Earth and Planetary Science Letters

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1 Earth and Planetary Science Letters 287 (2009) Contents lists available at ScienceDirect Earth and Planetary Science Letters journal homepage: Methane sources and production in the northern Cascadia margin gas hydrate system J.W. Pohlman a,, M. Kaneko b, V.B. Heuer c, R.B. Coffin d, M. Whiticar e a U.S. Geological Survey, Woods Hole Science Center, 384 Woods Hole Rd, Woods Hole, MA 02543, USA b Department of Earth and Planetary Sciences, Kyushu University, Fukuoka, , Japan c Organic Geochemistry Group, Dept. of Geosciences & MARUM Center for Marine Environmental Sciences, University of Bremen, D Bremen, Germany d Naval Research Lab, Marine Biogeochemistry, Washington D.C., 20375, USA e School of Earth and Ocean Sciences, University of Victoria, P.O. Box 3055, Victoria, BC, Canada V8W 2Y2 article info abstract Article history: Received 24 November 2008 Received in revised form 24 August 2009 Accepted 29 August 2009 Available online 26 September 2009 Editor: M.L. Delaney Keywords: methanogenesis methane stable isotopes gas hydrate Cascadia margin Integrated Ocean Drilling Program (IODP) The oceanographic and tectonic conditions of accretionary margins are well-suited for several potential processes governing methane generation, storage and release. To identify the relevant methane evolution pathways in the northern Cascadia accretionary margin, a four-site transect was drilled during Integrated Ocean Drilling Program Expedition 311. The δ 13 C values of methane range from a minimum value of 82.2 on an uplifted ridge of accreted sediment near the deformation front (Site U1326, 1829mbsl, meters below sea level) to a maximum value of 39.5 at the most landward location within an area of steep canyons near the shelf edge (Site U1329, 946 mbsl). An interpretation based solely on methane isotope values might conclude the 13 C-enrichment of methane indicates a transition from microbially- to thermogenically-sourced methane. However, the co-existing CO 2 exhibits a similar trend of 13 C-enrichment along the transect with values ranging from 22.5 to The magnitude of the carbon isotope separation between methane and CO 2 (ε c =63.8 ±5.8) is consistent with isotope fractionation during microbially mediated carbonate reduction. These results, in conjunction with a transect-wide gaseous hydrocarbon content composed of > 99.8% (by volume) methane and uniform δd CH4 values ( 172 ±8) that are distinct from thermogenic methane at a seep located 60 km from the Expedition 311 transect, suggest microbial CO 2 reduction is the predominant methane source at all investigated sites. The magnitude of the intra-site downhole 13 C-enrichment of CO 2 within the accreted ridge (Site U1326) and a slope basin nearest the deformation front (Site U1325, 2195mbsl) is ~5. At the mid-slope site (Site U1327, 1304 mbsl) the downhole 13 C-enrichment of the CO 2 is ~25 and increases to ~40 at the nearshelf edge Site U1329. This isotope fractionation pattern is indicative of more extensive diagenetic alteration at sites with greater 13 C-enrichment. The magnitude of the 13 C-enrichment of CO 2 correlates with decreasing sedimentation rates and a diminishing occurrence of stratigraphic gas hydrate. We suggest the decreasing sedimentation rates increase the exposure time of sedimentary organic matter to aerobic and anaerobic degradation, during burial, thereby reducing the availability of metabolizable organic matter available for methane production. This process is reflected in the occurrence and distribution of gas hydrate within the northern Cascadia margin accretionary prism. Our observations are relevant for evaluating methane production and the occurrence of stratigraphic gas hydrate within other convergent margins. Published by Elsevier B.V. 1. Introduction Marine gas hydrate occurs along continental margins under conditions of low temperature and moderate pressure (> m water depth) where the concentration of low molecular weight gases (primarily methane) exceeds pore water solubility Corresponding author. address: jpohlman@usgs.gov (J.W. Pohlman). (Kvenvolden, 1993; Xu and Ruppel, 1999). Enhanced primary production, efficient organic matter burial and tectonic fluid expulsion provide ideal conditions for producing methane and concentrating it within the gas hydrate stability field of convergent margins (Hyndman and Davis, 1992; Kastner, 2001). Accordingly, accretionary convergent margins stretching along ~29,000 km of oceanic subduction zones (von Huene and Scholl, 1991) may contain as much as twothirds of the marine gas hydrate reservoir (Kastner, 2001). Expedition 311 of the Integrated Ocean Drilling Program (IODP) was designed to investigate the systematic evolution of the gas hydrate system across the northern Cascadia margin (NCM) accretionary prism. To that end, a 33-km long transect of five boreholes X/$ see front matter. Published by Elsevier B.V. doi: /j.epsl

2 J.W. Pohlman et al. / Earth and Planetary Science Letters 287 (2009) (IODP Sites U1325 U1329), including an off-transect cold seep (Site U1328), was drilled from an uplifted ridge near the deformation front to a shallow near-shelf location at the landward limit of gas hydrate occurrence (Fig. 1). The margin-wide transect approach employed during Expedition 311 distinguishes it from previous Ocean Drilling Program (ODP) investigations (Legs 146 and 204) that targeted gas hydrate occurrences in this region. For example, Leg 204 focused on regional structural and stratigraphic gas hydrate accumulations around Hydrate Ridge, a large accretionary ridge offshore Oregon (USA) (Tréhu et al., 2004). A spatial and mechanistic understanding of methane production in convergent margins is critical for assessing the formation, occurrence and distribution of a large component of the global gas hydrate inventory. The NCM accretionary prism is largely comprised of sediment that has been scraped off the subducting Juan de Fuca plate since Eocene time (~43 Ma) (Hyndman, 1995). At the deformation front, layered Pleistocene hemipelagic sediments are folded and faulted into anticlinal ridges (e.g., the frontal ridge at Site U1326; Fig. 1B). Further tectonic compression of the accreted material results in a landward series of folds and thrusts. Sediment compaction during accretion drives upward migration of methane-charged fluids into the bottom of the gas hydrate stability zone and has been implicated as the primary cause for prominent bottom simulating reflectors (BSRs) (Fig. 1B) and the occurrence of gas hydrate within this accretionary prism (Hyndman and Davis, 1992). A surprising result from IODP Expedition 311 is that gas hydrate was often more prevalent near the top of the gas hydrate occurrence zone (GHOZ) and the thickness of the GHOZ decreased upslope of the deformation front (Malinverno et al., 2008). The GHOZ, the vertical interval where gas hydrate is present, is to be distinguished from the gas hydrate stability zone (GHSZ), where the conditions for gas hydrate stability are satisfied, but gas hydrate may not be present. The observed distribution and thickness of the GHOZ indicate that other factors, e.g., sediment grain size and production of methane within the gas hydrate stability zone, might also influence gas hydrate formation within the NCM (Malinverno et al., 2008; Torres et al., 2008). The primary source for methane thus far recovered from gas hydrate-bearing convergent margins is microbial methane produced by CO 2 reduction (also referred to as carbonate reduction or hydrogenotrophic methanogenesis) (Claypool et al., 1985; Kvenvolden and Kastner, 1990; Whiticar et al., 1995; Milkov et al., 2005). Variable, though generally limited, thermogenic contributions have also been identified in gas hydrate-bearing convergent margins (Claypool et al., 1985; Kvenvolden et al., 1990; Berner and Faber, 1993; Whiticar et al., 1995; Milkov et al., 2005; Pohlman et al., 2005). Whiticar et al. (1995) inferred the presence of thermogenic gases below the gas hydrate stability zone at ODP Leg 146 Site 889 (located 250 m from Site U1327) from the presence of higher order (C 2+ ) hydrocarbons and the stable carbon isotope signatures of methane and CO 2. At a high gas flux region on Hydrate Ridge (ODP Leg 204 Sites ), Milkov et al. (2005) determined that thermogenic methane transported along a high permeability subvertical horizon from a sediment depth of 2 to 2.5 km supplied as much as 20% of the methane in shallow gas hydrate accumulations. The presence of a gas hydrate-bearing hydrocarbon seep in Barkley Canyon (Pohlman et al., 2005), located 60 km from and at a similar water depth as Site U1329, is evidence for a potential thermogenic gas contribution along the Expedition 311 transect. The prevalent empirical framework for characterizing gas origins is that microbial gas is dominated by methane with δ 13 C signatures ranging from 90 to 60 and thermogenic gas contains higher concentrations of C 2+ hydrocarbons and methane with δ 13 C signatures ranging from 50 to 30 (Whiticar, 1999). In some cases, however, preferential consumption of 12 CO 2 during CO 2 reduction, i.e., the kinetic isotope effect (KIE), has been shown to enrich, or fractionate, the residual CO 2 pool with 13 CO 2 to the extent that methane generated from it acquires a δ 13 C signature characteristic for thermogenic methane (Claypool et al., 1985; Kvenvolden and Kastner, 1990; Whiticar et al., 1995). Likewise, acetoclastic methanogenesis (i.e., methane production by acetate fermentation) may also produce 13 C-enriched methane that could be interpreted as an admixture of microbial and thermogenic methane based on δ 13 C- values alone (Whiticar, 1999). Finally, it has been argued that nonthermogenic contributions of ethane, propane and, perhaps, butane are significant in convergent margins (Kvenvolden et al., 1990; Hinrichs et al., 2006). In this study, we utilize the stable isotope content of methane (δ 13 C and δd) and CO 2 (δ 13 C) as well as the hydrocarbon composition from dissolved gas samples to constrain methane origins along the IODP Expedition 311 transect. Furthermore, we evaluate factors that control organic matter preservation and discuss how the quality of the organic matter may influence the carbon isotope mass balance and occurrence of gas hydrate in the northern Cascadia margin accretionary prism. Fig. 1. A. Site map of the northern Cascadia margin (offshore Vancouver Island, Canada), the region studied during IODP Expedition 311. The shaded area covering about 50% of the continental slope is the region where the occurrence of gas hydrate has been seismically inferred from bottom simulating reflectors (BSR) (Spence et al., 2000). B. Four of the drill sites the focus of this study are located along the multichannel seismic line (panel A), which shows the seismic structure of the subducting oceanic plate and the accretionary prism. The BSR, an indicator of the bottom of the gas hydrate stability zone (GHSZ), is most prominent near the mid-slope Site U1327. Distances between the deformation front and sites along the transect are indicated above the image.

3 506 J.W. Pohlman et al. / Earth and Planetary Science Letters 287 (2009) Methods 2.1. Sampling Sediment cores recovered with the advanced piston corer (APC) and extended core barrel (XCB) were transferred to the catwalk for gas sampling. Gas expansion voids formed from porewater dissolved gases during core depressurization and temperature rise were extracted on the catwalk by piercing the clear core liner with a steel penetration tool connected via a 3-way valve to a 60 ml plastic syringe (Becton Dickinson). A total of 160 gas voids (also referred to as dissolved gas samples to more accurately reflect their in situ origins) sampled at an average frequency of one per 6m core penetration were collected for isotopic analysis along the Expedition 311 transect (see Table S1) Laboratory analysis The void gas concentrations (vol%) of methane through pentanes (C 1 C 5 ), oxygen, nitrogen and CO 2 were determined onboard the JOIDES Resolution using an HP 6890 gas chromatograph (GC) equipped with a 60 m 0.32 mm ID DB-1 capillary column, a thermal conductivity detector and a flame ionization detector (Pimmel and Claypool, 2001; Riedel et al., 2006). Gas concentrations were determined relative to certified standard gas mixtures with an analytical precision of 1%. The remaining syringe gas was transferred through an 18 g needle into evacuated 30 ml serum vials (Wheaton) sealed with 10 mm thick butyl septa and stored at 4 C until isotopic analysis. The stable carbon isotope compositions of methane from 160 samples and CO 2 from 83 samples were analyzed by irm-gc/ms using a Thermo Finnigan Trace GC ultra gas chromatograph (GC) coupled to a DELTA Plus XP isotope ratio mass spectrometer (IRMS) via a GC combustion III interface (900 C) and routine open split. For deuterium analysis, methane from 69 samples was pyrolyzed at 1440 C to molecular hydrogen and carbon. To obtain a signal for each compound within the working dynamic range of the IRMS (1 10 V at m/z 44), gas volumes ranging from 0.5 to 5000 μl were cryogenically concentrated ( 196 C) onto a 0.53 mm ID capillary column handpacked with 80/120 Poraplot-Q. GC separation with a 30 m 0.32 mm ID Restek Rt-Q Plot column for methane carbon analysis was performed at 10 C to separate methane from residual nitrogen. For CO 2 carbon isotope analysis and methane deuterium analysis, GC separation was achieved with an oven temperature of 50 C. The carbon and hydrogen isotope ratios are reported in δ-notation (δ 13 C, δd) relative to the VPDB and VSMOW standards, respectively. The 1σ precision for δ 13 CH 4 and δ 13 CO 2 is ±0.3 and for δd-ch 4 is ±1. distance from the deformation front (Fig. 2, Table S1). At all sites, near-surface methane is 13 C-depleted (δ 13 C CH4 range= 84.0 to 72.8 ), and deeper methane becomes progressively 13 C-enriched down each borehole. At the downslope Sites U1326 and U1325 (located ~5 and 11 km, respectively, from the deformation front, Fig. 1B) the δ 13 C CH4 gradually increases to near constant values of ±0.2 (n =9) and ±0.6 (n =8) for the bottom 50mbsf (meters below seafloor) (Fig. 2). Similar profiles have been reported from DSDP and ODP Legs at Blake Ridge (Borowski, 2004 and references therein) and Hydrate Ridge (Milkov et al., 2005) for sites not strongly influenced by migration conduits. In conjunction with elevated ratios of methane to ethane (C 1 /C 2 ), which range from 800 to (Fig. 3, Table S1), the δ 13 C CH4 values from U1326 and U1325 suggest a predominantly microbial origin for methane (Whiticar, 1999). At the mid-slope Site U1327 (~21 km from the deformation front) and the upslope Site U1329 (~38 km from the deformation front), the 13 C-depleted δ 13 C CH4 values in the shallow sedimentary sections contrast with δ 13 C CH4 values of 46.9 at mbsf and 39.5 at mbsf, respectively (Fig. 2). An interpretation based solely on the methane carbon isotopes from Sites U1327 and U1329 is that the profiles reflect mixing between a deep thermogenic and shallow microbial source. However, numerous microbially mediated processes may also yield methane with δ 13 C values that fall within the typical thermogenic range. For example: 1) methane oxidation preferentially consumes 12 CH 4, leaving the residual methane pool 13 C-enriched (Alperin et al., 1988); 2) methyl-type fermentation, e.g., acetoclastic methanogenesis, can generate methane with δ 13 C CH4 values as low as 50 (Whiticar et al., 1986); and 3) the KIE during CO 2 reduction can enrich (or fractionate) the residual CO 2 with 13 C such that methane produced from this source acquires a thermogenic-like signature (Claypool et al., 1985). We address each of these possibilities by evaluating the hydrocarbon composition, stable carbon isotopes and hydrogen isotopes from void gas samples to differentiate thermogenic and microbial contributions and determine how diagenetic alterations can influence the δ 13 C of the subsurface gases within and immediately below the base of the gas hydrate stability zone along the Expedition 311 transect Thermogenic sources C 2 C 5 hydrocarbons are present in the void gas from all sites, with a maximum concentration of 1150 ppm (v/v) for C 2 from 200mbsf at Site U1327 (Riedel et al., 2006). Nevertheless, methane is the dominant gas, accounting for >99.8% of the hydrocarbon content of all samples from all sites (Riedel et al., 2006). The majority of the 2.3. Sedimentation rate calculations Sedimentation rate calculations for the potential gas hydrate occurrence zone (pghoz) the sediment interval between the base of the sulfate reduction zone (SRZ) and the base of GHSZ is based on the diatom biostratigraphic zones identified during Expedition 311 (Akiba et al., 2009). The average sedimentation rates for the pghoz were calculated as the depth difference between the base of GHSZ and the SRZ divided by the age difference between the same horizons. Maximum sedimentation rates are those reported by Akiba et al. (2009) for the Ma interval representing North Pacific Diatom Zone 11 (NPD11). The thickness of that interval varied from 32 m (Site U1329) to 140 m (Site U1326). 3. Results and discussion 3.1. Methane origins in the northern Cascadia margin The δ 13 C CH4 profiles of void gas along the transect exhibit distinct patterns of 13 C-enrichment with increasing sediment depth and Fig. 2. Depth profiles of δ 13 C CH4 and δ 13 C CO2 from the IODP Expedition 311 transect. The dashed lines intersecting the δ 13 C CH4 profiles represent the depths of the bottom simulating reflector (BSR).

4 J.W. Pohlman et al. / Earth and Planetary Science Letters 287 (2009) Fig. 3. Relationship between the C 1 /C 2 ratios and δ 13 C values of methane for interpreting gas sources (adapted from Bernard et al., 1978). The mixing curve (dashed line) was calculated from the most prevalent microbial endmember in this study (δ 13 C= 70,C 1 /C 2 =10 6 ) and the thermogenic endmember from the nearby Barkley Canyon hydrocarbon seep (δ 13 C= 30, C 1 /C 2 =81; Pohlman et al., 2005). paired δ 13 C CH4 and C 1 /C 2 data from Sites U1327 and U1329 are outside the expected range of microbial sources, thermogenic sources and the standard mixing scenario between the two endmembers (Fig. 3) (Bernard et al., 1978). However, the source fields in Fig. 3 are constructed from data acquired from representative freshwater and marine systems. Secondary effects resulting from oxidation, gas migration, substrate depletion and advanced source rock maturity generate values that fall outside the field ranges. For example, similar combinations of elevated C 1 /C 2 ratios and δ 13 C CH4 are possible for gas sources that have migrated from source rocks in advanced stages of maturation (Coleman et al., 1977; Schoell, 1980). To differentiate if the combination of elevated C 1 /C 2 ratios and 13 C- enriched CH 4 values represents a dry (methane dominated) thermogenic gas source, we consider the corresponding methane hydrogen isotope data (δd CH4 ). Dry thermogenic methane produced from a mature marine or humic source rock is characterized by δd CH4 values that are typically more enriched than 150 (see Fig. 4 for dry thermogenic gas δ 13 C CH4 and δd CH4 fields from Schoell, 1980). Consistent with that trend, thermogenic methane from Barkley Canyon, a hydrocarbon seep located ~60 km from Site U1329, has an average δd CH4 of 140 ±2 (n=4) (Pohlman et al., 2005). By contrast, the average δd CH4 for all samples acquired along the Expedition 311 transect is 172±8 (n =69; Table S1). Dissimilar δd CH4 values for the Expedition 311 transect methane in comparison to known dry thermogenic gas sources and Barkley Canyon indicate thermogenic methane is not a significant constituent of the gas inventory within the gas hydrate stability zone of the Expedition 311 transect Microbial alterations and sources Methane oxidation. Preferential utilization of the lighter isotope 12 C during anaerobic methane oxidation (AOM) is capable of enriching the residual methane in 13 C to attain the δ 13 C values we measured (Alperin et al., 1988). However, methane oxidation is eliminated as a possibility because at all sites the 13 C-enriched methane is below the depth of the SRZ (Riedel et al., 2006), which is the environment where AOM occurs (Hoehler et al., 1994). Fig. 4. C D diagram for interpreting gas sources. The shaded areas indicate fields for microbial endmembers, thermogenic endmembers and mixing regions for the endmember fields after Whiticar (1999). The position of the fields for dry thermogenic gas (Schoell, 1980) and the Barkley Canyon thermogenic endmember (Pohlman et al., 2005) in relation to the sample pairs are not consistent with a dry thermogenicmicrobial methane mixing system. A more suitable explanation is carbon isotope fractionation (effect indicated by arrow) and constant deuterium fractionation of formation water during microbial CO 2 reduction Methyl-type fermentation. Acetoclastic methanogenesis is a methyl-fermentation pathway that produces methane with δ 13 C values as 13 C-enriched as 50 (Whiticar et al., 1986). The stable carbon isotopic composition of acetate at one transect location (Site U1329), at least, suggests acetoclastic methanogenesis (Heuer et al., 2009). However, an increasing contribution from acetoclastic methanogenesis as an explanation for the δ 13 C CH4 variability within the margin is unlikely. The δd CH4 synthesized during acetoclastic methanogenesis values is controlled primarily by the deuterium (D) content of the acetate methyl group and fractionation during its conversion to methane, and secondarily by the δd of the formation water (Whiticar et al., 1986). The KIE during fermentative acetogenesis causes extensive D-depletion in the acetate (Whiticar et al, 1986). Empirical observations of δd CH4 values from systems where acetate is the dominant methane source range from approximately 280 to 400 (Whiticar et al., 1986). Every δd CH4 value measured during this study (Fig. 4, Table S1) is enriched with deuterium by more than 100 relative to the maximum value in that range. Furthermore, the δd values are indistinguishable among sites (Fig. 4), which suggests a common source, not a systematic transition between processes, as the δ 13 C variability might imply CO 2 reduction. Constant δd CH4 values in tandem with systematic 13 C-enrichment in the methane and CO 2 are evidence that CO 2 reduction is the primary source of methane within the Expedition 311 transect. For methane formed via CO 2 reduction in marine sediment, a simple linear relationship exists between the δd CH4 and the formation water (δd H2 O), the primary source of hydrogen in methane formed during carbonate reduction (Daniels et al., 1980, Whiticar et al., 1986, Valentine et al., 2004), δd CH4 = δd H2 O 180ð 10 Þ: ð1þ The δd H2 O of global seawater has been relatively constant at ~0 (Lécuyer et al., 1998) within the 43 Ma period that the NCM has been developing (Hyndman, 1995). Thus, methane formed by CO 2 reduction in the NCM accretionary prism would be expected to have δd values within the range of 170 to 190. Indeed, the δd CH4 of -172 ±8 (n=69) measured in this study (Table S1) is relatively constant and within the range expected for CO 2 reduction. The constant δd values relate back to the effectively limitless pool of formation water supporting CO 2 reduction. In contrast, the δ 13 Cof the CO 2 supporting CO 2 reduction is much more limited in size and thus may be substantially fractionated during CO 2 reduction. Fractionation resulting from the KIE is frequently expressed as the kinetic isotope fractionation factor, α c,defined as α c = 12 k = 13 k; ð2þ

5 508 J.W. Pohlman et al. / Earth and Planetary Science Letters 287 (2009) where 12 k and 13 k are the rate constants for 12 CO 2 and 13 CO 2 reduction to methane, i.e., 12 CO2 12 k 12 CH4 and CO 2 k 13CH4 : ð3þ As 12 CO 2 reacts slightly faster than 13 CO 2 during CO 2 reduction (i.e., 12 k> 13 k), α c ranges from ~1.04 to ~1.09 (Valentine et al., 2004; Penning et al., 2005). The kinetic fractionation factor, α c, may be expressed as the instantaneous isotope separation factor ε c, where ε c =10 3 ðα 1Þ δ 13 C CO2 δ 13 C CH4 : ε c approximates the fractionation between the instantaneously produced methane and its substrate CO 2 during carbonate reduction as the δ 13 C difference between CO 2 CH 4 pairs. For this discussion, we utilize ε c because it allows us to describe how the KIE influences the carbon isotope ratios of CO 2 and methane with the δ 13 C values of the CO 2 CH 4 pairs measured in this study (Table S1). The magnitude of downhole 13 C-enrichment of methane at each site is similar to that of the CO 2 (Fig. 2). The average δ 13 C difference between methane and CO 2, ε c, at each site along the transect (61.1 to 66.2) is consistent with the range of 65 to 75 typically observed for carbonate reduction in marine sediments (Whiticar, 1999). While the downhole 13 C-enrichment trends are easily discerned in the δ 13 C CH4 - profiles, scatter in the δ 13 C CO2 profiles, particularly for Sites U1325 and U1326, partially obscures the downhole 13 C-enrichment in the CO 2 profiles. This variability is an artifact of equilibrium isotope fractionation that occurs during degassing of CO 2 during core recovery (Emrich et al., 1970). Given that variability, we do not attempt to interpret the downhole and cross-margin differences in ε c from this dataset. We assume constant isotopic fractionation between the CO 2 and methane (Whiticar, 1999). When viewed as a closed system, the site-specific 13 C-enrichment of CO 2 and methane can be described by a simplified version of the Rayleigh distillation function that incorporates ε c (Whiticar, 1999): δ 13 C CO2 ;t = δ 13 C CO2 ;i ε c ln f δ 13 C CH4 ;t = δ13 C CO2 ;i ε cð1+ lnf Þ; ð6þ where δ 13 C CO2,t and δ 13 C CH4,t represent the isotope ratio for the residual CO 2 and accumulated product methane, respectively, at time t when fraction f of the initial CO 2 remains. The value for δ 13 C CO2,i is approximately the δ 13 C of remineralized sedimentary organic matter (Blair and Aller, 1995). As the fraction f of the initial CO 2 substrate pool decreases, the δ 13 C of the residual CO 2 (δ 13 C CO2,t) and product methane (δ 13 C CH4,t) become progressively more 13 C-enriched. Thus, the extent of fractionation that occurs within a closed system is a methane production effect whereby more 13 C-enrichment indicates greater substrate depletion of the initial CO 2 pool. To estimate the fraction of the initial CO 2 pool consumed by methanogenesis (1 f) at each site, we evaluate Eq. (5) using these inputs: 1) 25.4 for δ 13 C CO2,i, the average δ 13 C of sedimentary organic matter along the Expedition 311 transect (Kim and Lee, 2009), 2) the average ε c from CO 2 CH 4 pairs from each site as calculated by Eq. (4), and 3) the average δ 13 C CO2 from the bottom 50 m of each site for δ 13 C CO2,t (Table S1). According to the model results, the 13 C- enrichment of CO 2 along the Expedition 311 transect represents consumption of ~20 22% of the initial CO 2 at the seaward Sites U1326 and U1325, ~41% at the mid-slope Site U1327 and ~52% at the landward Site U1329. These model results demonstrate that CO 2 reduction is capable of driving the observed fractionation of CO 2 and methane within the NCM, in situ methanogenesis is occurring within the GHSZ and a larger fraction of the initial CO 2 pool has been ð4þ ð5þ converted to CH 4 at the landward locations. Similar CH 4 and CO 2 values have been reported from slope sediments of the Middle America Trench (Claypool et al., 1985); however, the extent to which the CO 2 and methane are 13 C-enriched in the NCM is extreme with respect to the carbonate reduction field boundaries frequently utilized to differentiate methane production pathways and oxidative effects in marine and freshwater systems (Whiticar et al., 1986; Whiticar, 1999) (Fig. 5). Because the natural system is open to additional inputs and outputs of CO 2 (Claypool et al., 1985; Wigley et al., 1978), the closed system Rayleigh distillation model is a crude approximation of the carbon isotope mass balance. For example, from balanced organic matter fermentation (Eq. (7)) and CO 2 reduction (Eq. (8)), equivalent amounts of CO 2 and methane are produced (Eq. (9)), 2CH 2 O þ 2H 2 O 2CO 2 þ 4H 2 ðfermentationþ ð7þ CO 2 þ 4H 2 CH 4 þ 2H 2 O ðco 2 ReductionÞ ð8þ 2CH 2 O CO 2 þ CH 4 ðnetþ ð9þ In terms of the CO 2 isotope mass balance, the addition of CO 2 with a δ 13 C similar to organic matter (~25 ) will diminish the magnitude of the isotope separation factor, ε c, resulting from CO 2 reduction (Eq. (8)). Furthermore, precipitation of calcite (CaCO 3 ) and dolomite (CaMg(CO 3 ) 2 ) removes carbonate (CO 3 2 ) from the total CO 2 pool, which may limit concentration dependent rates of methane production (Claypool et al., 1985). Both processes counteract the 13 C- enrichment of CO 2 resulting from CO 2 reduction. Thus, the closed system approach taken here places a lower bound on the fraction of the initial CO 2 pool removed from the system. To better constrain the carbon mass balance within this system requires an integrated diagenetic model that includes open-system Rayleigh equations (Wigley et al., 1978) with components for organic matter degradation (Berner, 1980; Canfield, 1994; Middelburg, 1989), alkalinity/carbonate production, methane production, diffusion and fluid advection (Wallmann et al., 2006). Fig. 5. Relationship between paired (co-existing) δ 13 C CH4 and δ 13 C CO2 values for interpreting gas sources and isotopic shifts resulting from production and oxidation (adapted from Whiticar, 1999). The diagonal lines indicate carbon iso-fractionation lines. The line connecting the paired samples from Site U1329 demonstrates a depthdependent production effect resulting from substrate depletion. Near-surface samples plot in the 13 C-depleted area and shift to more 13 C-enriched values downcore. The dashed line extends the CO 2 reduction field to account for the values measured in this study. The margin-scale trajectory of the CO 2 CH 4 sample pairs depends upon the location of each site relative to the deformation front, which is consistent with greater substrate depletion at sites landward of the deformation front.

6 J.W. Pohlman et al. / Earth and Planetary Science Letters 287 (2009) Organic matter constraints on gas hydrate distribution Gas hydrate along the Expedition 311 transect Downhole log and chloride data from Expedition 311 (Malinverno et al., 2008) revealed a more complicated distribution of gas hydrate than predicted by the fluid expulsion model (Hyndman and Davis, 1992) or recycling of methane at the bottom simulating reflector (BSR) (Dillon and Paull, 1983). Rather than being concentrated at the bottom of the stability field near the BSR, gas hydrate is widely distributed within the GHSZ. The region of the GHSZ where gas hydrate is present (i.e., the gas hydrate occurrence zone, GHOZ), decreases in thickness from 217 m at Site U1326 (nearest the deformation front) to 119 m at the mid-slope Site U1327 and is absent at the landward Site U1329 (Table 1, data from Malinverno et al., 2008). Malinverno et al. (2008) developed a simple diagenetic model that utilized different combinations of fluid advection velocities and methane generation rates to match the distribution of gas hydrate within the stability field. Attaining model methane concentrations needed to form gas hydrate in the upper portion of the GHSZ required higher rates of in situ methane generation or higher advective methane fluxes as would be generated by higher rates of fluid advection. Consistent with advection as a mechanism for transporting methane into to the GHSZ from greater depth (Hyndman and Davis, 1992; Ruppel and Kinoshita, 2000), the highest gas hydrate saturation at Site U1325 occurs immediately above the base of the GHSZ (Malinverno et al., 2008). Fluid expulsion models of the northern Cascadia accretionary margin that utilize different porositydepth functions predict maximum fluid advection velocities at the deformation front in one case (Wang et al., 1993) and maximum velocities ~30 km landward from the deformation front in the other (Hyndman et al., 1993). While the observed distribution is more consistent with the Wang et al. (1993) results, additional research is required to constrain to what extent fluid advection influences the distribution of gas hydrate across this accretionary prism. With respect to the potential role that in situ methane generation plays in the distribution of gas hydrate, sedimentation rates that decrease with distance from the deformation front (Akiba et al., 2009) are likely to affect the quality of organic matter buried within the accretionary prism and, hence, in situ methane production. The CO 2 and CH 4 methane carbon isotope data discussed in Section 3.1 strongly suggest in situ production of methane within the gas hydrate stability zone. The downhole magnitude of CO 2 fractionation resulting from in situ methanogenesis from each site correlates with the average sedimentation rate in the GHSZ (r 2 =0.81) and the thickness of the GHOZ (r 2 =0.92) across the margin (data from Table 1). Thus, factors that control in situ methane production, e.g., organic matter sources and burial preservation, may also be related to the occurrence and distribution of gas hydrate Sedimentary organic carbon content The average bulk sedimentary organic carbon content for all sites (0.64±0.23 wt.%, Riedel et al., 2006) is about 50% less than that from the Blake Ridge gas hydrate system (Paull et al., 1996), but theoretically sufficient to produce enough methane to saturate the pore water, as required for gas hydrate formation. For example, Clayton (1992) calculated that 0.2% TOC in a 1 km deep sediment system is sufficient to saturate the pore water, and Kastner (2001) reasoned that mineralization of 0.1% of marine organic carbon can concentrate methane to ~200 mm, which exceeds the solubility concentration of methane within the entire GHSZ of the NCM (Riedel et al., 2006). Along the transect, as the thickness of the GHOZ decreases (Malinverno et al., 2008), the site-specific average TOC content increases systematically from 0.40 wt% at the seaward Site U1326 to 0.88 wt% at landward Site U1329 (Table 1, data from Riedel et al., 2006). If in situ methane production is an important source of methane within the GHOZ, as the carbon isotope ratios of CO 2 and methane from this study and other studies suggest (Kvenvolden and Barnard, 1983; Claypool et al., 2006), the increasing wt% of TOC (the base substrate pool for methanogenesis) along a transect where the abundance of gas hydrate diminishes implies that the bulk organic carbon concentration is not the major factor controlling organic matter reactivity in this setting. Indeed, other studies suggest that quality not quantity defines the reactivity of organic matter (Canfield, 1994) and how rapidly methane may be generated from it (Penning and Conrad, 2007). Along the Expedition 311 transect, the relationship between high sedimentation and low TOC most likely reflects increased dilution of organic matter by inorganic clastic material transported farther offshore by turbidity currents. Gas hydrate resource assessment models that utilize the wt% (or quantity) of the sediment organic matter when calculating methane production (e.g., Frye et al., 2008) might consider this effect Preservation and reactivity controls The reactivity, or quality, of sedimentary organic matter is fundamentally controlled by its chemical composition (Hedges and Keil, 1995), matrix effects that regulate the bioavailability of the metabolizable organic matter (Hedges and Keil, 1999) and the history of exposure to different redox conditions (Canfield, 1994). Because labile (more reactive) compounds tend to be degraded before refractory (less reactive) compounds, the diagenetic exposure time is a useful and robust metric for reactivity (Berner, 1980; Middelburg, 1989). The exposure time of the organic matter is often parameterized by the sedimentation rate (Berner, 1980; Canfield, 1994; Malinverno et al., 2008), which determines how long the organic matter was exposed to various redox conditions. Oxic respiration, nitrate reduction, iron and manganese reduction, sulfate reduction and methane production all contribute to organic matter degradation, but sulfate reduction and oxic respiration are thought to dominate in normal sedimentary marine environments (Henrichs and Reeburgh, 1987; Hartnett et al., 1998). The average δ 13 C of sedimentary organic matter along the Expedition 311 transect is 25.4 ±1.0, which suggests a fairly uniform admixture of marine and terrigenous material (Kim and Lee, Table 1 Site information, sedimentation rates and organic matter residence times for the IODP Expedition 311 transect. Site Water depth a Distance from deformation front a Average TOC a Base GHSZ a Thickness of the GHOZ b Sedimentation rate in pghoz Base of SRZ a Time in SRZ c Thickness of pghoz d Time in pghoz c (m) (km) (wt%) (mbsf) (m) (m/ma) (mbsf) (ka) (m) (Ma) U U U U TOC: total organic carbon; GHSZ: gas hydrate stability zone; SRZ: sulfate reduction zone; pghoz: potential gas hydrate occurrence zone. a From Riedel et al. (2006). b From Malinverno et al. (2008). c From Eq. 10. d pghoz (m)=ghsz (m) SRZ (m).

7 510 J.W. Pohlman et al. / Earth and Planetary Science Letters 287 (2009) ). Thus, the relative reactivity of organic matter buried in the methanogenic zone across the margin is more likely a reflection of its depositional and post-burial diagenetic history, not the chemical composition of the source organic matter. At sedimentation rates below ~10 m/ma, prolonged exposure to aerobic oxidation leaves scant labile organic matter for supporting anaerobic metabolism (Canfield, 1994). Sedimentation rates in the range of m/ma in deep marine environments are characterized by some degree of sulfate reduction, and those in excess of 50 m/ma by both sulfate reduction and methanogenesis (G. Claypool, pers. comm.). Canfield (1994) suggests that at sedimentation rates above ~300 m/ma, most of the organic matter is preserved for utilization by anaerobic processes. The average sedimentation rate for material within the potential gas hydrate occurrence zone (pghoz) decreases across the margin from 263 m/ma at Site U1326 near the deformation front to 80 m/ma at the most landward Site U1329 (Fig. 6, Table 1). We define the pghoz as the region of the GHSZ where methanogenesis is the favored redox process and describe it as the region of the GHSZ where gas hydrate may occur. The thickness of the pghoz is the difference between the thickness of the GHSZ and the thickness of the SRZ. The pghoz is differentiated from the GHOZ because gas hydrate does not occur within the SRZ. We emphasize sedimentation rates in the pghoz because organic matter preserved within that horizon would support in situ methane production that could be incorporated into gas hydrate. All of the sediment along the transect was deposited at average sedimentation rates exceeding 50 m/ma (Fig. 6), the threshold sedimentation rate above which metabolizable organic matter preservation is enhanced. Sediments within the pghoz at Sites U1326 and U1325 that were deposited at rates in the range of m/ma should have resulted in even more enhanced preservation of metabolizable organic matter (Canfield, 1994). To demonstrate how exposure time to oxic and sulfate reducing conditions may have influenced preservation of organic matter presently within the pghoz, the average residence time that organic Fig. 6. Sedimentation rate controls on organic matter preservation in the northern Cascadia margin accretionary prism. Average sedimentation rates within the potential gas hydrate occurrence zone (pghoz) decrease with distance from the deformation front. High sedimentation rates at Site U1326 and U1325, particularly those in excess of 300 m/ma (indicated by dashed line), limit the exposure of organic matter to aerobic conditions, thereby preserving labile organic matter for anaerobic processes (Canfield, 1994), including methanogenesis. matter spent within the aerobic oxidation-sulfate reducing redox zone is determined as: T rz = z rz v ; ð10þ where T rz is the residence time in the redox zone, z rz is the thickness of the redox zone and v is the average sedimentation rate for the pghoz. This calculation ignores compaction (Berner, 1980) and assumes that z rz has been constant (steady-state conditions) over the past 2 Ma, the time interval over which organic matter now in the pghoz was deposited (Akiba et al., 2009). The combination of sedimentation rates that progressively decrease from 270 m/ma to 80 m/ma and SRZ depths that increase from ~2.9mbsf to ~10.9mbsf along the transect (Table 1) leads to a dramatic increase in residence times within the oxic and sulfate reducing redox zones at the upslope sites (Fig. 7). The estimated residence times evaluated by Eq. (10) increase from ~10 ka at Site U1326 to ~100 ka at Site U1329 (Table 1). This order of magnitude increase in exposure time to oxic and sulfate reducing redox conditions is likely to have had a substantial influence on the quality of sedimentary organic matter preserved along the Expedition 311 transect. More labile organic matter supports higher methane production rates (Penning and Conrad, 2007) and is a potential explanation for the positive correlation between the thickness of the GHOZ (Malinverno et al., 2008) and average sedimentation rates within the pghoz (r 2 =0.97, data from Table 1). In contrast to the SRZ, which becomes deeper (or thicker) with distance from the deformation front, the pghoz becomes thinner, primarily as a result of decreasing hydrostatic pressure at sites with shallower water depths (Table 1). Consequently, despite decreasing upslope sedimentation rates, the residence time (T rz ) for organic matter in the methanogenic redox zone of the pghoz, as calculated by Eq. (10), remains relatively constant (1.0 to 1.3 Ma) at all sites (Fig. 7). Milkov et al. (2003) speculated that Blake Ridge may contain more gas hydrate than Hydrate Ridge sites with stratigraphic hydrate occurrences because a thicker pghoz at Blake Ridge allowed methane to accumulate over a longer time period. Similar residence times for organic matter in the pghoz across the NCM indicate an alternate factor that limits the occurrence of gas hydrate at the upslope locations, where shallower water depths require less methane to saturate the porewater. The upslope trend of increased exposure of organic matter to oxic and sulfate reducing conditions during burial suggests the organic matter residing in the pghoz is more highly degraded at the upslope sites, and, therefore, contains less metabolizable organic matter for supporting methanogenesis and, ultimately, gas hydrate formation. Substrate limitation quantified by the extent of 13 C-enrichment in the residual CO 2 and product methane may also be linked to the preservation and reactivity of the organic matter along the margin. The cross plot of δ 13 C data from CO 2 CH 4 sample pairs (Fig. 5) reveals that substrate depletion associated with methane production occurs both in the vertical (downhole) and horizontal (upslope) directions. Downhole 13 C-enrichment of the sample pairs (illustrated in Fig. 5 by the solid line connecting sample pairs from increasing sediment depth for Site U1329) is affiliated with sediment organic matter that becomes older and, presumably, more degraded (Whiticar, 1999). Similar isotopic trends have been identified at individual sites within other convergent and passive margin settings with gas-charged and gas hydrate-bearing sediment (Claypool et al., 1985; Kvenvolden and Kastner, 1990; Berner and Faber, 1993; Whiticar et al., 1995; Paull et al., 2000; Milkov et al., 2005). An unusual observation is that the same 13 C-enrichment pattern that occurs downhole, also occurs across the margin. Pairs from the landward sites U1325 and U1325 cluster in the region of the CO 2 reduction field where the production effect is less pronounced, while pairs from Sites U1327 and U1329 extend into regions where the effect of substrate depletion driven by methanogenesis is greater (Fig. 5). The similarity between the downhole and

8 J.W. Pohlman et al. / Earth and Planetary Science Letters 287 (2009) Joint Oceanographic Institutions (JOI), Inc. Funding for this research was provided by the U.S. Science Support Program (USSP), and Natural Science and Engineering Research Council (MJW). We thank the Captain and the crew of the Joides Resolution and the technical staff for their support at sea; in particular, C. Bennight and L. Brant for tireless assistance in the laboratory. Guidance and support from Tim Collett, Michael Riedel and Mitch Malone are appreciated. We also acknowledge R. Plummer and Paul Eby for technical support and assistance with the isotopic analysis. We thank Jim Bauer, George Claypool, Carolyn Ruppel and Bill Waite for productive discussions and valuable comments on an earlier version of the manuscript. Any use of a trade, product, or firm name is for descriptive purposes only and does not imply endorsement by the U.S. Government. Appendix A. Supplementary data Fig. 7. Residence time of organic matter (OM) within the sulfate reduction zone (SRZ) and potential gas hydrate occurrence zone (pghoz) along the Expedition 311 transect. The residence time of organic matter within the SRZ increases by an order of magnitude from Site U1326 to Site U1329 as a result of decreasing sedimentation rates (Fig. 6, Table 1)and increasing SRZ thickness (Table 1). Prolonged exposure to sulfate reduction leads to burial of more refractory organic matter in the pghoz at sites distal to the deformation front. By contrast, the residence time of organic matter in the pghoz remains constant because decreasing sedimentation rates are offset by a decreasing pghoz thickness and increasing SRZ thickness (Table 2). Because the residence time for organic matter in the pghoz is similar at all sites, more methane and, hence, gas hydrate is expected to accumulate at sites where organic matter preservation is enhanced by short SRZ residence times. upslope patterns of 13 C-enrichment (Fig. 5) is consistent with the conclusion that variable sediment deposition rates (Fig. 6) influence organic matter preservation (Fig. 7), which, in turn, influences spatial patterns of methane production and, perhaps, the distribution of gas hydrate within the margin. 4. Conclusions The prevalence of gas hydrate concentrated near the top of the gas hydrate occurrence zone in the NCM is not consistent with the fluid expulsion model, which states that gas hydrate in accretionary prisms is concentrated where upwardly migrating methane-charged fluids enter and penetrate the gas hydrate stability field. Alternative explanations that explain the recent observation of a more widespread gas hydrate occurrence within the gas hydrate stability zone are that the expelled fluids migrate to shallower depths (where pressure is lower) before the dissolved methane concentrations exceed solubility, or that methane produced in situ is the major methane source. Analysis of the isotopic and hydrocarbon composition of gas extracted from the sediments along the IODP Expedition 311 transect suggests that microbial methane generated by CO 2 reduction is the primary hydrocarbon at all sites and that a predominant fraction of the methane is formed within the GHOZ. Additionally, the quality of the organic matter, which is affected by variable residence times of sedimentary material within different redox zones during burial along the margin, may influence if and where sufficient methane can be produced in situ to form gas hydrate. In concert with the fluid expulsion model, our organic matter quality model complements other in situ production models to provide a more comprehensive explanation for the geological and biogeochemical factors that control the occurrence and distribution of gas hydrate within accretionary margins. Acknowledgements Samples and data were provided by the Integrated Ocean Drilling Program (IODP), which is funded by the U.S. National Science Foundation and participating countries under management of the Supplementary data associated with this article can be found, in the online version, at doi: /j.epsl References Akiba F., Inoue Y., Saito M., Pohlman, J., Data report: diatom and foraminiferal assemblages in Pleistocene turbidite sediments from the Cascadia Margin (IODP Expedition 311), Northeast Pacific. In: Riedel, M., Collett, T.S., Malone, M.J., and the Expedition 311 Scientists (Eds.). 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