Organic Geochemistry 113 (2017) 262–282

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Organic Geochemistry

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Invited Review clumped : Progress and potential for a new isotopic tracer ⇑ Peter M.J. Douglas a,b, , Daniel A. Stolper a,c, John M. Eiler a, Alex L. Sessions a, Michael Lawson d, Yanhua Shuai a,e, Andrew Bishop f, Olaf G. Podlaha g, Alexandre A. Ferreira h, Eugenio V. Santos Neto h, Martin Niemann i, Arne S. Steen j, Ling Huang e, Laura Chimiak a, David L. Valentine k, Jens Fiebig l, Andrew J. Luhmann m, William E. Seyfried Jr. n, Giuseppe Etiope o,p, Martin Schoell q, William P. Inskeep r, James J. Moran s, Nami Kitchen a a California Institute of Technology, Geological and Planetary Sciences, Pasadena, CA 91125, USA b McGill University, Earth and Planetary Sciences, Montreal, QC H3A 0E8, Canada c University of California, Berkeley, Earth and Planetary Science, Berkeley, CA 94720, USA d ExxonMobil Upstream Research Company, Spring, TX 77389, USA e Key Laboratory of Petroleum Geochemistry, Research Institute of Petroleum Exploration and Development, PetroChina, Beijing, China f University of California-Riverside, Department of Earth Sciences, Riverside, CA 92521, USA g Shell Global Solutions International B.V., Grasweg 2, 1031HW Amsterdam, The Netherlands h Division of Geochemistry, PETROBRAS Research and Development Center (CENPES), PETROBRAS, Rua Horácio Macedo, Ilha do Fundão, Rio de Janeiro, RJ 21941-915, Brazil i Statoil ASA, Martin Linges Vei 33, Fornebu, Norway j Statoil ASA, R&T ET PSM, PO Box 7200, NO-5020 Bergen, Norway k University of California-Santa Barbara, Earth Science, Santa Barbara, CA 93106, USA l Institut für Geowissenschaften, Goethe Universität, Frankfurt am Main, Germany m New Mexico Institute of Mining and Technology, Earth and Environmental Science, Socorro, NM 87801, USA n University of Minnesota-Twin Cities, Earth Sciences, Minneapolis, MN 55455, USA o Istituto Nazionale di Geofisica e Vulcanologia, Sezione Roma 2, Roma, Italy p Faculty of Environmental Science and Engineering, Babes-Bolyai University, Cluj-Napoca, Romania q Gas Consult International, Berkeley, CA 94704, USA r Montana State University, Land Resources and Environmental Sciences, Bozeman, MT 59717, USA s Pacific Northwest National Laboratory, Richland, WA 99354, USA article info abstract

Article history: The isotopic composition of methane is of longstanding geochemical interest, with important implica- Received 26 June 2017 tions for understanding petroleum systems, atmospheric concentrations, the global car- Received in revised form 11 July 2017 bon cycle, and life in extreme environments. Recent analytical developments focusing on multiply Accepted 22 July 2017 substituted isotopologues (‘clumped isotopes’) are opening a valuable new window into methane geo- Available online 16 August 2017 chemistry. When methane forms in internal isotopic equilibrium, clumped isotopes can provide a direct record of formation temperature, making this property particularly valuable for identifying different Keywords: methane origins. However, it has also become clear that in certain settings methane clumped Methane measurements record kinetic rather than equilibrium isotope effects. Here we present a substantially Clumped isotopes Geothermometry expanded dataset of methane clumped isotope analyses, and provide a synthesis of the current interpre- Petroleum systems tive framework for this parameter. In general, clumped isotope measurements indicate plausible forma- Biogeochemistry tion temperatures for abiotic, thermogenic, and microbial methane in many geological environments, which is encouraging for the further development of this measurement as a geothermometer, and as a tracer for the source of natural gas reservoirs and emissions. We also highlight, however, instances where clumped isotope derived temperatures are higher than expected, and discuss possible factors that could distort equilibrium formation temperature signals. In microbial methane from freshwater ecosystems, in particular, clumped isotope values appear to be controlled by kinetic effects, and may ultimately be use- ful to study methanogen metabolism. Ó 2017 Elsevier Ltd. All rights reserved.

⇑ Corresponding author at: McGill University, Earth and Planetary Sciences, Montreal, QC H3A 0E8, Canada. E-mail address: [email protected] (P.M.J. Douglas). https://doi.org/10.1016/j.orggeochem.2017.07.016 0146-6380/Ó 2017 Elsevier Ltd. All rights reserved. P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 263

1. Introduction the isotopic composition of abiotic gas was considered to be typi- cally enriched in 13C, with d13C values higher than 25‰. More d13 Methane (CH4) is an important component of the Earth’s carbon recent data indicates that C values of abiotic methane in serpen- cycle. As the primary constituent of natural gas (90%; (Hunt, tinized ultramafic rocks can be as light as 37‰ and abiotic 1979), methane extracted from geological reservoirs accounts for methane from Precambrian shields can be even lighter (Etiope approximately 20% of total global energy use (IEA, 2015). Methane and Sherwood Lollar, 2013). is also the second most important long-lived (i.e., excluding Post-generation processes can also impart notable isotopic frac- vapor) atmospheric greenhouse gas, and on a molar basis traps 28 tionations in methane. For instance, biological methane oxidation times as much heat as on 100-year timescales leads to an increase in both d13C and dD(Alperin et al., 1988; (Myhre et al., 2013). More generally, as one of the most common Whiticar, 1999) of the residual methane, while atmospheric reac- fluid forms of organic carbon, methane has played an important tions involving OH or Cl that consume methane lead to a large role throughout Earth history, both in facilitating the movement increase in D/H ratios in the residual gas (Gierczak et al., 1997; of reduced carbon between different environments, and as a Saueressig et al., 2001). Diffusion of methane through a gas phase metabolite for biotic communities. It has often been suggested that results in lower d13C and dD values of the gas that has diffused and methane played a role in the origin of life on Earth (Urey, 1952; elevated values in the remaining methane (Krooss et al., 1992; Russell et al., 2010; McCollom and Seewald, 2013), and could be Zhang and Krooss, 2001; Chanton, 2005). a signal of life on other planets (Formisano et al., 2004; While conventional stable isotope ratios often provide valuable Krasnopolsky et al., 2004; Atreya et al., 2007). clues about methane sources and sinks, several factors limit their Given the importance of methane, methods for identifying for- diagnostic ability. The empirically defined fields for d13C and dD mation processes and transport mechanisms are of great value. In values of different methane sources (Fig. 1) are not sharply defined, particular, the isotopic composition of methane, including both and there are clear cases of overlap between them (Martini et al., stable (i.e., 13C/12C and D/H ratios) and radioactive isotopes (i.e., 1996; Prinzhofer and Pernaton, 1997; Martini et al., 1998; Horita 14C and T), has been widely used as a tracer for sources and sinks and Berndt, 1999; Valentine et al., 2004; Etiope and Schoell, (Martell, 1963; Schoell, 1980; Whiticar et al., 1986; Lowe et al., 2014). Moreover, a number of studies have questioned whether 1988; Quay et al., 1999; Whiticar, 1999). For example, methane the isotopic fields associated with hydrogenotrophic and fermenta- produced by the thermal breakdown of organic matter during oil tive methane on this plot are indicative of those pathways, or of and gas formation generally has 13C/12C(d13C) and D/H (dD) ratios environmental or biological variables (Sugimoto and Wada, 1995; higher than methane produced by microorganisms (Schoell, 1980; Waldron et al., 1998, 1999; Conrad, 2005; Penning et al., 2005). Whiticar et al., 1986)(Fig. 1). In addition, different pathways of Finally, differentiating the effects of generation and post- microbial methanogenesis are thought to produce distinctive iso- generation fractionations, combined with mixing of two or more topic fractionations. Methane produced via CO2 reduction (or methane sources with different isotopic compositions, can be chal- hydrogenotrophic methanogenesis) shows especially low d13C val- lenging (Prinzhofer and Pernaton, 1997; Martini et al., 1998; ues, and methane produced by fermentation (or fermentative Whiticar, 1999). Combining methane stable isotope data with gas methanogenesis) has particularly low dD values (Whiticar et al., concentration or radiocarbon (14C) data, or with stable isotope 1986; Whiticar, 1999)(Fig. 1). Finally, abiotic methane can be gen- measurements of co-occurring phases such as water, H2,CO2,or erated over a wide range of temperatures by magmatic and gas– ethane, can often help to resolve these ambiguities (Bernard water–rock reactions that do not directly involve organic matter et al., 1978; James, 1983; Coleman et al., 1995; Hornibrook et al., or microbes (Etiope and Sherwood Lollar, 2013; Etiope and 1997; Waldron et al., 1999; Townsend-Small et al., 2012). Never- Schoell, 2014). Until a few years ago, and based on limited data, theless, there is a clear need for additional tracers to help differen- tiate between various sources and post-generation processes. The clumped isotope composition of methane has great poten- tial to complement conventional measurements including both isotopic and gas composition measurements. ‘Clumped isotope’, as used here, refers to molecules with two or more rare, generally heavy stable isotopes (Eiler, 2007, 2013). For methane, this implies either a 13C and one or more D substitutions, or two or more D sub- stitutions, in the same molecule. For a population of methane molecules that are in isotopic equilibrium with one another, the abundance of multiply-substituted isotopologues relative to a stochastic (random) distribution is a function of temperature (Stolper et al., 2014a; Webb and Miller III, 2014; Wang et al., 2015). This relationship allows methane clumped isotope abun- dances to be used as a geothermometer to constrain gas formation temperatures (Stolper et al., 2014b; Wang et al., 2015). The tem- perature dependence of clumped-isotope abundances is not unique to methane and exists for all studied materials including

CO2 (Eiler and Schauble, 2004), carbonate-bearing minerals (Ghosh et al., 2006), O2 (Yeung et al., 2012), and N2O(Magyar et al., 2016). Clumped isotope abundance in systems that are not in internal isotopic equilibrium can also be used to provide con- straints on non-equilibrium processes including the chemical kinetics of various reactions, which we discuss here as well (Daëron et al., 2011; Saenger et al., 2012; Stolper et al., 2015; Fig. 1. Plot comparing methane d13C and dD values, after Etiope (2015) and Etiope Wang et al., 2015; Yeung et al., 2015). and Sherwood Lollar (2013), based on Schoell (1980) and new empirical data. MH – microbial hydrogenotrophic; MF – microbial fermentation; ME – microbial in In this review, we present a new and substantially expanded evaporitic environments. database of clumped isotope compositions for methane from a 264 P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 13 13 diverse set of formation environments, and discuss the implica- d13 ¼ Rsample RVPDB ð Þ C 13 2 tions of these new data alongside previously published datasets RVPDB (Ono et al., 2014; Stolper et al., 2014a, 2014b; Inagaki et al., where 2R and 13R are the ratios D/H and 13C/12C respectively. Delta 2015; Stolper et al., 2015; Wang et al., 2015; Douglas et al., values are commonly expressed as per mil (‰) values, which 2016; Young et al., 2017). The purpose of this review is to explore implicitly includes a multiplicative factor of 1000 (Coplen, 2011). the potential of clumped isotopes to decipher methane origins. We All of the previously unpublished clumped isotope data pre- begin with a brief overview of the two different measurement sented in this paper are combined measurements of the two techniques currently available. We then discuss the different pro- -18 isotopologues performed using the prototype Ultra, and cesses that control methane clumped isotope values. Finally, we are expressed using D18 notation (Stolper et al., 2014a): review the broad patterns of clumped isotope abundance in methane from different environments, compare these data with 18R D ¼ 1 ð3Þ conventional isotope and gas composition measurements, and dis- 18 18R cuss potential applications involving the atmosphere and the sur- face of other planets. where: ð½13CH Dþ½12CH D Þ 18 ¼ 3 2 2 ð Þ R 12 4 2. Analytical methodology ½ CH4

2.1. Measurement techniques for methane clumped The specified isotope ratios are measured from the corresponding ion beam current ratios, and are standardized by comparison with 18 ⁄ Two distinct measurement techniques were developed over the a gas of known isotopic composition. R is the ratio expected for past five years for methane clumped isotope analysis. The first a random distribution of isotopes among all isotopologues, and is 13 2 employs high-resolution dual-inlet mass spectrometers with an calculated using the measured R and R values for the sample electron ionization source. The first such instrument developed (Stolper et al., 2014a): was the Thermo MAT-253 Ultra (hereafter referred to as the 18R ¼ð62R2Þþð42R13RÞð5Þ ‘Ultra’), described in detail by Eiler et al. (2013). Methane clumped isotope compositions measured by the prototype version of the The prefactors 6 and 4 in Eq. (5) derive from the symmetry numbers D Ultra combine the abundances of the two mass-18 isotopologues of the mass-18 methane isotopologues (Stolper et al., 2014a). 18 13 12 values are reported as per mil (‰) deviations from a calculated ref- of methane ( CH3D and CH2D2), but distinguish the two mass- 13 12 erence frame, where 0‰ represents a random distribution of 17 isotopologues ( CH4 and CH3D) (Stolper et al., 2014a). Given 18 18 ⁄ the low natural abundance of D, the combined mass-18 ion current methane isotopologues (i.e., R= R ) – this is equivalent to the 13 D value of a gas internally equilibrated at infinite temperature. is primarily (98%) determined by the abundance of CH3D 18 (Stolper et al., 2014a). More recently, a prototype version of a As gases cannot, in practice, be equilibrated at infinite temperature, D larger-radius high-resolution isotope ratio mass spectrometer, all samples are calibrated against a laboratory standard with a 18 13 value of 2.981‰, as described by Stolper et al. (2014a). Most of the the Nu Instruments Panorama, that can routinely resolve CH3D 12 new data presented in this paper were measured during a period in and CH2D2 was developed (Young et al., 2016, 2017). A newer, D 18 production version of the Thermo Ultra employs an improved which we observed a linear dependence of 18 on R in heated gas beam-focusing and detection design that also allows it to resolve samples, i.e. the measured state of clumping depends slightly on the 13 the two mass-18 isotopologues (Clog et al., 2015). level of C or D enrichment, which it clearly should not. A correc- The other measurement technique employs long path-length tion for this dependence was applied as detailed by Douglas et al. laser spectroscopy using mid-infrared frequencies. Spectroscopic (2016). measurements of methane clumped isotopes were first performed We also discuss a smaller subset of previously published data using a difference-frequency-generation laser (Tsuji et al., 2012), measured using the TILDAS spectroscopy technique or the Nu but this technique gave relatively poor precision (20‰). More Panorama mass spectrometer (about 20% of the total dataset), D13 recently, Ono et al. (2014) developed a tunable infrared laser direct which are reported as CH3D values, as defined by Ono et al. D13 D absorption spectroscopy (TILDAS) method that can measure (2014). We have not converted CH3D values to 18 values. In 13 the case of methane inferred to have formed in isotopic equilib- CH3D abundance with greatly improved precision (0.25‰). This technique uses two quantum cascade lasers tuned to four different rium we primarily discuss the data in terms of equivalent temper- 12 13 12 13 ature (see below), in which case we employ the distinct isotopologues of methane ( CH4, CH4, CH3D, and CH3D). 12 temperature calibrations for each measurement. In the case of Measurement of CH2D2 is not currently possible with production 12 D version laser spectroscopy systems, but is possible in principle and kinetic fractionations, the contribution of CH2D2 to 18 values may be developed in the future. In addition, development of cavity is uncertain, and therefore an accurate conversion between these ringdown spectroscopy for methane isotopologues, including measurements is not straightforward. However, we expect that clumped isotope species, is ongoing (Bui et al., 2014). The sensitiv- such small differences are unlikely to influence the broad patterns ity and precision of the mass spectrometric and spectroscopic of abundance that we seek to outline here. D methods for methane clumped isotope analyses are broadly simi- 18 values can be related to equivalent temperature (which has lar. Inter-calibration of these two measurement techniques has physical meaning as an environmental temperature if the sample not yet been performed, and is a key priority for future research. has achieved internal isotopic equilibrium), via the equation (Stolper et al., 2014a): ! ! 2.2. Nomenclature 6 6 D ¼ : 10 þ : 10 : ð Þ 18 0 0117 2 0 708 2 0 337 6 Conventional carbon and isotopic compositions are T T expressed using delta notation relative to standard mean ocean 13 The analogous relationship for D CH3D is: water (VSMOW) and Pee Dee Belemnite (VPDB), respectively: ! ! 6 6 2 2 13 10 10 Rsample RVSMOW D CH D ¼0:0141 þ 0:699 0:311 ð7Þ d ¼ ð Þ 3 2 2 D 2 1 T T RVSMOW P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 265 as derived from the calculations of Webb and Miller III (2014).We and 4.7) determined using various methods. Methods for reservoir hereafter refer to such estimated temperatures as T18 or T13CH3D val- temperatures for the published Haynesville and Marcellus Shales, ues. When referring to both measurements together we refer to D18 Gulf of Mexico, Powder River Basin, Western Pacific, and Birchtree or T18, as this is the more general term. Mine samples were detailed previously. Briefly, temperatures from the Gulf of Mexico are borehole temperatures corrected using pro- prietary formulas (Stolper et al., 2014b), temperatures from the 2.3. Sample preparation Haynesville and Marcellus Shale are borehole temperatures with an upwards correction of 10% (Stolper et al., 2014b), temperatures Samples analyzed for this study were prepared using the proto- from the Powder River Basin and Birchtree Mine samples were col described by Stolper et al. (2014a, 2014b). In brief, mixed gas measured in associated formation (Bates et al., 2011; samples were introduced from either a steel cylinder, an aluminum Wang et al., 2015; Young et al., 2017), and temperatures from cylinder, or a glass serum vial sealed with a butyl stopper, into a the Western Pacific are estimated from the local geothermal gradi- glass vacuum line. Gas samples were first exposed to liquid nitro- ent (Inagaki et al., 2015). Shallow marine methane samples from gen to trap H2O, CO2, and H2S. The gases in the headspace (includ- the Santa Monica Basin and the Beaufort Sea (Fig. 2) were com- ing CH4,O2, and N2) were then exposed and transferred to a 20 K pared with measured bottom water temperatures (Stolper et al., cold trap, and residual gases, including He and H2, were pumped 2015; Douglas et al., 2016). away. The cold trap was then sealed, heated to 80 K, cooled to 45 Reservoir temperature measurements for the Qaidam and Son- K, and opened to vacuum to remove N2 and O2. This temperature gliao Basins (Supplementary Table) were based on long-term drill- cycling was repeated until < 2.67 Pa of gas remained in the cold stem tests, which is a robust method for determining virgin forma- trap at 45 K, corresponding to a purity of CH4 of 99.8% (Stolper tion temperatures (Hermanrud et al., 1991; Peters and Nelson, et al., 2014a). The cryostat was then heated to 70 K, and CH4 was 2012). Reservoir temperature estimates for the Milk River Forma- transferred to a PyrexTM breakseal containing molecular sieve (EM tion (Supplementary Table) were calculated using the local Science; type 5A) immersed in liquid N2. Prior to introduction to geothermal gradient (estimated using a basin model), the surface the mass spectrometer dual inlet, samples were heated using a temperature, and the depth of the sampled well. An example of heat gun or copper block set to 150 °C for 2–3 h to ensure mini- this methodology is provided by Nunn (2012). Aside from drill- mal isotopic fractionation when transferring CH4 from the molec- stem tests, the methods used for reservoir temperature estimation ular sieves (Stolper et al., 2014a). are approximate, and their uncertainties are not consistently quan- tified. Therefore we applied a conservative 20 °C error to these esti- 2.4. Natural gas reservoir temperature measurements and estimates mates (Peters and Nelson, 2012; see Sections 3.1 and 4.7). For temperatures derived from drill-stem tests we applied a 10% error. For a subset of samples we compared clumped isotope derived temperature estimates with independent measurements or esti- mates of the natural gas reservoir temperature (see Sections 3.1 3. Processes Controlling Methane Clumped Isotope Values

3.1. Formation or re-equilibration temperature

13 D18 or D CH3D values of methane in internal isotopic equilib- rium are predicted to vary as monotonic functions of the tempera- ture of equilibration (Fig. 2), as described by Eqs. (6) and (7) above (Stolper et al., 2014a; Webb and Miller III, 2014; Wang et al., 2015). In most environments internal isotopic equilibrium is not depen- dent on isotopic exchange between methane molecules, but instead is produced via isotope-exchange reactions with other

phases, including H2O, H2,orCO2. However, it is not necessary for methane to be in carbon or hydrogen isotope equilibrium with co-occurring molecules to achieve internal isotopic equilibrium, as long as hydrogen isotope exchange reactions between methane and other molecules are reversible, allowing the distribution of iso- topes in C-H bonds in methane to reach equilibrium. For more detailed discussions of the theory of equilibrium clumped isotope fractionation see Wang et al. (2004), Eiler (2007), Eiler et al. (2014), Stolper et al. (2014a), Wang et al. (2015), Young et al. (2017) and Stolper et al. (in press). The clumped isotope temperature dependence for methane formed in isotopic equilibrium is grounded in statistical thermody- namics (Ono et al., 2014; Stolper et al., 2014a; Webb and Miller III, 2014; Wang et al., 2015; Young et al., 2017) and has been empiri- cally validated using various methane samples having known equi- libration or formation temperatures (Fig. 2). These include Fig. 2. The relationship between D18 values and formation temperature for methane formed in internal isotopic equilibrium. The black line indicates the methane equilibrated using a metal catalyst (e.g., Ni or Pt) at tem- theoretical prediction (Stolper et al., 2014a). The data depicted are either naturally peratures between 150 and 500 °C(Ono et al., 2014; Stolper et al., occurring methane with well-constrained formation temperatures (See Section 2.4), 2014a; Wang et al., 2015; Douglas et al., 2016); methane from or experimentally derived methane (Stolper et al., 2014a, 2014b, 2015; Douglas pyrolysis experiments performed at 360 and 600 °C(Stolper et al., 2016). A similar relationship exists for D13CH D values (Ono et al., 2014; 3 et al., 2014b); and methane from environmental samples inferred Webb and Miller III, 2014; Wang et al., 2015). 20 °C error bars in formation temperature were applied to the Gulf of Mexico and Haynesville Shale samples, to have formed in isotopic equilibrium and with independently whereas x-axis errors for other data are smaller than the markers. constrained formation temperatures (Stolper et al., 2014b, 2015; 266 P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282

Douglas et al., 2016). Additional measurements of samples gener- oping clumped isotope equilibrium in thermogenic methane, and 13 ated between 50 and 150 °C would improve the D18- how this relates to d C and dD values (see detailed discussion in temperature calibration. Stolper et al., in press). The environmental samples shown in Fig. 2 are all from con- texts where methane formed at or near the sampling environment, and where significant migration of gas from other environments is 3.2. Kinetic isotope effects during methane generation unlikely (Stolper et al., 2014b, 2015; Douglas et al., 2016). For example, the Haynesville Shale is considered to be both the source Methane produced by methanogens in a number of pure culture and the reservoir for generated , and its current tem- experiments is characterized by clumped isotope values signifi- perature is within 17 °C of modeled maximum burial temperatures cantly lower (5.4 to 2.3‰; Fig. 3) than would be expected based 13 (Curtis, 2002; Stolper et al., 2014b). on the experimental temperatures. Such ‘low’ D18 and D CH3D

D18 values vary as a non-linear function of temperature (Fig. 2), values correspond to apparent formation temperatures that are and the typical precision of D18 measurements (0.25‰;1r) cor- either significantly higher (216–620 °C) than those of the actual responds to varying temperature errors. For example,±0.25‰ cor- growth temperatures (25–85 °C), or, in the case of negative D18 responds to an uncertainty of ±8 °C at an inferred temperature of values, do not correspond to any possible formation temperature 25 °C, whereas at an inferred temperature of 200 °C the corre- (Fig. 2)(Stolper et al., 2014b, 2015; Wang et al., 2015; Douglas sponding uncertainty is ±21 °C. The formation or re-equilibration of methane in internal iso- topic equilibrium, whether in the laboratory or environment, requires processes that allow the C-H bonds in methane to reversi- bly exchange. In a laboratory setting, nickel or platinum catalysts have been used to equilibrate methane C-H bonds at temperatures from 150 to 500 °C over a period of one to two days (Ono et al., 2014; Stolper et al., 2014a; Wang et al., 2015; Douglas et al., 2016). Methane from the Marcellus Shale with a measured well temperature of 60 ± 10 °C (See Section 2.4) yields a T18 value of 200 °C, which is generally consistent with inferred formation conditions. This suggests that methane formed at a higher temper- ature in isotopic equilibrium, consistent with modeled maximum burial temperatures, and was subsequently uplifted and stored for millions of years at lower temperatures without experiencing re-equilibration (Stolper et al., 2014b). This suggests that internal re-equilibration of methane either does not occur, or proceeds very slowly, at temperatures below 200 °C in shale reservoirs in the absence of significant amounts of metal catalysts. Internal isotopic equilibrium of methane in natural environ- ments may depend on factors other than storage temperature. For example, H-exchange rates (the main process driving equili- bration) may be enhanced by the presence of certain clay minerals or other ‘natural catalysts’ (Alexander et al., 1982, 1984; Horita, 2001). It has been argued that enzymatic catalysis during anaero- bic oxidation of methane by marine archaea induces carbon iso- tope equilibrium between CH4 and CO2 (Yoshinaga et al., 2014), and this process might also induce CH4 clumped isotope equilib- rium (Stolper et al., 2015). Conventional (dD and d13C) isotope values in thermogenic methane are generally thought to be controlled by kinetic isotope effects, as opposed to equilibrium fractionation (e.g. Sackett, 1978; Tang et al., 2000; Ni et al., 2011). However, the inference of clumped isotope equilibrium in thermogenic methane is not incon- sistent with kinetic isotope effects controlling its d13CordD values (Stolper et al., in press). This is because it is not necessary for D D13 a methane to fully equilibrate with an external carbon or Fig. 3. Plot of clumped isotopes values ( 18 or CH3D) vs. H2O-CH4 for microbial hydrogen-bearing phase in order for reversible isotope exchange methane samples. Deep subsurface microbial methane samples are not plotted since dD values are uncertain. Equilibrium values for D vs. a ,as reactions to equilibrate the distribution of isotopes within the C- H2O 18 H2O-CH4 calculated by Stolper et al. (2015) are shown by the solid black line. Samples

H bonds of methane. Evidence for clumped isotope equilibrium analyzed for D18 values are shown by solid markers, whereas samples analyzed for 13 in both pyrolysis experiments and natural gas samples has been D CH3D values are shown by open markers. The samples are categorized by found in several studies (Stolper et al., 2014b; Wang et al., 2015; environment and geographic region. Marine microbial methane samples plot near the equilibrium line, but samples from other categories exhibit a negative trend Young et al., 2017), including consistent results from the measure- a 13 with lower clumped isotope values and higher H2O-CH4 values than the equilibrium ment of two multiply substituted isotopologues ( CH3D and line. The dashed line indicates the predicted trend for decreasing enzymatic 12 CH2D2; Young et al., 2017). As discussed in the following section, reversibility of methanogenesis at 20 °C based on a model of kinetic isotopic effects some recent pyrolysis experiments have yielded clumped isotope (Stolper et al., 2015). For freshwater microbial methane there appear to be results indicating distinctive kinetic fractionations, in which case differences in this trend in different geographic regions. A single pure culture of a fermentative methanogen clearly deviates from the model prediction, with a low the T18 value is much higher than the experimental temperature D18 value relative to its aH2O-CH4 value. The cross in the lower left indicates (Shuai et al., in revision). Additional experimental and theoretical representative x and y error bars. Data from Stolper et al. (2015), Wang et al. (2015), studies are needed to better understand the mechanisms for devel- Douglas et al. (2016). P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 267 et al., 2016; Young et al., 2017). Similarly, microbial methane sam- To first order, the ‘reversibility of methanogenesis’ hypothesis pled from freshwater ecosystems, cow rumen, and serpentiniza- predicts that faster rates of methanogenesis (per cell), stimulated tion zones also yields either negative clumped isotope values or by larger chemical potential gradients between methane precur- values that are so low that they cannot plausibly be interpreted sors and methane, will lead to lower D18 values. While it is unclear as formation temperatures (Stolper et al., 2015; Wang et al., to what extent other environmental or biological variables modu-

2015; Douglas et al., 2016)(Fig. 3). These studies hypothesized that late this relationship, it is possible that D18 values could serve as an the low observed clumped isotope values are the result of kinetic indicator of specific growth rate of methanogens (Stolper et al., isotope effects arising during microbial methane generation. 2015; Wang et al., 2015). Such a proxy could improve understand- Specifically, these studies proposed that the low D values (either ing of the biogeochemistry and ecology of microbial methanogen- 13 D18 or D CH3D) result from the expression of kinetic isotope esis in different environments. effects during irreversible, enzymatically-catalyzed hydrogenation There are also preliminary indications that kinetic isotope of C from CO2 or CH3 groups. It has been hypothesized that the effects can be induced in high-temperature catagenetic reactions, degree of enzymatic reversibility dictates how low the D value will at least in the laboratory. While early hydrous pyrolysis experi- be, with the least amount of reversibility being linked to the lowest ments performed with shale samples produced D18 values that D values. We refer to this as the ‘reversibility of methanogenesis’ were within error of the predicted equilibrium value for the hypothesis, which is based on and consistent with earlier work experimental temperature (Fig. 2)(Stolper et al., 2014b), more relating D/H and 13C/12C isotope fractionation in microbial recent experiments performed with coals at faster heating rates methanogenesis to the chemical potential gradient between the have yielded D18 values lower than predicted for equilibrium, reactants, such as H2,CO2, or methyl groups, and products (CH4) and in some cases negative ‘anti-clumped’ values (Shuai et al., of methanogenesis (Valentine et al., 2004; Penning et al., 2005). in revision). While mechanisms for these deviations remain A clear relationship emerges for microbial methane in which incompletely understood, Shuai et al. (in revision) hypothesize

D18 values correlate with aH2O-CH4 (i.e., the hydrogen isotope frac- that the observed non-equilibrium fractionation is a result of tionation between methane and water; Fig. 3). This relationship breaking carbon-carbon bonds during cracking of aliphatic hydro- indicates that the same processes controlling the D18 values also carbons, whereas demethylation of kerogen produces methane set the dD value of methane relative to the source water. Impor- with equilibrium D18 values. Regardless of the mechanism, it is tantly, when D18 values indicate formation temperatures consis- clear that under some experimental conditions thermogenic tent with the environmental temperatures of methanogenesis methane can be generated that does not form in internal isotopic

(D18 of 5.5–7, corresponding to temperatures between 50 and 0 equilibrium. While most naturally occurring thermogenic °C), the aH2O-CH4values are also usually consistent with formation methane samples have T18 values that are consistent with plausi- in isotopic equilibrium at that temperature (Fig. 3). Based on this, ble formation temperatures, kinetic isotope effects may help to quantitative models have been created that relate the reversibility explain some cases where T18 values appear to be too high, as of methanogenesis to both the bulk isotopic compositions of discussed below (Section 4.5.1). 13 microbial methane (d C and dD) as well as D18 values (Fig. 3). Recent experiments generating methane via Sabatier reactions These models are capable of describing, to first order, the co- using ruthenium catalyst also resulted in strong kinetic isotope 12 variation between aCH4-H2O values and D18 values (Stolper et al., effects, as indicated by large negative deviations in D CH2D2 13 2015; Wang et al., 2015). However, the models applied thus far and smaller deviations in D CH3D relative to the experimental have multiple free parameters and do not supply unique solutions. temperature (Young et al., 2017). The authors of this study pro- In any case, the agreement of these models with available data posed that the mechanism responsible for kinetic isotope effects from pure cultures and natural microbial methane samples indi- observed in these experiments, and possibly in microbial methane cates that they are useful descriptions of the key processes that as well, involves quantum tunneling effects associated with hydro- control the microbial formation of methane (Fig. 3). gen isotope fractionation. These models were developed to characterize fractionations in hydrogenotrophic methanogenesis. To date, there have been two 3.3. Mixing effects published analyses of fermentative (specifically methylotrophic) methanogenesis from a pure culture (Douglas et al., 2016; As with clumped isotope values for other gases, mixing between Young et al., 2017), only one of which published the dD value end-members that differ in their conventional isotope values (i.e., 13 13 of the culture media water. The result from that study deviates dD, d C) can show non-linear variation in D18 and D CH3D values. 13 significantly from the model predictions, with low D18 relative This non-linearity results from the definition of D18 and D CH3D to aCH4-H20 (Douglas et al., 2016)(Fig. 3). Possible explanations values in reference to the stochastic distribution of mass-18 iso- for this deviation are discussed in detail by Douglas et al. topologues, which is a non-linear polynomial function of dD and 13 (2016). One plausible explanation is that fermentative methane d C values (Eq. (5); Fig. 4). The non-linearity of mixing in D18 is derives a portion of its hydrogen atoms (50–75%) from methyl negligible when end-member dD and d13C values are similar, but groups (Pine and Barker, 1956; Waldron et al., 1999), and there- becomes progressively larger as end-member dD and d13C values fore may not express aCH4-H20 values as large as those expressed become more widely spaced. Depending on differences in bulk by hydrogenotrophic methanogenesis, which derives all of its composition, the resultant D18 values can be either larger or smal- hydrogen atoms from H2 in isotopic equilibrium with water ler than the expected value for conservative mixing of D18. The cur- (Daniels et al., 1980; Valentine et al., 2004). Indeed, correcting vature induced in D18 by mixing of end-members for two- for this effect in methylotrophic methanogenesis experiments component mixing is diagnostic of the mixing process (Douglas leads to aCH4-H20 values that more closely agree with the model et al., 2016). predictions (Fig. 8 in Douglas et al., 2016). Some methane sam- The D18 of some mixed methane samples does not correspond ples obtained from freshwater and serpentinization environments to meaningful formation temperatures (Douglas et al., 2016). How- also deviate from the model prediction similarly to the methy- ever, in some cases the non-linearity of mixing can lead to D18 val- lotrophic methane, although to a lesser degree (Fig. 3). This is ues that provide a diagnostic fingerprint of mixing (Fig. 4). If consistent with the observation that fermentative methanogene- multiple samples of varying mixing ratio can be measured, the cal- sis occurs widely in freshwater environments (Ferry, 1993; culation of a D18 mixing curve can provide useful constraints on Borrel et al., 2011). the isotopic compositions, and potentially formation temperatures, 268 P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282

13 Fig. 4. Hypothetical examples of mixing and diffusion effects for D18 values, adapted from Douglas et al. (2016). Plots show mixing relationships in d C-D18 space (A) and dD-

D18 space (B) for mixtures of methane with varying end-member compositions. In the mixing examples (black lines) the end-member D18 values remain fixed, but the end- 13 13 13 member d C and dD values vary. For mixtures where d C and dD values are relatively similar, mixing in D18 is essentially linear (solid line); as the d C and dD values of the mixing end-members become increasingly widely spaced the non-linearity of mixing in D18 becomes more pronounced (dashed lines). Trajectories for gas-phase interdiffusion of methane in air are shown by the red arrows; the arrow shows the direction of isotopic fractionation of the escaping methane.

13 of the mixing end-members (Stolper et al., 2015; Douglas et al., residual methane decreasing in D CH3D as aerobic oxidation pro- 2016). Alternatively, if the mixtures have similar dDord13C values, ceeds, while d13C and dD values increase. These results are consis- 13 the inferred temperature of the mixture will reflect a pseudo- tent with a prediction that the fractionation factor for CH3D average of the end member D18 temperatures (Stolper et al., during aerobic methane oxidation is approximated by the product 2014b, 2015). of the 13C/12C and D/H fractionation factors. The effect of aerobic methane oxidation on clumped isotopes in open-system natural environments has not been studied, and could be highly dependent 3.4. Other kinetic isotope effects on the interaction of oxidation and transport processes (Wang et al., 2016). Anaerobic methane oxidation employs a biochemical In addition to equilibrium temperature effects, kinetic effects pathway that is distinct from aerobic methane oxidation, and some associated with methane generation, and mixing effects, a number evidence suggests it leads to partial carbon isotope equilibration of D of other processes could also have significant effects on 18 values. residual methane (Holler et al., 2011; Yoshinaga et al., 2014), but Diffusion of methane in either a vacuum (i.e., gases following Gra- the effect of this process on clumped isotope values has not been ham’s law of diffusion) or diffusion at significant gas pressures in studied. which interactions between particles takes place (i.e., gas-phase Similarly, reactions with OH, and to a lesser extent Cl, are D interdiffusion) is predicted to increase 18 values of the gas that major sinks for atmospheric methane (Khalil and Rasmussen, d d13 escapes relative to the residual, but to decrease D and C values 1983; Kirschke et al., 2013). There have been laboratory observa- of the escaping gas (Fig. 4). Specifically, diffusion of methane tions of kinetic isotope effects for reactions between both OH D through air is predicted to increase 18 in the escaping gas by and Cl with 13CH D and 12CH D (Gierczak et al., 1997; Feilberg ‰ d d13 3 2 2 1.5 , and to decrease D and C values in the escaping gas by et al., 2005; Joelsson et al., 2014, 2015), which are discussed below ‰ 19 . Diffusion through a liquid or solid has unknown effects in Section 5.1. on D18. Similar diffusive fractionations have been described for multiply-substituted isotopologues of CO2 (Eiler and Schauble, 2004), O2 (Yeung et al., 2012), and N2O(Magyar et al., 2016). Frac- 4. Methane clumped isotope data from environmental samples tionation during diffusion through a liquid phase, which is an important transport mechanism for methane in aquatic environ- 4.1. General patterns of isotopic variation in environmental methane ments, is unknown both theoretically and experimentally. We note that a sample of methane from gas bubbles at Killarney Lake in Clumped isotope measurements of 250 methane samples

Alaska presented an anomalously high D18 value (9.6‰) and a have been reported to date, including data from this paper, repre- low d13C value (88.76‰) relative to other Alaskan lake samples, senting diverse Earth environments. Of these, 135 measurements a result that is consistent with diffusive fractionation (Douglas are presented for the first time here. The totality of D18 and 13 et al., 2016). D CH3D data span a wide range of values from 5.4‰ (pure cul- Aerobic and anaerobic microbial methane oxidation are critical ture of a methylotrophic methanogen; Douglas et al., 2016)to sinks for methane in many environments, especially in surficial 10.1‰ (a natural gas sample from the Songliao Basin in eastern marine and terrigenous environments (Valentine and Reeburgh, China; this study). The lowest value for an environmental sample 13 2000; Le Mer and Roger, 2001; Gupta et al., 2013). Methane oxida- is 3.4‰ (D CH3D), observed in methane from the Cedars serpen- tion generally leads to enrichment of both d13C and dD in the resid- tinization zone of Central California (Wang et al., 2015). 13 ual methane (Whiticar et al., 1986; Whiticar, 1999). Recent closed- Plotting the three isotopic parameters (D18 or D CH3D, dD, 13 13 system batch-culture experiments constrain how D CH3D values d C) for all samples from natural environments (i.e., not labora- are affected by aerobic methane oxidation (Wang et al., 2016), with tory experiments) reveals a broadly triangular distribution P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 269

(Fig. 5). Samples with the highest D18 values generally also have values (Fig. 5). In one, decreasing D18 values correspond to increas- the lowest d13C values and intermediate dD values. These samples ing d13C and dD values (solid line in Fig. 5). This sample set includes represent primarily marine and ‘deep’ biosphere (i.e., organisms most of the thermogenic, volcanic and hydrothermal methane, and living in buried sedimentary strata) microbial methane, but also some samples from serpentinization systems. In general, T18 values include some samples that are inferred to be from mixed sources. for these samples indicate plausible methane formation tempera-

There are two broad but distinct trends of data with decreasing D18 tures (Fig. 6), and we infer that this axis of variability primarily

Fig. 5. Plots showing the isotopic distribution of environmental methane (i.e., not produced in laboratory experiments) samples analyzed for clumped isotopes to date. (A) 13 13 13 13 Three-dimensional plot of d C vs. dD vs. clumped isotopes (D18 or D CH3D); (B) d C vs. clumped isotopes; (C) dD vs. d C; (D) dD vs. clumped isotopes. Empirical fields for different methane sources as defined in Fig. 1 are shown in (C). General trends for equilibrium fractionation (solid line) and kinetic isotope fractionation (dashed line) are indicated in each plot. Data from Stolper et al. (2014b), Inagaki et al. (2015), Stolper et al. (2015), Wang et al. (2015), Douglas et al. (2016), Young et al. (2017) and this study. 270 P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 corresponds to equilibrium isotope fractionation. However, some of methane on the early Earth, its role in climatic conditions, and of the samples that plot along this trend are influenced by mixing the development of life (Emmanuel and Ague, 2007). Methane in effects. volcanic and hydrothermal systems is generally thought to derive

In the other trend, decreasing D18 values correspond to decreas- from (i) the high-temperature breakdown of buried and/or sub- ing dD values, as well as a relatively small increase in d13C values ducted organic matter (similar to thermogenic methane); (ii) abio- (dashed line in Fig. 5). This group of samples include some that genic synthesis of methane at moderate to high temperatures 13 yield negative or ‘anti-clumped’ D18 or D CH3D values. Samples (200 to 600 °C); or (iii) mantle-derived methane (Welhan, falling along this trend are mainly microbial methane derived from 1988; Emmanuel and Ague, 2007; Proskurowski et al., 2008). The freshwater ecosystems and serpentinization-zone microbial com- relative fraction of these sources varies and primarily depends on munities and we interpret this trend to represent kinetic isotope the amount of sedimentary rocks in the volcanic system effects related to the differential reversibility of microbial (Welhan, 1988). Carbon isotope ratios are often applied to differen- methanogenesis, as discussed in Section 3.2. tiate the source of hydrothermal methane, with high d13C values Some samples do not conform neatly to either of the two iso- (>20‰) often interpreted as indicative of abiogenic methane topic trends discussed above. For example, unconventional oil- (Welhan, 1988; Fiebig et al., 2015), although experimental and nat- associated thermogenic methane samples (See Section 4.4) tend ural data suggest revisions to this interpretive framework (Horita to have lower dD values, relative to their D18 values, compared to and Berndt, 1999; Etiope and Sherwood Lollar, 2013; Etiope and other thermogenic methane samples. Samples from deep crustal Schoell, 2014).

fluids (Wang et al., 2015; Young et al., 2017) (see Section 4.4), There are currently eight CH4 clumped isotope measurements inferred to be abiotic methane (Sherwood Lollar et al., 2008), have from hydrothermal and volcanic systems, including four samples relatively low dD values but high d13C values, making them unique from two distinct marine hydrothermal vents, and four from three in these plots relative to other sample types (Fig. 5). different terrigenous hydrothermal systems (Supplementary Table). The four samples from marine systems, one from the Guay- 4.2. Hydrothermal and volcanic methane mas Ridge in the Gulf of California (Wang et al., 2015) and three from the Juan de Fuca Ridge (Supplementary Table), yield forma- Methane emissions from hydrothermal vents and volcanoes are tion temperatures from 304 ± 40 to 365 ± 50 °C. These inferred a relatively small component of the global methane budget, con- temperatures are within error of vent fluid temperatures, which tributing about 2–9 Tg/yr (Lacroix, 1993; Etiope et al., 2008), com- are estimated at 299 ± 5 °C for the Guaymas Ridge (Reeves et al., pared to a total global flux of 540–680 Tg/yr (Kirschke et al., 2013). 2014), and 324 °C for Bastille and 335 °C for Lobo (maximum mea- However, methane from these systems can provide important sured temperatures) at the Main Endeavour Field on the Juan de insights into the carbon-cycle geochemistry of these high- Fuca Ridge in July 2014. Methane emitted from Guaymas Ridge temperature environments. Modern volcanic and hydrothermal was previously interpreted to form from thermocatalysis of buried methane sources may be an important analogue for the generation organic matter (Welhan, 1988), and the low d13C values (44 to

Fig. 6. Plot shows the distribution of formation temperatures inferred from clumped isotope measurements for different categories of environmental methane. The red lines indicate the median value, the blue box indicates the first and third quartiles, the black whiskers indicate the maxima and minima within 1.5 interquartile range of the first and third quartile, and individual points indicate outlier values beyond this limit. The number samples per category is listed. The arrow for freshwater microbial indicates that some values extend beyond 600 °C. Some samples in this category have negative clumped isotope values, which do not correspond to any temperature. Microbial methane 13 samples from serpentinization sites are not shown since these samples have negative D CH3D values. Data from Stolper et al. (2014b), Inagaki et al. (2015), Stolper et al. (2015), Wang et al. (2015), Douglas et al. (2016), Young et al. (2017) and this study. P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 271

50‰) and relatively high dD values (96 to 106‰) of all four of interpreted as indicators of microbial methanogenesis accompa- these samples are consistent with this mechanism (Fig. 1; Fig. 5). nied by strong kinetic isotope effects (Fig. 3). A previous study of The clumped isotope data suggest that, at least in the studied sys- the geochemistry of the Cedars serpentinization site suggested that tems, methane derived from the thermal cracking of organic mat- contributions from both abiotic and microbial methane were likely ter in hydrothermal systems forms at significantly higher (Morrill et al., 2013), but the clumped isotope data are consistent temperatures than those commonly observed in natural gas reser- with a dominantly microbial, strongly non-equilibrium methane 13 voirs (Fig. 6). source. Interestingly, in a plot of D CH3D vs. aH20-CH4, the Cedars The four samples from terrigenous hydrothermal systems indi- samples plot near a sample from a pure culture of fermentative cate generally higher temperatures (364 ± 49 °C at Pantelleria, methanogens grown with a methanol substrate (Fig. 3). Morrill Italy, 444 ± 79 °C at Nisyros, Greece, 347 ± 45 and 578 ± 109 °Cat et al. (2013) suggested that microbial acetogenesis occurs in these Washburn Spring, Yellowstone, USA) (Supplementary Table). It springs, potentially providing a ready source of acetate for fermen- has been suggested that CH4 and CO2 at Nisyros and Pantelleria tative methanogenesis. However, it is currently unknown whether occur in isotopic equilibrium, with carbon isotope fractionation fermentation of methanol versus acetate yields similar clumped pointing to hydrothermal reservoir temperatures of 320–360 °C isotope compositions. and 540 °C, respectively (Fiebig et al., 2004, 2013). Accounting Methane sampled from wells in the Coast Range Ophiolite for analytical error, the apparent clumped isotopic equilibration Microbial Observatory (CROMO) is significantly more enriched in temperature for Nisyros is only slightly higher than that deter- d13C(26 to 27‰) and dD(157.5 to 169.5‰) than that at mined by Fiebig et al. (2004). Additional clumped isotope data Cedars, and has higher D13CH3D values (5.2–4.4‰) consistent with for CH4 from Nisyros could be used to further evaluate the hypoth- a formation temperature between 42 ± 11 and 76 ± 13 °C(Wang esis of CH4 and CO2 occurring in equilibrium. In the case of Pantel- et al., 2015). It remains unclear if the methane sampled from leria, the apparent clumped isotopic temperature is significantly CROMO is abiotic (Wang et al., 2015), but assuming the methane lower than that based on the apparent carbon isotope fractiona- formed in isotopic equilibrium, the inferred temperatures suggest tion. This discrepancy either indicates the absence of carbon iso- formation within the ophiolite (peridotite) nappe, as generally tope equilibration between CH4 and CO2 in this system, or points considered for continental serpentinization sites (Etiope et al., to the importance of H-isotope re-equilibration during the ascent 2016). of the gases, which could re-set the apparent clumped isotope Methane emitted at the Chimaera seeps in Turkey (Etiope et al., temperature. 2011) is characterized by highly enriched d13C(12‰) and dD At Washburn Springs in Yellowstone, carbon isotope fractiona- (116‰) values that suggest a predominantly abiotic origin. tion between CH4 and CO2 indicates a formation temperature of The D18 values in these samples (2.2–2.3) suggest an equilibrium 286 °C(Moran et al., 2017), while the temperature of the formation temperature of 235 ± 29 °C(Fig. 6, Supplementary hydrothermal reservoir is estimated to be 360 °C(Truesdell et al., Table). Earlier application of an H2-CH4 dD geothermometer 1977). One of the clumped isotope analyses indicates a tempera- implied that methane at Chimaera formed at 50 °C(Etiope ture similar to that of the hydrothermal reservoir, while the other et al., 2011), which is clearly at odds with the D18-derived temper- is considerably higher. The wide range of inferred temperatures ature. The local geothermal gradient suggests maximum tempera- from the Yellowstone samples is intriguing, as it could suggest tures of 80–100 °C at the base of the ophiolite nappe, within which either differential sources of methane despite similar dD and d13C abiotic methane is considered to have formed (Etiope et al., 2011). values, or differences in the extent of re-equilibration within the It cannot be excluded, however, that methane formed near the hydrothermal reservoir. However, this difference will need to be metamorphic sole (high-grade metamorphic rocks within the validated with further analyses. Overall, the terrigenous ophiolite complex) during ophiolite obduction, which likely expe- hydrothermal data are compatible with high-temperature (> 300 rienced higher temperatures (Etiope et al., 2016). Furthermore, 13 12 °C) abiotic methane in these systems. analysis of D CH3D, both alone and in tandem with D CH2D2, indicates methane forming at or near internal isotopic equilibrium 4.3. Methane from low-temperature serpentinization sites at a temperature of 128 ± 10 °C(Wang et al., 2014; Young et al., 2017). The disagreement between the clumped isotope data and

Water–rock interactions in low-temperature terrigenous ser- the H2-CH4 dD geothermometer indicate that the H2 and CH4 emit- pentinization sites (i.e. <150 °C; (Etiope et al., 2011) are also ted at the seep are not in isotopic equilibrium. thought to be a source of abiotic methane, with an uncertain global The difference in inferred temperatures between the different flux (Etiope et al., 2016). The most widely discussed mechanisms clumped isotope measurements for the Chimaera seep samples is for methane production in serpentinization zones are Sabatier or both noteworthy and problematic. We suggest there may be three Fischer-Tropsch type reactions in which inorganic carbon (mainly possible causes for this difference. First, leakage of gas from imper-

CO2) combines with molecular hydrogen to form methane (and lar- fectly sealed sample containers could lead to a diffusive isotope ger hydrocarbons) and water (Emmanuel and Ague, 2007; effect and a depletion in D18 in the residual gas, leading to artifi- McCollom, 2013; Etiope and Schoell, 2014). In many cases, how- cially high T18 values. Indeed, earlier analyses at Caltech of gas ever, it is difficult to rule out contributions of microbial or migrated sampled from Chimaera appeared to demonstrate such an effect, 13 thermogenic methane in serpentinization zones (Etiope et al., with progressively lower D18 and higher dD and d C values in 2011, 2013). sample containers with lower methane concentrations. The sam- Clumped-isotope compositions of methane from five different ples presented here, however, were analyzed later and were stored low-temperature serpentinization systems have been analyzed, in better-sealed sample containers. An argument against leakage is with widely varying results. These include two sites from central the similarity of dD and d13C values between samples analyzed at California (Cedars and CROMO), and sites from Turkey (Chimaera), Caltech and at UCLA (Supplementary Table; Young et al., 2017). Portugal (Cabeço de Vide), and Italy (Acquasanta). In the Cedars Second, the difference could reflect unresolved discrepancies in serpentinite site from central California, Wang et al. (2015) clumped isotope reference frames and standardization between 13 observed strongly negative D CH3D values (2.4 to 3.4‰) that laboratories and methodologies. Third, it is possible, but unlikely, clearly indicate non-equilibrium isotope fractionation during that the difference reflects true isotopic heterogeneity in Chimaera methane generation or migration. These values, alongside low gases. Resolving this difference will require additional interlabora- d13C(63.8 to 68.0‰) and dD(333.1 to 342.0‰) values, were tory calibration and standardization. 272 P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282

13 12 Recent coupled analyses of D CH3D and D CH2D2 indicate Stolper et al. (in press) for a more detailed discussion of the appli- that while the Chimaera methane appears to form in isotopic equi- cation of clumped isotopes in natural gas. In general, thermogenic librium, methane at other low-temperature serpentinization sites methane D18 values indicate formation temperatures from 100 to (Cabeço de Vide, Portugal and Acquasanta, Italy) is not in isotopic 300 °C(Fig. 6), which broadly corresponds to estimates of the tem- 13 equilibrium, and that therefore D CH3D-derived temperatures perature window for catagenesis of buried organic matter (Quigley from these sites likely do not reflect formation temperatures and Mackenzie, 1988; Seewald, 2003). A few samples from marine (Young et al., 2017). The mechanisms for disequilibrium at these seeps in the Santa Barbara basin, which are believed to be thermo- sites are not clearly identified, but could reflect kinetic effects genic in origin, indicate formation temperatures around 90 °C induced by surface catalysis, diffusion, and/or mixing effects. (Supplementary Table), but we cannot rule out a contribution of microbial methane in these samples. There are a number of ther- 4.4. Deep fracture fluid methane mogenic methane samples from unconventional reservoirs with

T18 estimates greater than 300 °C. These temperatures are higher Methane samples from crustal fluids in Precambrian shield for- than expected for catagenetic methane formation in sedimentary mations from Canada and South Africa were analyzed by Wang basins, and we discuss possible explanations for this discrepancy et al. (2015) and Young et al. (2017). Methane in these samples below (Section 4.5.1). 13 was characterized by relatively high d C values between 32 In general, T18 values for methane from oil-associated, conven- and 42‰, and low dD values between 323 and 421‰ com- tional reservoirs (mean 171 °C, minimum 87 °C) are lower than pared to other categories of environmental methane (Fig. 5). those for other thermogenic categories (mean between 205 and 13 D CH3D values vary between 3.1 to 6‰, corresponding to temper- 230 °C, and minimum 140 °C) (Fig. 6). This difference in temper- atures from 25 ± 7 to 150 ± 20 °C. Based on previous observations ature distributions could be interpreted to mean that methane of methane and higher isotope values, these samples found in oil-associated conventional reservoirs is often produced are believed to be abiotic (Sherwood Lollar et al., 2008). The via lower temperature cracking of larger organic molecules associ- inferred formation temperatures from Kidd Creek in Canada are ated with the ‘oil-window’ (Quigley and Mackenzie, 1988; generally higher than measured groundwater temperatures (30 Seewald, 2003), whereas other, higher temperature processes °C) (Sherwood Lollar et al., 2008), but could represent methane (e.g., secondary cracking of larger hydrocarbons produced from that formed deeper in the crust and subsequently migrated to its kerogen and later stages of kerogen cracking) dominate methane current thermal environment. However, there is evidence that production in the other reservoir categories. However, the upper most of these methane samples formed out of isotopic equilibrium, range of T18 values (maximum of 298 °C; Fig. 6) in conventional, 13 based both on comparison of D CH3D values and aCH4-H2O values oil-associated reservoirs suggests that methane from higher- 12 (Wang et al., 2015), as well as comparison of D CH2D2 and temperature cracking of shorter-chain hydrocarbons (relative to 13 D CH3D values (Young et al., 2017). The proposed mechanisms co-generation of methane with oil) can also be significant in some for disequilibrium in these systems are kinetic isotope effects gen- reservoirs. However, as discussed below (Section 4.5.1), tempera- erated during catalyst-mediated Fischer-Tropsch reactions (Young tures at the upper range of the T18 distribution for all of the cate- et al., 2017). gories of thermogenic methane may not reflect accurate formation temperatures, and instead could be influenced by kinetic 4.5. Thermogenic methane isotope effects during formation, transport, or extraction. We observe a weak but significant positive correlation (R2 = 0.2; 13 Thermogenic methane formed by high-temperature breakdown p < 0.001) between d C values and T18 estimates for methane in all 13 of organic matter in deeply buried sediments (>1 km), is a major natural gas reservoirs (Fig. 7A). The trend between d C and T18 component of economically recoverable petroleum systems. It is varies between basins, likely due to differences in the d13C value also a significant contributor to atmospheric methane, both of the source kerogen, among other factors (see Stolper et al., in through natural emissions from seeps and mud volcanoes (50–60 press for further details). We also observe a weak, but significant, 13 Tg per year) (Etiope et al., 2008; Etiope, 2012; Schwietzke et al., positive correlation between d C and T18 in conventional reser- 2016), and through anthropogenic emissions related to fossil fuel voirs (R2 = 0.29; p < 0.001). These relationships are consistent with extraction, distribution, and use (98–150 Tg per year) (Denman the concept that the d13C of methane increases with source rock et al., 2007; Schwietzke et al., 2016). We discuss several subcate- maturity (Schoell, 1980). However, such a relationship is not gories of thermogenic methane, primarily defined by the type of observed in unconventional reservoirs, and in fact there is a nega- 13 reservoir in which the gas is stored. Conventional gas reservoirs tive relationship between d C and T18 in non-associated, uncon- generally occur in porous sedimentary rocks capped by an imper- ventional methane samples (R2 = 0.23, p < 0.01). One possible meable stratum (Hunt, 1979). Unconventional gas accumulations explanation for this negative relationship is that it reflects a non- are generated and reservoired within the source rock, which are linear mixing trend between methane formed at different maturi- typically organic-rich mudrocks (Curtis, 2002; Peters et al., 2015). ties, as described by Stolper et al. (in press). We do not observe a

Unconventional natural gas is typically extracted by hydraulic frac- correlation between dD and T18 values in any of the thermogenic turing of mudrock strata, which releases gas, and in some cases oil. methane categories (Fig. 7B). Notably, unconventional, associated

We further subdivide natural gas samples into those that are asso- methane samples generally have lower dD values for a given T18 ciated with significant deposits of liquid petroleum (‘oil- than methane from other types of reservoirs. associated’) and those that are not (‘non-associated’). Associated We also compared T18 values with ratios of methane to heavier gases are either dissolved in oil (‘solution gas’) or present as a gas- hydrocarbons, i.e., ‘gas wetness’, calculated here as [C1]/[C2+C3] eous phase overlying a liquid phase in the reservoir (a ‘gas cap’). (Fig. 8). Gas wetness is often used as a qualitative indicator of ‘Non-associated’ gases are present in the gaseous phase in the the maturity of hydrocarbon systems, with higher temperature reservoir and are not in contact with any liquid hydrocarbons in catagenesis and the associated cracking of smaller hydrocarbons the subsurface. generally producing ‘drier’, more methane-rich gases (Fig.1B), Thermogenic natural gas samples comprise by far the largest set although initial gases formed at low temperatures are also often of clumped isotope measurements to date, making up approxi- relatively ‘dry’ (Bernard et al., 1978; Hunt, 1979). In our dataset, mately 45% of the total dataset. We provide a general overview oil-associated gases (excluding samples from marine hydrocarbon of patterns observed in these data here, but refer the reader to seeps in the Santa Barbara Basin) generally have relatively low P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 273

13 Fig. 7. Plot of (A) d C vs. T18 and (B) dD vs. T18 values for samples from natural gas reservoirs or seeps. The regression line is fit to the entire dataset depicted. Here we only plot samples of an inferred mixed source from natural gas wells. Data from Stolper et al. (2014b, 2015), Wang et al. (2015), Young et al. (2017) and this study.

[C1]/[C2+C3] values ranging from 1 to 13. This corresponds to a wide range of methane concentrations, from 50% to 93%. There

is no clear trend in these samples between gas wetness and T18, although interestingly some of the ‘wettest’ conventional samples

have relatively high T18 temperatures (>160 °C) that are above the typical oil-window (Hunt, 1979). Gas from hydrocarbon seeps in the Santa Barbara Basin are

characterized by higher [C1]/[C2+C3] values for a given T18 value than other conventional, oil-associated samples. They also follow a trend similar to that of mixed microbial and thermogenic gases,

although with relatively high [C1]/[C2+C3] values. While these gases are generally thought to originate from underlying conven- tional hydrocarbon reservoirs (Hornafius et al., 1999), this pattern is consistent with molecular fractionation (i.e., an increase in the C1/C2+ ratio during migration) and/or secondary methanogenesis, which are both typical of many seeps (Etiope et al., 2009). Non-

associated gases span a much wider range of [C1]/[C2+C3] values 3 (as high as 10 ), but gas wetness is not clearly related to T18 values. Dry gases with [C1]/[C2+C3] greater than 20 likely represent methane dominantly generated beyond the oil window (>160 °C),

which is generally consistent with their T18 values. Methane samples from terrigenous Type-III kerogen in the Rotliegend formation (conventional, non-associated), span a range

of T18 values from 193 to 267 °C (with the exception of one sample with a lower T18 of 144 °C), and a narrow range of [C1]/[C2+C3] val- ues, from 52 to 68. Excluding the outlier, these samples indicate a consistently high formation temperature of 225 ± 26 °C(1r), which is equivalent to the analytical error at this temperature range. Terrigenous hydrocarbon sources that are older than the Cenozoic, including coal, often generate dominantly dry gas regardless of the thermal maturity of the rock (Schoell, 1983).

The T18 values in this case indicate dry gas formation from coal Fig. 8. Plot of [C ]/[C +C ] vs. T values for samples from natural gas reservoirs or 1 2 3 18 source rocks at high temperatures (200 °C), although as discussed seeps. Not all natural gas reservoir samples had gas composition data available. We only plot samples of an inferred mixed source from natural gas wells. Data from below (Section 4.5.1) there are several possible factors that could Stolper et al. (2014b, 2015), Wang et al. (2015) and this study. lead to T18 values that could be higher than the actual formation 274 P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 temperature. In the case of the Rotliegend Formation the effect of components stored at high pressures be heated during extraction gas diffusion from the reservoir or source rock could be important. to avoid this effect. Second, high clumped isotope temperatures may be the result of secondary isotope effects related to methane transport, either

4.5.1. Anomalously high T18 values in thermogenic methane occurring naturally or as a result of oil and gas extraction. Gas dif- D In some samples from natural gas reservoirs, T18 values indicate fusion, in particular, is predicted to decrease 18 values in the formation temperatures that are unrealistically high (Figs. 7 and 8). residual gas. For example, diffusive loss of 30% of the methane in In some cases these temperatures exceed the nominal upper bound a reservoir, assuming inter-gas diffusion in air, would cause the of thermogenic gas generation (300 °C) (Quigley and Mackenzie, residual methane D18 to increase by 0.6‰. If methane originally 1988; Seewald, 2003), e.g., in samples from the Eagleford and Bak- formed at 230 °C(D18 of 2.3‰), this diffusive loss would lead to ken shales and the Appalachian Basin. an increase in T18 values to 305 °C(D18 of 1.7‰) (Section 3.4). We suggest four plausible explanations for this phenomenon. The high apparent temperatures measured for some methane sam- First, high apparent clumped-isotope temperatures could be an ples could thus indicate that the analyzed sample is residual gas analytical artifact, either related to sample preparation or analysis. remaining in the source or reservoir rock after diffusive loss of a

In two samples from the Eagleford Shale that indicated high T18 significant fraction of the total gas in the reservoir. We note, how- values when prepared using the standard technique, we subse- ever, that inter-gas diffusion is not a realistic model for petroleum quently re-extracted gas samples while heating the sample cylin- systems, and that isotope effects for diffusion through permeable der to 85 °C(Fig. 9). This was done because these samples rock or liquid are poorly constrained (Prinzhofer and Huc, 1995; contain non-trivial quantities of C4+hydrocarbons, which would Zhang and Krooss, 2001), and unknown for clumped isotopes. It be in the liquid phase at the pressures of the sampled cylinders. is unknown whether either hydraulic fracturing or subsequent

The heated extractions yield lower T18 values, with a decrease in gas separation induce measurable isotopic fractionations, and it T18 from 120 to 220 °C(Fig. 9). While the mechanism for this shift is possible that these production processes could also contribute is unknown, we suspect that liquid hydrocarbons in the sample to high T18 values. cylinders may retain some methane, and that this absorption or Third, high temperature cracking experiments have in some dissolution causes an isotopic fractionation leading to apparently cases produced methane with T18 values that are significantly high temperatures. In contrast, a sample of mixed thermogenic higher than the experimental temperature (Shuai et al., in revi- and microbial gas from the Gulf of Mexico, which indicated a sig- sion), as discussed above in Section 3.2. These data may indicate nificantly lower T18 value (70 ± 9 °C) when originally extracted at that thermogenic cracking of sedimentary organic matter can pro- D room temperature, did not show a significantly different T18 value duce methane with non-equilibrium 18 values under specific cir- (58 ± 8 °C) when extracted at 85 °C(Fig. 9). The gas in this cylinder cumstances, and that high apparent temperatures in some systems was at a lower pressure, and therefore C4+hydrocarbons were less could be a result of this non-equilibrium isotope effect. likely to be liquids and to adsorb methane. Most samples discussed Finally, high T18 values could represent true formation temper- here were not stored in high-pressure cylinders containing signif- atures for methane originating at greater depths than is currently icant quantities of C4+ hydrocarbons, and this is unlikely to be an predicted by models of oil and gas generation (Seewald, 2003). issue for our interpretation of D18 data, except in the cases of the The variability in T18 values observed in many of the studied sed- Eagleford and Bakken shale samples. Regardless, we recommend imentary basins might then reflect differential contributions of that all samples from natural gas reservoirs with substantial C4+ methane from shallow versus deep environments.

Fig. 9. Plot shows the difference in T18 values for three natural gas samples between methane extracted without heating the steel cylinder and methane extracted while heating the cylinder to 85 °C. In a sample from the Gulf of Mexico with a relatively low T18 value the difference between treatments is within error. For two samples from the

Eagleford Shale, however, heating the sample cylinder led to substantially lower T18 values. P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 275

The heated cylinder extraction and non-equilibrium results methane from distinct formation environments. This mixing was from pyrolysis experiments raise the possibility that analytical first identified by linear co-variation of dD, d13C, and D14C values, artifacts or kinetic isotope effects during methane generation and was originally thought to reflect mixing of microbial methane may be responsible for some of the highest apparent temperatures produced in lake sediments, and thermogenic methane from that have been observed. Furthermore, isotope effects during gas underlying strata (Walter Anthony et al., 2012). Clumped isotope diffusion or separation may also possibly alter D18 signals, analysis revealed elevated D18 values, which correspond to nega- although the signature of these processes has not been clearly tive apparent temperatures (i.e., below 0 °C) in two of these sam- identified in natural samples. It remains to be seen whether evi- ples. The data also exhibit a non-linear trend in D18-dD space 14 13 dence supports methane forming at temperatures greater than (Fig. 10). When constrained by C data, the D18-dD and D18-d C 300 °C and whether this represents a substantial contribution to two end-member mixing curves fit to these samples suggest a high natural gas reservoirs. temperature component that formed at around 60 °C(Douglas et al., 2016). Methane emitted in nearby Prince William Sound

4.6. Mixing of methane from distinct sources has a similar T18 value of 73 ± 10 °C, consistent with this formation temperature. These relatively low temperatures, combined with

Mixtures of methane from different sources are common in nat- high [C1]/[C2] values, suggest that the high-temperature end- ural gas reservoirs and seeps (Schoell, 1983). Clumped isotope data member in this setting may originate from deep subsurface can provide new insights into such mixtures and the characteris- microbes instead of thermogenic gas generation (Douglas et al., tics of the end-members, and serve as a basis for quantitative 2016). These data comprise the first instance in which non-linear apportionment. As discussed above, mixing can be non-linear in mixing of D18 values in a natural environment results in noticeably 13 D D18 or D CH3D values (Section 2.3, Fig. 4), with the non- elevated 18 values (Fig. 10), but such an effect is likely to occur in linearity becoming increasingly pronounced as the difference in other settings where methane end members produced in deep and dD and d13C values between the end-members becomes larger. surficial environments mix, especially when the dD values of the An example of clumped isotope analyses to constrain methane end-members are substantially different. mixing is in the Antrim Shale of Michigan (Stolper et al., 2015). Pre- Methane in offshore natural gas reservoirs in the Diana-Hoover viously, hydrogen isotope ratios were interpreted as indicating that Field of the Gulf of Mexico yields T18 values from 52 to 118 °C the methane in this shale was dominantly (>80%) microbial, (Fig. 10). 118 °C is a plausible thermogenic formation temperature though thermogenic gas is also clearly present in the basin based in this environment and is consistent with maturity estimates from on significant quantities (20% in some samples) of C2+ hydrocar- biomarker ratios, but the observed lower temperatures are likely bons (Martini et al., 1996). The clumped isotope data, however, the result of mixing with microbial methane produced by oil indicate two end-members of methane, one forming around 140 biodegradation. Methane from the nearby Hadrian South Field, °C or above, which is above the known range of microbial interpreted as resulting from oil biodegradation, yields T18 values methanogenesis, while the other formed at temperatures around between 34 and 48 °C, within analytical uncertainty of measured 20 °C(Stolper et al., 2015)(Fig. 10). Subsequent analysis of noble reservoir temperatures (Stolper et al., 2014b), and lower than the gas concentrations further support the contribution of significant values observed in the Diana-Hoover Field. In the Diana-Hoover 13 quantities of thermogenic methane to the Antrim Shale gas reser- Field T18 inversely correlates with both dD and d C(Figs. 7 and voir (Wen et al., 2015). 10), suggesting that the microbial methane component is relatively In southeastern Alaska a selection of methane samples emitted enriched in D and 13C relative to the thermogenic component. from seeps near Lake Eyak also exhibited signs of mixing of Previous research has shown that thermophilic methanogens

13 Fig. 10. Plots of (A) dD vs. d C and (B) dD vs. T18 for three suites of gases inferred to be mixtures of methane from high- and low-temperature formation environments. Mixing models for the Antrim Shale (Stolper et al., 2015) and Southeast Alaska (Douglas et al., 2016) were described previously. The mixing trend for the Diana/Hoover Fields was based on end-members from inferred pure microbial (Stolper et al., 2014b) and thermogenic methane (this study) from nearby fields in the Gulf of Mexico. The non- linearity of mixing of T18 values for the Antrim Shale and Diana/Hoover Fields is subtle, but is much more pronounced in Southeast Alaska, where the end-members differ much more in dD and d13C. 276 P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282

(environments between 50 and 75 °C) produce methane with rel- pure-culture incubations demonstrate that methanogens can grow atively high dD and d13C values (Schoell, 1980; Valentine et al., at higher temperatures up to 122 °C, at least under laboratory 2004) compared to mesophilic organisms (environments < 50 °C), conditions (Takai et al., 2008). Clumped isotope data exist for which could account for this effect. The variable isotopic composi- microbial methane samples from several deep subsurface tions in the Diana-Hoover Field likely represent differential mixing environments, including biodegraded oil reservoirs, terrigenous of thermogenic and microbial methane having similar d13C values. and marine coal seams, and organic-rich lacustrine shales, and

Alternatively, this variability be related to the residence time of span a range of T18 values from 34–95 °C(Fig. 11a). thermogenic gas and oil in relatively low-temperature reservoirs. T18 values from deep subsurface methanogens are generally Longer residence times may correspond to a greater proportion within error of, or higher than, estimated reservoir temperatures of microbial methane produced by oil degradation and conse- for the sampled gases (Fig. 11a). In cases where T18 values exceed quently lower T18 values. the reservoir temperatures it is possible that methane formed at Finally, in two samples from Songliao Basin in Eastern China we deeper burial depths, and then moved to shallower, cooler environ- have also observed very high D18 values of 6.9‰ and 10.2‰, corre- ments either through gas migration or uplifting of the strata. For sponding to T18 temperatures of 5 and 77 °C(Supplementary example, methane from the Milk River Formation of Alberta yields Table), with the latter D18 value the highest observed in any sam- T18 values for two samples of 69 ± 10 and 74 ± 10 °C, but the cur- ple thus far. These inferred temperatures are much lower than the rent well temperature is estimated to be 12 °C(Fig. 11b). Compar- measured well temperatures (25 and 32 °C, respectively), but are ison with a basin thermal model suggests that the T18 values are plausible as a result of non-linear variation in D18 for mixtures of most consistent with methane formation during maximum burial methane end-members having widely differing dD and d13C values temperatures of the Milk River Formation from 40 to 50 million (Figs. 4 and 10). Methane in these samples was previously inferred years ago, between 60 and 87 °C(Fig.11b). A plausible explanation to be dominantly microbial with a minor thermogenic component is that methane that formed early in the burial history at cooler based on carbon isotope and gas composition data (Zhang et al., temperatures migrated out of the strata prior to lithification, and

2011), a conclusion that is consistent with the observed high D18 that only late-formed methane from the deepest burial depths values. However, without additional samples we are unable to pro- was retained within the strata following lithification. Later, as vide further constraints on the end-member compositions and the formation was uplifted, no new microbial methanogenesis their fractional contribution to these samples. occurred, possibly because lithification made conditions unfavor- able for methanogens.

4.7. Deep subsurface microbial methane In other settings where methane T18 values are higher than reservoir temperatures, such as the Qaidam basin, a similar process Microbial methanogenesis occurs in a number of deep (>100 m can be invoked. It is also possible that there has been migration of below the land surface or seafloor) subsurface environments, gas from deeper methane-producing strata to overlying strata. Gas including buried organic-rich sediments, coal beds, and oil reser- migration may be a more likely scenario in the Qaidam Basin, voirs (Wilhelms et al., 2001; Stra˛poc´ et al., 2011; Valentine, where different wells have widely varying temperatures (from 19 2011). It has been proposed that methanogens survive tempera- to 73 °C), but in many cases the T18 values are higher than the tures up to 80–90 °C(Wilhelms et al., 2001). However, reservoir temperatures. In the uplift or gas migration scenarios

Fig. 11. (A) Comparison of T18 values vs. estimated reservoir temperatures for deep subsurface microbial methane samples. T18 values are either within error of, or significantly higher than, reservoir temperatures. T18 values higher than reservoir temperatures may indicate uplift of methane forming strata to cooler thermal environments, or migration of gas from deeper strata. Methodologies for reservoir temperature measurements or estimates are detailed in Section 2.4. We applied a conservative 20 °C error to reservoir temperatures, with the exception of drill-stem tests from the Qaidam Basin, where we applied a 10% error. (B) Comparison of the thermal history of the Milk River Formation with methane T18 values. The T18 values suggest that methane predominantly formed during the deepest burial of the sediments, prior to subsequent uplift. Data from Stolper et al. (2014b), Inagaki et al. (2015), Wang et al. (2015), Young et al. (2017); and this study. P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 277 there may be significant mixing of methane formed at different petition for substrates by sulfate-reducing microbes, marine temperatures. Additionally, in some cases there may be a compo- methanogenesis is dominantly hydrogenotrophic and strongly nent of thermogenic gases in these samples, which could raise substrate (H2) limited (Valentine, 2011). Such conditions may the inferred temperatures. Furthermore, we cannot rule out that enhance the reversibility of methanogenesis by reducing the deviations from reservoir temperatures could be caused by kinetic chemical potential gradient between methane and its precursors. isotope fractionations during microbial methanogenesis, which This leads more readily to reversibility in the enzymes generating have been clearly observed in freshwater and cultured microbial methane by methanogens (Valentine et al., 2004; Conrad, 2005), methanogenesis. One sample of putative microbial methane from thus allowing methane to achieve internal isotopic equilibrium the Niobrara Formation in Colorado demonstrated slight disequi- via rapid hydrogen exchange with water (Stolper et al., 2015; 13 12 librium in D CH3D-D CH2D2 space, which potentially explains Wang et al., 2015). a higher than expected T13CH3D value for this sample (142 °C) Another possible explanation is that activation of C-H bonds (Young et al., 2017). Finally, as noted above (Section 2.4) well tem- during anaerobic oxidation proceeds reversibly (Hallam et al., perature estimates are approximate, and errors in these values 2003) such that C-H bonds are broken and reformed faster than could explain some discrepancies with the T18 values, although the net rate of methane consumption and methane is re- they are unlikely to account for the largest observed differences equilibrated to the temperatures of anaerobic methane oxidation (i.e., > 50 °C in the Milk River Formation and Qaidam Basin; Fig. 11). (Stolper et al., 2015). This hypothesis is partially supported by One further example of deep subsurface microbial methane is a the observation that anaerobic methane oxidation promotes large gas seep (CH4 flux of 65 L/min) in a permafrost hosted lake on exchange of carbon isotopes between CH4 and CO2 (Yoshinaga the North Slope of Alaska. The methane emitted from this seep has et al., 2014). If this is the case, D18 temperatures in marine aT18 of 9 ± 10 °C, within error of the temperature of lake sediments methane may record the temperature of oxidation rather than (0 °C). However, the D18 value is higher than that observed in formation. most other lacustrine methane, as almost all other samples of this type, including from the same region of Alaska, record substantial 4.9. Freshwater microbial methane kinetic fractionations (Wang et al., 2015; Douglas et al., 2016). Additional evidence, including the relatively high dD and d13C val- Methane from freshwater environments exhibits a wide range 14 ues for this sample, the absence of C, the high methane flux, and of D18 values from 9.6 to 1‰ (Fig. 3). The high end of this range the fact that coal seams underlie the lake, all support the hypoth- corresponds to a T18 of 50 °C, while negative D18 values do not esis that the seep emits microbial coal-bed methane formed within correspond to any temperature. Most methane samples from fresh- or beneath the permafrost, which extends to about 300–400 m water environments have low D18 values that correspond to T18 below the surface in this area (Walter Anthony et al., 2012). temperatures much hotter than expected based on environmental temperatures (Fig. 6). The exceptions to this pattern include exam- 4.8. Marine microbial methane ples discussed above, e.g. methane hypothesized to have been pro- duced in sub-permafrost coal seams, mixtures of thermogenic and Microbial methanogenesis occurs in shallow (<100 m), unlithi- microbial methane emitted from a lake in southeast Alaska, and

fied marine sediments across the global ocean (Valentine, 2011). methane with an elevated D18 value (9.6‰) from a lake near Fair- Thus far, clumped isotopes have been measured in microbial banks, Alaska, whose origin is uncertain but could represent the methane from a number of seeps, pore fluids, and gas hydrates effects of diffusion (Douglas et al., 2016).

(Stolper et al., 2015; Wang et al., 2015; Douglas et al., 2016). T18 These exceptions notwithstanding, the current interpretation of values for these samples range from 0 to 42 °C(Fig. 6), almost all low D18 values in freshwater microbial methane is that they reflect of which are within error of seafloor or sediment temperatures. kinetic isotope fractionations expressed during methane genera- 13 In addition, comparison of aH20-CH4 and D18 or D CH3D values tion related to the differential reversibility of methanogenesis indicates that all marine microbial methane samples analyzed thus (Stolper et al., 2015; Wang et al., 2015; Douglas et al., 2016), as far formed at or near isotopic equilibrium, both internally and in described in Section 2.2. If this hypothesis is correct, then the vari- terms of hydrogen exchange between water and methane ation in D18 values between samples reflects differences in enzy- (Fig. 3). The depths at which these methane samples were gener- matic reversibility of methanogens in a given environment. It is ated is unknown, but the clumped-isotope temperatures provide possible that some of the observed variability is caused by mixing bounds on these depths if the local geothermal gradient is known. of methane produced with different kinetic isotope effects, or by In the case of methane sampled from the Beaufort Shelf, where post-formation processes, including oxidation and diffusion permafrost extends to a depth of 700 m (Paull et al., 2011), (Douglas et al., 2016; Wang et al., 2016). Different pathways of methane forming in relatively deep settings could be consistent methanogenesis may also influence clumped isotope values. For with inferred formation temperatures that are within error of sea- example, it has been hypothesized that the clumped isotope value floor temperatures (0 °C). In this setting the 14C content of of methane from fermentative pathways could be partly influenced methane is much lower than that of sedimentary organic matter, by isotopic clumping in methyl substrates (Wang et al., 2015; suggesting a deeper, older organic carbon source for Douglas et al., 2016). methanogenesis. Given that methanogens in freshwater environments can Gas hydrates sampled from Hydrate Ridge in the North Cascadia express significant kinetic isotope fractionations during methane

Margin record T18 values from 12 to 42 °C, whereas sediment tem- generation, clumped isotope measurements in freshwater micro- peratures range from 3 to 17 °C(Wang et al., 2015). It is possible bial methane cannot be used to infer formation temperature. How- that the highest temperature sample represents methane migrated ever, the wide range of observed D18 values suggests that such data from higher temperature environments, but the dD and d13C values could provide new insights into the biogeochemistry of freshwater of these samples suggest a dominantly microbial source. methanogenic environments, and in particular insights into the A key question in light of these data is why marine microbial cell-specific rates of methanogenesis and chemical conditions methane samples (in addition to deep subsurface microbial (e.g., thermodynamic gradients) driving microbial methane) seem to form close to equilibrium, whereas freshwater methanogenesis. and pure culture microbial methane samples clearly record kinetic In most freshwater microbial methane samples, D18 and 13 fractionations. One possible explanation is that, in the face of com- D CH3D are negatively correlated with aH20-CH4 (Fig. 3). However, 278 P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 in some cases freshwater samples deviate from the trend predicted sources will need to be modeled and constrained. However, it is by the reversibility model, with lower D18 values for a given aH20- possible that these non-linear mixing relationships could provide CH4 value. These deviations may be related to the importance of fer- distinctive fingerprints for different categories of mixtures, and mentative methanogenesis pathways in some freshwater environ- help to resolve ambiguities in dD and d13C values that can occur ments, since a fermentative methanogen culture experiment in atmospheric mixtures. expressed a larger deviation in the same direction as that shown Clumped isotopes in methane from ice cores present a second by these freshwater samples (Fig. 3)(Douglas et al., 2016). In com- exciting possibility to distinguish past changes in methane paring freshwater biogenic samples from different regions we sources and sinks. The development of methane 14C measure- 13 observe somewhat different trends in D18 (or D CH3D) versus ments in ice samples from the Greenland and Antarctic ice sheets aCH4-H2O space. For example, samples from Alaska have lower D18 (Petrenko et al., 2008, 2009; 2016) is promising in this regard, for a given aCH4-H2O relative to samples from lakes and wetlands since clumped isotope analyses require similar amounts of sam- in Sweden, Massachusetts, and California. It is possible that these ple to 14C measurements. However, as discussed above in regards trends are an artifact of how water dD values were estimated to atmospheric methane, this application will also require new (Douglas et al., 2016), but it is also possible that they represent dif- techniques to purify methane mixed with large concentrations ferences in either the dominant pathways of methanogenesis or of N2 and O2. the extent of post-formation oxidation or diffusion in these ecosystems. 5.2. Methane on other planets

Methane is also highly relevant to the study of planetary chem- 5. Potential future applications istry beyond Earth. Methane can be a significant component of planetary atmospheres (Atreya et al., 2003; Formisano et al., 5.1. Atmospheric methane 2004; Swain et al., 2008; Mumma et al., 2009). For example, methane exists in liquid form on Titan (Lunine and Atreya, 2008). Atmospheric methane is a compelling target for future methane 13 13 CH3D has been detected spectroscopically in Titan’s atmosphere, clumped isotope analyses. d C values of atmospheric methane, 13 although no D CH3D value was calculated (Bézard et al., 2007). and to a lesser extent dD and D14C values, have been valuable tools Methane is an attractive target in the search for life on other to apportion sources of atmospheric methane (Quay et al., 1999; planets because it can both be produced and consumed by Fisher et al., 2011; Townsend-Small et al., 2012), and for detecting microbes. The detection of , and variations in temporal variability in the strength of those sources (Bousquet its atmospheric concentration, has generated great interest as a et al., 2006; Sapart et al., 2012; Nisbet et al., 2014). However, ambi- potential biosignature (Formisano et al., 2004; Oze and Sharma, guities remain in apportioning methane sources based on conven- 2005; Mumma et al., 2009; Webster et al., 2013, 2015). Clumped tional isotope parameters, and additional isotopic data based on isotope measurements of Martian methane could be useful to eval- multiply substituted isotopologues could provide further con- uate its origin, since this method can be helpful in differentiating straints. However, there are analytical challenges facing the devel- thermogenic, volcanic, abiotic and microbial methane (Figs. 5 and opment of clumped isotope measurements of atmospheric 6). In addition, clumped isotope values do not depend on the dD methane. Currently all of the available measurement techniques or d13C values, which may be difficult to interpret on another plan- require relatively large (between 500 and 50 mmol) and highly etary body. With current measurement technology, however, it is purified aliquots of methane (Ono et al., 2014; Stolper et al., difficult to imagine either making a measurement using a deployed 2014a). Analysis of atmospheric methane will require removing instrument on Mars, or returning a sufficiently large sample for a N2 and O2 from very large atmospheric samples, or measuring laboratory measurement. For example, measuring D18 to a preci- methane clumped isotopes in the presence of abundant N and 2 sion of 1‰ in relatively high concentration pulses of methane O , neither of which has been demonstrated yet. 2 detected at Gale Crater (7.2 ppb; (Webster et al., 2015) would Despite the lack of atmospheric measurements, there are some require approximately 3.5 million liters of Martian atmosphere. theoretical and laboratory constraints on the clumped isotope However as technologies continue to develop, particularly in terms composition of atmospheric methane (Gierczak et al., 1997; of spectroscopic measurements, it is possible that this situation Feilberg et al., 2005; Joelsson et al., 2014., 2015; Whitehill et al., could change within the next 10–20 years. 2017). This work suggests that atmospheric chemical reactions 12 13 cause large enrichments in both CH2D2 and CH3D relative abundance. The dominant atmospheric methane sink is reaction 6. Conclusions and outlook with OH (85%, (Kirschke et al., 2013), and laboratory studies 13 indicate a for CH3D reacting with OH of A substantial body of clumped isotope data for methane now 1.343 at tropospheric temperatures (Whitehill et al., 2017), while exists, and provides a basis for understanding the major biogeo- 12 that for CH2D2 reacting with OH is estimated to be 1.81 chemical controls on this property in nature (Fig. 5). Two distinct (Gierczak et al., 1997). Similarly, the observed kinetic effect of Cl processes are recognized, broadly characterized as equilibrium 13 reacting with CH3D (1.579) (Whitehill et al., 2017) is much smal- versus kinetic fractionations. Methane appears to form in isotopic 12 ler than that for Cl reacting with CH2D2 (2.45) (Feilberg et al., equilibrium in a variety of environments, including hydrothermal, 12 2005). The substantially larger kinetic isotope effects for CH2D2 thermogenic, abiogenic, and marine microbial methane sources. In suggests that measurements of this isotopologue in atmospheric these samples, the clumped isotope composition generally pro- methane could be useful to constrain the extent of atmospheric vides insights into sample formation temperature. As discussed sink reactions, and could be used to correct atmospheric methane above (Section 4.5.1), however, there are several examples of 13 13 D CH3D, dD, and d C values for fractionations related to atmo- clumped isotope data that yield formation temperatures higher spheric sinks. This correction could be used to more accurately than expected given other independent constraints, and clumped identify the isotopic signatures of atmospheric methane sources. isotope data should always be interpreted in the context of the In any analysis of atmospheric methane, non-linear mixing of geological and geothermal conditions of the studied system. In clumped isotope values will be an important consideration, since addition, further efforts are needed to constrain the D18-T relation- the products of complex mixtures involving many methane ship below 150 °C. P.M.J. Douglas et al. / Organic Geochemistry 113 (2017) 262–282 279

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