INTRODUCTION

Background

Ocean Drilling Program (ODP) Leg 164 was devoted to investigating amounts and in situ characteristics of natural gas hydrates contained in marine sediments. Sites 994, 995, and 997 were drilled along a ~10-km-long transect across the crest of the Blake Ridge (Fig. 1) to depths of 700-750 mbsf. A well-developed bottom simulating reflector (BSR) occurs under Sites 995 and 997, whereas Site 994 has no underlying BSR. Sites 994, 995, and 997 were drilled through the depth of the acoustically detected BSR (~450 m). Four short-hole sites (total depth of 49-63 mbsf) were also drilled along the flanks and crest of the Cape Fear (Sites 991, 992, and 993) and Blake Ridge diapirs (Site 996; Fig. 1; Paull, Matsumoto, Wallace, et al., 1996).

Gas hydrate nodules were recovered, the largest of which was a ~30-cm-thick horizon of massive gas hydrate. Shipboard gas hydrate decomposition experiments revealed that evolved gases are composed of ~99% CH₄ and ~1% CO₂ (Lorensen et al., Chap. 25, this volume).

Sediments from the three deep sites (Sites 994, 995, and 997) and Site 996 are very gassy and undergo vigorous expansion during core recovery. Because gas hydrates actively decompose during core recovery and the curation process, and because obvious gas loss occurs before sampling, various proxy measurements are used to define the in situ distributions of gas and gas hydrate. Pore-water profiles from the three sites show freshening of chloride to depths of ~200 mbsf. From 200 to 450 mbsf, chloride concentrations are highly variable, and characterized by local, anomalous excursions toward lower values. These anomalies indicate that variations of up to 14% in the amount of gas hydrate contained in adjacent samples occur throughout this zone and that finely disseminated gas hydrates, on average, occupy more than 1% of the sedimentary section between 200 and 450 mbsf at all three sites, regardless of whether a BSR is present (Sites 995 and 997) or not (Site 994). Well logs show distinct zones of higher electrical resistivity and sonic velocity that coincide with chloride anomaly occurrences (Collett and Ladd, Chap. 19, this volume). Vertical seismic profiles indicate that the velocity of sediments above the BSR are not significantly elevated above normal sediment velocities. However, acoustic velocities as low as 1400 m/s were measured immediately beneath the BSR at Site 997 (Holbrook et al., 1996) and are attributed to the accumulation of free gas.

Methane gas volumes measured using the ODP pressure core sampler (PCS) tool are in excess of in situ CH₄ gas saturation in pore water, demonstrating that free gas exists intermittently throughout the sedimentary section below the base of gas hydrate stability. PCS data indicate that the amount of free gas in the sediments beneath the base of gas hydrate stability may exceed the volume stored in the hydrates above (Dickens et al., 1997).

The results of Leg 164 confirm that marine gas hydrates and the associated dissolved and gaseous phases represent a major methane reservoir. In this paper, we consider the origins of interstitial CH₄ and associated interstitial CO₂ and their relationship to sedimentary organic matter.

Sources of Interstitial Low Molecular-Weight Hydrocarbon Gases and CO₂

In normal marine sediments, distinct sequences of microbial zones occur with sediment burial that influence methane accumulation. In the upper portion of the sedimentary section, microbial CH₄ production is at first inhibited by the presence of sulfate (Martens and Berner, 1974). Neither significant CH₄ production nor accumulation will occur until the sulfate is substantially depleted. Microbial CH₄ production commences near the base of the sulfate-reduction zone, usually resulting in increasing pore-fluid CH₄ concentrations with increasing depth. This effect is caused by local CH₄ production and the addition of CH₄, which migrates from greater depths. If CH₄ comes from greater depths, either by diffusion or advection, it may have been produced by either microbial processes or thermal hydrocarbon cracking (Tissot and Welte, 1984).

Methane

Microbial methane is produced primarily by CO₂ reduction (CO₂ + 4H₂ CH₄ + 2H₂O) and acetate fermentation (CH₃COOH CH₄ + CO₂) with CO₂ reduction primarily dominate in marine systems (Fenchel and Blackburn, 1979). CO₂ reduction is dependent on a supply of dissolved H₂, whereas acetate fermentation is usually limited by the amount of acetate available. Both limiting reactants are derived in part from the breakdown of larger organic molecules comprising sedimentary organic matter. Considerable carbon isotope fractionation (commonly 60-70) is associated with methanogenesis (e.g., Rosenfeld and Silverman, 1959; Whiticar et al., 1986).

Thermogenic methane is produced as a consequence of thermocatalytic degradation of kerogen at temperatures over ~120° C (Tissot and Welte, 1984), which, with typical geothermal gradients, requires burial depths significantly greater than ~1 km. Because less fractionation occurs during thermal cracking of kerogen (Schoell, 1983), the carbon isotopic composition of CH₄ produced from thermal cracking of kerogen is closer to the carbon isotopic composition of its parent sedimentary organic matter.

Interstitial CO₂

Interstitial dissolved CO₂ may be derived from organic sources, which include anaerobic oxidation of organic matter, and abiogenic thermal decarboxylation reactions (Claypool and Kaplan, 1974). The isotopic composition of the CO₂ that is added to the dissolved inorganic carbon (DIC) pore-water pool by these processes will reflect that of the organic matter from which it is derived (Presley and Kaplan, 1968). Carbon in the DIC pool may also be supplied from inorganic sources, primarily carbonate dissolution. The isotopic composition of carbon from dissolution of biogenic carbonate rarely varies more than a few parts per million (~2 ppm) from 0 (Anderson and Arthur, 1983).

Dissolved CO₂ sinks include methanogenesis (CO₂ reduction) and authigenic carbonate formation (Rodriguez et al., Chap. 30, this volume). Authigenic carbonates are precipitated at or near isotopic equilibrium with the dissolved pore-water species (Anderson and Arthur, 1983). However, methanogenesis preferentially uses ¹²CO₂ to form methane, shifting the remaining CO₂ pool toward enriched ¹³C values.

Low Molecular-Weight Hydrocarbon Gases

The ratio of low molecular-weight hydrocarbon gases (e.g., methane/ethane ratio) and ¹³CCH₄ values are commonly used to establish the methane's source. Microbial CH₄ typically has ¹³CCH₄ values that range from -90 to -55 (PDB) and methane to ethane ratios of >1000 (Bernard et al., 1977). Thermogenic CH₄ typically has ¹³CCH₄ values more positive than -55 and methane to ethane ratios less than 100 (Bernard et al., 1977). In addition, DCH₄ values are used to distinguish between the two major microbial methane-production pathways (Schoell, 1980; Whiticar et al., 1986). The hydrogen in CH₄ comes solely from surrounding water if the CH₄ has been produced by CO₂ reduction, whereas three-fourths of the hydrogen in CH₄ comes from organic matter and only one-fourth from water if the CH₄ is produced by acetate fermentation. Thus, in natural systems, CH₄ produced via acetate fermentation characteristically has D values more negative than -250 SMOW), with typical values between -355 and -290, whereas CH₄ produced from CO₂ reduction is characterized by D values more positive than -250 with typical values near -191 ± 19 (Whiticar et al., 1986).

Organic Matter Base Line

In marine sediments, the source for CH₄ and CO₂ carbon is typically sedimentary organic matter. Thus, to understand variations in the isotopic composition of these mobile carbon phases in a sedimentary sequence, it is necessary to establish the elemental and isotopic composition of the organic matter and to document any significant stratigraphic changes in the parental organic matter. Although only a small fraction of the original organic matter in an ancient sediment may survive, the isotopic composition of the remaining organic matter in the host sediments is believed to faithfully reflect the original organic matter source (Meyers, 1994).

Characteristically, the ¹³C and ¹⁵N composition of bulk sedimentary organic matter depends on whether the carbon and nitrogen came from the atmosphere or seawater, and on the specific biochemical pathways. Tropospheric CO₂ carbon is ~7 (Wahlen et al., 1993), whereas marine bicarbonate is ~0 (Anderson and Arthur, 1983). The carbon fractionation associated with the most common photosynthetic pathway (C3 pathway) is ~-20 (Deines, 1980). Terrestrial organic matter typically has ¹³C values more negative than -24, whereas marine organic matter typically has values more positive than -23 (Meyers, 1994). Similarly, ¹⁵N values of bulk sedimentary organic matter are also indicative of the source: values more positive than 7 indicate marine sources, and values more negative than 3 are characteristic of terrestrial sources (Cifuentes et al., 1988).

The C:N value of organic matter also reflects its source. Typically marine algae have atomic C:N values between 4 and 10, whereas organic matter produced by terrestrial plants (C3 and C4 pathways), typically have C:N values that exceed 20 (Meyers, 1994).

Previous Work on Blake Ridge Methane Sources

Deep Sea Drilling Project (DSDP) Sites 102, 103, 104, and 533 were also drilled on the Blake Ridge (Shipboard Scientific Party, 1972; Gradstein and Sheridan, 1983). Isotopic measurements of the CO₂ gas, DIC, CH₄, and organic carbon were reported in unprecedented detail from Site 533 to the depth of 399 mbsf (Brooks et al., 1983; Claypool and Threlkeld, 1983; Galimov and Kvenvolden, 1983). ODP Leg 164 extends the isotopic stratigraphies initially outlined at Site 533 to greater depths (~750 mbsf) and to other sites along the Blake Ridge.