii)), rather than the inclusion of thermogenic methane (possibility (i)) or the predominance of acetoclastic methanogenesis (possibility (iii)).

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The δD(CH4 ) values of methane in the core samples and dissolved methane in the well water samples were positively correlated with the δD(H2 O) values of the pore water and well water, respectively

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(Fig. 8). The data fell mainly in the range δD(CH4 ) = δD(H2 O) − 180 ± 10, a typical relationship found when hydrogenotrophic methanogenesis predominates in in situ groundwater (Whiticar et al., 1986). This

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observation is consistent with the inference from αCO2(g)–CH4(g), as mentioned above. Hence, the main source of methane in the sampling location is likely hydrogenotrophic methanogenesis. The presence of

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Methanobacterium (a genus of hydrogenotrophic methanogen), as the dominant Archaea in the well water (Ueno et al., in press), further supports this consideration.

The larger contribution of hydrogenotrophic methanogenesis, rather than acetoclastic

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methanogenesis, in the sampling location implies an abundance of hydrogen relative to acetate, despite some pore water containing acetate at remarkably high concentrations (Fig. 7). The inference of the higher

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bioavailability of hydrogen is supported by bioenergetic considerations in the following section. 4.4. Bioenergetics

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The deviation of ΔGr (T) from zero is a measure of the disequilibrium of a reaction: the more negative the value of the deviation, the greater the excess of reactants relative to the equilibrium condition. The energetic requirement to synthesize adenosine triphosphate (ATP), a common form of energy stock in

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microbes, requires the ΔGr (T) of the ATP synthetic reaction to be more negative than −20 kJ mol−1 to −10 kJ mol−1 (Hoehler et al., 2001; Jackson and McInerney, 2002; Schink, 1997). In this study, the ΔGr (T) values for ATP synthesis by hydrogenotrophic and acetoclastic methanogenic reactions are calculated using Eqs. (1) and (2), respectively, for well water samples. Here, the in situ H2 (aq), CH4 (aq), and CO 2 (aq) concentrations were obtained by atmospheric sampling (Tamamura et al., 2018), which was corrected for the effects of effervescence on dissolved gas concentrations during well water sampling, under the assumption of the water being gas-saturated in situ. The value of ΔGr (T) was also calculated using dissolved gas concentrations in the sampled well water without such correction. The actual ΔGr (T) would be between the values derived from the corrected and uncorrected dissolved gas concentrations. The resultant ΔGr (T) values for acetoclastic methanogenesis, −16.1 kJ mol−1 to −7.2 kJ mol−1 (Fig. 9), were close to the thermodynamic limit (−20 kJ mol−1 to −10 kJ mol−1 ), as frequently encountered for methanogenesis in subsurface environments (Jackson and McInerney, 2002). On the other hand, ΔGr (T) values for hydrogenotrophic methanogenesis, −55.8 kJ mol−1 to −18.5 kJ mol−1 (Fig. 9), were more negative

ACCEPTED MANUSCRIPT than the thermodynamic limit, indicating greater bioavailability of hydrogen relative to acetate in the well water. This condition could make hydrogenotrophic methanogenesis the dominant source of methane in the lignite seams, in harmony with the discussion in Section 4.3. As argued in the following sections, these thermodynamic conditions favorable for methanogenesis would have been established by groundwater advection supplying substrates for methanogens. 4.5. Thermodynamic control of fermentation In an anaerobic environment, microbial transformation of SOM to methane generally proceeds in a

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stepwise manner: (I) hydrolysis of SOM by extracellular bacterial enzymes; (II) fermentation of the resulting dissolved organic matter to fatty acids, alcohols, and hydrogen; (III) secondary fermentation of those

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products to acetate and hydrogen, and (IV) the products are then utilized by methanogens (Alperin et al., 1994; Brüchert and Arnosti, 2003; Conrad, 1999; Morris et al., 2013; Schink, 1997). Because a sin gle

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microbe cannot carry out all the steps, various microbes must contribute to anaerobic SOM degradation (Schink, 1997).

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When a reaction approaches equilibrium, the net rate of the forward reaction declines owing to “thermodynamic inhibition” of the reaction (e.g., Dale et al., 2006). The fermentation reactions (steps (II) and (III)), which often have low equilibrium constants, tend to suffer from thermodynamic inhibition

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(Conrad, 1999; Heider et al., 1998; Schink, 1997), because the concentration ratio of (products)/(reactants) at equilibrium is low for such reactions. Culture experiments have indeed shown that the fermentative degradation of organic molecules (e.g., fatty acids and benzoate) are suppressed by increasing the

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concentrations of the products (e.g., hydrogen and acetate) (Ahring and Westermann 1988, Chin and Conrad, 1995; Fukuzaki et al., 1990; Rothfuss and Conrad, 1993; Van Lier et al., 1993; Warikoo et al., 1996).

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To prevent thermodynamic inhibition, fermentation products (especially hydroge n and acetate) must be consumed by respiratory microbes (e.g., methanogens) in the system (Schink, 1997). If the fermentation bacteria are located close (< 10 μm) to the respiratory microbes, their mutual interaction is

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called syntrophy (Ishii et al., 2005; McInerney et al., 2008; Morris et al., 2013; Stams and Plugge, 2009). When fermentation products are actively consumed by respiratory microbes, hydrogen and acetate concentrations in groundwater are usually lower than 0.2 µg L−1 (Chin and Conrad, 1995; Hoehler et al., 1998; Lovley and Goodwin, 1988; Rothfuss and Conrad, 1993) and 10 mg L−1 (Christensen et al., 2000; Rothfuss and Conrad, 1993; Warikoo et al., 1996), respec tively. However, pore water in some sediments contains these products at higher concentrations (Chapelle, 2001; Chapelle and Bradley, 1996; Heuer et al., 2006; Wellsbury et al., 1997; this study), implying the relative inactivity of respiratory microbes in these pores (Chapelle and Bradley, 1996; McMahon and Chapelle, 1991). For the small pores (<1–10 µm in diameter) in sediments, Chapelle (2001) suggested that some mechanisms inhibit microbial respiratory activity relative to bacterial fermentation activity, resulting in elevated concentrations of the fermentation products in these pores. Here, at least some types of fermentation likely suffered from thermodynamic inhibition.

ACCEPTED MANUSCRIPT 4.6. Thermodynamic control of methanogenesis When methane from hydrogenotrophic methanogenesis predominates in shale and coal seams, not only the isotopic fractionation factors (αCO2(g)–CH4(g) and αH2O(l)–CH4(g)) but also the clumped isotope (i.e., multiply-substituted isotopologues) composition of methane often coincide with those at theoretical equilibrium (Gutsalo, 2008; Penning et al., 2005; Schoell, 1980; Stolper et al., 2015; Valentine et al., 2004; Vinson et al., 2017; Wang et al., 2015). These observations indicate that (hydrogenotrophic) methanogenesis frequently proceeds close to the thermodynamic equilibrium in subsurface environments (Gutsalo, 2008; Penning et al., 2005; Schoell, 1980; Stolper et al., 2015; Valentine et al., 2004; Wang et al., 2015) at the

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bioenergetic limit (ΔGr (T) values of −20 kJ mol−1 to –10 kJ mol−1 ) (Hansen et al., 2001; Hoehler et al., 1998, 2001; Jackson and McInerney, 2002; Jakobsen et al., 1998). Hence, either an increase in reactant (i.e.,

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hydrogen and acetate) concentration or a decrease in product (i.e., methane) concentration is required for methanogenesis to proceed thermodynamically.

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In gas-saturated groundwater, the CH4 (aq) concentration is stabilized at its solubility limit; this commonly occurs in shale gas and CBM deposits, and is also shown in this study (Section 2.5). Here, the

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methanogenesis is thermodynamically possible only by an increase in the reactants’ (H2 (aq) and acetate) concentrations, rather than a decrease in the product’s (CH4 (aq)) concentration, because the CH4 (aq) concentration is stabilized at the solubility limit. This contrasts with the case during fermentation, whereby

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removal of products relaxes the reaction’s thermodynamic inhibition (Section 4.5). The next section relates this difference in the thermodynamic regulation of fermentation and methanogenesis to the role of meteoric

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water recharge for secondary biogenic methane generation. 4.7. Role of meteoric water recharge in the generation of secondary biogenic methane

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Vitrinite reflectance (Ro ) is a common indicator of the thermal maturity of SOM: from 0.2 for peat to 4.0 for anthracite (Taylor et al., 1998). Microbes in sediments are generally sterilized at above 80 °C (Strąpoć et al., 2011; Valentine, 2011; Wilhelms et al., 2001), which roughly corresponds to the thermal

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maturity of SOM, as Ro > 0.7 (Selley, 1998). Meteoric water recharge can aid secondary biogenic methane generation by introducing microbes into a thermally sterilized basin ( Ro > 0.7) after its geological uplift to shallower and cooler environments (Hamilton et al., 2014; Scott et al., 1994; Strąpoć et al., 2011). Alternatively, biogenic methane generation in response to meteoric water recharge has also been observed in thermally unsterilized basins (Ro < 0.7); e.g., the Michigan Basin (Martini et al., 1996, 1998, 2008), Illinois Basin (Martini et al., 2008; McIntosh et al., 2002), and Powder River Basin (Flores et al., 2008; Rice et al., 2008). Therefore, some other effects could also be responsible for biogenic methane generation in response to meteoric water recharge (Flores et al., 2008). In this study, the presence of organic acids at remarkably high concentrations in some pore water samples (Fig. 7) suggests the relative inactivity of respiratory microbes (that would consume the fermentation products) in these pores, where at least some types of fermentation could suffer from thermodynamic inhibition. Here, removal of the fermentation p roducts, either by advection, diffusion, or a combination of the two (i.e., dispersion), would contribute to relaxing the thermodynamic inhibition of the

ACCEPTED MANUSCRIPT fermentation. The relative magnitudes of advective and diffusive transport are evaluated using a model fermentation product (acetate) and model pore system (planar fracture) as follows. Let d [m] and l [m] be the aperture and length of a planar fracture, respectively (Fig. 10a). Let ΔP [Pa] be the pressure drop of water flowing through the fracture. Then, the average flow velocity (v [m s−1 ]) of the water through the fracture is given by (Lucia, 1983) 𝑑 2 ∆𝑃 𝑣= ( ) , ( 5) 12𝜂 𝑙

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where η [N s m−2 ] is the viscosity of water (8.9 × 10−4 N s m−2 at 25 °C). Using this velocity, the advective

𝑐𝑑 2 ∆𝑃 = 𝑐𝑣 = ( ) , ( 6) 12𝜂 𝑙

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where c [mol m−3 ] is the acetate concentration in the water.

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𝐹advective

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flux of acetate through the fracture per unit area (Fadvective [mol s−1 m−2 ]) is

For diffusive flux, let c [mol m−3 ] be the acetate concentration at the left edge of the fracture (i.e.,

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the site of fermentation) (Fig. 10b). Let the acetate diffuse along the fracture (length l, aperture d) to its right edge (i.e., the site of methanogenesis), where the acetate concentration is assumed to be zero due to consumption by the methanogens. This leads to an acetate concentration gradient along the fracture of c/l

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follows Fick’s first law of diffusion:

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[mol m−4 ]. Then, the diffusive flux of acetate through the fracture per unit area (Fdiffusive [mol s−1 m−2 ])

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𝑐 𝐹diffusive =– 𝐷 ( ) , (7) 𝑙

where D [m2 s−1 ] is the diffusion coefficient of acetate, set here as 7.9 × 10−10 m2 s−1 , which is the average of values reported by McMahon and Chapelle (1991) (2.76 × 10−10 m2 s−1 ) and Michelson et al. (1989) (1.30 × 10−9 m2 s−1 ).

Subsurface microbes inhabit pores wider than 1 μm (Männik et al., 2009; Rebata-Landa and Santamarina, 2006). Hence, the model assumes fracture apertures (Fig. 10a and b) of at least this size: 1, 5, 10, and 20 μm, which are within the range common for sediment pore diameters (e.g., Liu et al., 2017; Yang et al., 2017; this study) and coal cleats (e.g., Busse et al., 2017; Karacan and Okandan, 2000; Laubach et al., 1998; Weniger et al., 2016; this study). The hydraulic gradient through the fracture is assumed to be 0.01 (i.e., 1 m/100 m), corresponding to a ΔP/l of 98 Pa m−1 in Eq. (6). From this, the Fadvective/Fdiffusive ratio is illustrated as a function of the fracture length (l) in Fig. 10c. It is not dependent on the acetate concentration (c), because the term cancels in the ratio of Eqs. (6) and (7). In our sampling location, the hydraulic gradient (0.034, Section 4.1) is more than three times greater than the model assumption (0.01). Hence, the

ACCEPTED MANUSCRIPT contribution from Fadvective relative to Fdiffusive is expected to be greater in the study area than in the model evaluation below. From Fig. 10c, the advective flux of acetate exceeds its diffusive flux when the fracture length is longer than 9 cm, 4 mm, 900 μm, and 300 μm for fracture apertures of 1, 5, 10, and 20 μm, respectively. Hence, thermodynamic inhibition of fermentation at the site of the fermentation will be relaxed by advective (rather than diffusive) transport of the fermentation products when the sites of fermentation and methanogenesis are separated by millimeters or more. The occurrence of such spatial separation of the two sites is supported by the presence of organic acids at high concentrations (>10 mg L−1 ) in some pore water samples (Fig. 7). To be specific, let the site of the fermentation be a cubic pore of 5 μm side length

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(corresponding to a volume of 1.3 × 10–16 m3 ; Fig. 11) connected to the site of methanogenesis through a planar fracture of aperture 1 μm, width 5 μm, and length 10 μm (Fig. 11). Let the acetate concentration at the

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sites of fermentation and methanogenesis be 10 mg L−1 (i.e., 1.7 × 10−1 mol m−3 ) and zero, respectively (Fig.

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11), resulting in an acetate concentration gradient along the fracture of (1.7 × 10−1 [mol m−3 ])/(10 × 10−6 [m]) = 1.7 × 104 [mol m−4 ]. Given a diffusion coefficient of acetate of 7.9 × 10−10 m2 s−1 , the diffusive flux of acetate through the fracture per unit area is 1.3 × 10−5 mol m−2 s−1 according to Eq. (7). Hence, the

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“absolute” diffusive flux of acetate through the fracture is 1.3 × 10−5 [mol m−2 s−1 ] × (5 × 10−6 [m] × 1 × 10−6 [m]) = 6.5 × 10−17 [mol s−1 ]. This flux is maintained only when the production rate of acetate at the

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cubic fermentation site is 5.2 × 10−1 mol m−3 s−1 (i.e., absolute diffusive flux/pore volume), which is over three orders of magnitude higher than the reported range of subsurface microbial metabolism (10 −11 to 10−4 mol m−3 s−1 ) (Brüchert and Arnosti, 2003; Hines et al., 1994; Michelson et al., 1989; Novelli et al.,

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1988; Wellsbury et al., 1997). For the calculated fermentation rate to be of plausible magnitude (i.e., < 10−4 mol m−3 s−1 ), the fracture length in this example (i.e., the distance between the sites of fermentation and

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methanogenesis) must be longer than 5.5 cm.

Considering that (1) smaller sediment pores would selectively harbor fermentation bacteria, rather than respiratory microbes (Chapelle, 2001; Chapelle and Bradley, 1996; McMahon and Chapelle, 1991); (2)

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relatively limited phylogenetic groups of fermentation bacteria are capable of syntrophic metabolism (Sieber et al., 2012) versus the diverse fermentation bacteria found in sediments and coals (e.g., Penner et al., 2010); and (3) microbes are distributed patchily (separated on scales of micrometers to centimeters) in subsurface sediments (Brockman and Murray, 1997; Carreira et al., 2013; Fredrickson et al., 1997; Krumholz et al., 1997; Rogers and Apte, 2004) with remarkable compositional heterogeneity (Fredrickson et al., 1991; Guo et al., 2015; Klein et al., 2008; Lawson et al., 2015; Nickels et al., 1981; Wei et al., 2013), the physical separation between the active sites of fermentation and microbial respiration (e.g., methanogenesis) by some distance (> millimeters) would not be rare, as indicated by analyses of Cretaceous sediment of the Atlantic Coastal Plain (Chapelle and Bradley, 1996; McMahon and Chapelle, 1991). When the active sites of fermentation and methanogenesis are suitably separated, the fermentation products will be transported mainly by advection, rather than diffusion, to the site of the methanogenesis, as mentioned above. Hence, groundwater flow (advection) induced by meteoric water recharge (Schwartz and Zhang, 2003) could stimulate biogenic methane generation by (i) relaxing the thermodynamic inhibition of

ACCEPTED MANUSCRIPT fermentation at the site of the fermentation, and (ii) enhancing the transport rate of the fermentation products to the site of the methanogenesis. Consistently, in situ formation of biogenic methane (Section 4.3) under thermodynamically favorable conditions (Section 4.4) has been revealed at our sampling location, where groundwater flows (Section 4.1). The importance of advection to the stimulation of microbial activity has also been reported in tidal sediments (Beck et al., 2009; Riedel et al., 2011; Seidel et al., 2012). 4.8. Implication for MECBM From the above discussion, shale and coal seams can plausibly have at least some fermentation

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and methanogenesis sites suitably separated. This study suggests that groundwater flow, induced by meteoric water recharge, could be a key factor for successful MECBM, relaxing the thermodynamic inhibition of

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fermentation at the site of the fermentation, and enhancing the rate of transport of the fermentation products

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to methanogens. Further research is required to test this mechanism. 5. Conclusions

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1. Pore water compositions, isotopic indices, and bioenergetic (ΔG r (T)) considerations confirm in situ biogenic methane formation in the zone where meteoric water and saline groundwater mix and flow in the Tempoku Coal Field, Japan.

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2. Some pore water contained biogenic organic acids such as acetate, propionate, and succinate at remarkably high concentrations (~200 mg L−1 ), implying: (i) the thermodynamic inhibition of fermentation at fermentation sites, and (ii) the spatial separation of the active sites of fermentation and methanogenesis.

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3. Meteoric water recharge could stimulate biogenic methane generation by inducing the advective transport of the fermentation products, thus (i) relaxing the thermodynamic inhibition of fermentation, and (ii)

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enhancing the transport rate of fermentation products to the site of methanogenesis. 4. The maintenance of groundwater flow will contribute to MECBM.

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Acknowledgements

This study was performed in the mining area of Mitsubishi Materials Corporation, which greatly supported this study by constructing boreholes, completing them for well water collection, and providing core samples. We are especially grateful to N. Ozawa, S. Yamamoto, T. Gonda, and H. Shigeyoshi of this company for their valuable contributions to this study. We also thank T. Yokoyama, S. Miyaike, S. Matsuo, and I. Takeda from Mitsubishi Material Techno Corporation for their critical assistance in the core and groundwater sampling. H. Wakahama from Dia Consultants Corporation helped with groundwater sampling. Professor T. Naganuma of Hiroshima University and Professor Y. Fujii of Hokkaido University provided valuable comments that assisted this research. We thank Dr. S. Shimizu, M. Akatsuka, M. Muramoto, T. Endo, H. Tada, M. Miyako, K. Nishizawa, T. Nakanishi, T. Higashikawa, E. Ando, and H. Takahashi of the Horonobe Research Inst. for the Subsurface Environment (H-RISE) for their assistance. This study was directed by the research project of H-RISE, prepared in 2011 by Dr. Y. Ishijima and initiated in 2012 by Y. Ohmi.

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Table 1 Lithology and porosity of the core samples. Depth Borehole Formationa Lithology [m] 15.3 S Clay–silt 21.0 S Silt 26.0 S Fine sand 32.6 S Silt–fine sand 25-1 39.5 S Fine sand 47.0 S Silt 52.0 S Silt 53.0 S Lignite 57.6 S Fine sand 10.6 O Silt 20.4 O Silt 30.5 O Silt 40.4 O Silt 51.0 O Medium–coarse sand 58.0 S Lignite 61.5 S Silt 70.0 S Silt–fine sand 81.5 S Fine sand 25-2 91.0 O Medium–coarse sand 100.5 S Fine sand 111.5 S Silt 120.5 S Fine sand 130.0 S Medium sand 135.6 S Lignite 142.4 S Lignite 150.0 S Silt 159.0 S Silt 163.4 S Lignite 16.0 O Silt 27.7 O Silt 40.0 O Silt 51.5 O Silt 61.5 O Medium sand 71.5 S Silt 79.6 S Silt 87.5 S Fine sand 104.0 S Medium sand 25-3 119.6 S Clay–silt 130.5 S Medium sand 143.6 S Silt 151.5 S Silt 160.0 S Lignite 161.0 S Lignite 183.0 S Silt–fine sand 190.5 S Silt 199.8 S Lignite 203.4 S Silt a

S and O represent the Soya Formation and the Onishibetsu Formation, respectively

Porosity [%] N.D.b 25.0 21.0 19.5, 18.5c 22.4 20.2 N.D. 8.40, 22.3c 16.6 26.2 25.7 25.0 25.8 27.0c 16.4 19.7 18.6 20.0, 23.8c 28.6c 20.8 20.2 23.2, 23.9c 22.9 N.D. 20.9 N.D. N.D. 7.35 28.5 N.D. N.D. 26.0 28.0c 18.9 18.6 21.1 24.8, 29.3c 17.5 23.4 25.0 N.D. 8.70 N.D. 19.9 20.8 8.40 22.5

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From helium gas expansion porosimetry

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b

ACCEPTED MANUSCRIPT Fig. 1 (a) Distribution of the Soya Formation. (b) Age, stratigraphic relationships, and paleoenvironment of the Soya Formation. (c) Geological and topographic (inset (d)) maps showing the locations of boreholes 25 -1, 25-2, and 25-3. Fig. 2 Geological section along the three boreholes. The location of the section is shown on Fig. 1c.

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Fig. 3 Pore size distribution of representative lignite and sediments with the largest (2.13 μm) and smallest

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(0.0067 μm) median pore sizes.

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Fig. 4

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Vertical profiles of Cl– concentration in pore water and well water for each borehole. Fig. 5

Vertical profiles of (a) δ18 O(H2 O) and (b) δD(H2 O) in pore water and well water, and vertical profiles of

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concentrations ((c) CH4 , (d) CO 2 ) and isotopic compositions ((e) δ13 C(CH4 ), (f) δ13 C(CO2 )) of headspace gas in a vial containing pulverized core sample or well water. Symbols indicate samples for different boreholes. Data for lignite core and well water are distinguished from those for sediment core and pore

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Fig. 6

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water, respectively, by grizzled and blackened symbols, respectively.

Relationship betweenδD(H2 O) and δ18 O(H2 O) for pore water and well water. The isotopic compositions of stream water flowing near the boreholes (Fig. 1d) are also plotted. The local meteoric water line (δ

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D(H2 O) = 7.0δ18O(H2 O) + 6.9) is from Teramoto et al. (2006). The symbols reflect samples from different boreholes. Data for well water are distinguished from those for pore water by blackened symbol. Fig. 7

Vertical profiles of formate, acetate, propionate, oxalate, malonate, and succinate concentrations in pore water and well water for each borehole. Data for well water are labeled with arrows. Fig. 8 Relationship betweenδD(CH4 ) from the core samples and well water with respect to δD(H2 O) from the pore water and well water, respectively. The symbols reflect samples from different boreholes. Data for well water are distinguished from those for core samples by blackened symbol. Fig. 9

ACCEPTED MANUSCRIPT Δ G r (T) [kJ mol–1 ] for hydrogenotrophic and acetoclastic methanoge nesis in well water. Values are calculated from the dissolved gas concentrations corrected and uncorrected for the effect of effervescence during well water sampling. ΔGr (T) for acetoclastic methanogenesis in 25-2 well water is not shown, because its acetate concentration was below the detection limit. Fig. 10 Acetate flux through a model fracture of length l and aperture d, via (a) advection and (b) diffusion. (c) Ratio of advective flux/diffusive flux for acetate through fractures of various length (l). For advective flux, the

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hydraulic gradient through the fracture is assumed to be 0.01. For diffusive flux, the acetate concentration at the methanogenic site (right edge of the fracture) is assumed to be zero owing to consumption by

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methanogens.

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Fig. 11

Hypothetical case with sites of fermentation (cubic pore of side length 5 μm) and methanogenesis connected through a planar fracture (aperture 1 μm, width 5 μm) of length 10 μm. The acetate concentration is assumed

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to be 10 mg L–1 at the site of the fermentation and zero at the site of the methanogenesis. The main text

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discusses the plausibility of the distance (10 μm) between the two sites.

ACCEPTED MANUSCRIPT Highlights Study of role of meteoric water recharge in forming biogenic methane in coal seams.

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Advective transport (AT) relaxes the thermodynamic inhibition of fermentation.

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AT enhances the provision rate of fermentation products to methanogens.

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Meteoric water recharge promotes AT, activating both fermentation and methanogenesis.

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