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Figure 1: Region of coastal region along the Texas-Louisiana Shelf in the Gulf of Mexico that has a large abundance of surface outcropping of methane hydrate.

These BSR (bottom simulating reflectors) features are prevalent at the base of continental margins at sub-bottom water depths of about 300 m. The problem with estimating the quantity of methane stored in these reservoirs is in determining how much methane—free or in hydrates-will produce the detected change in sound velocity. In other words, some sort of calibration needs to be developed between gas quantity and seismic signal.

The primary focus of global ocean modeling is surface ocean primary production. Recently the models have been enhanced with the addition of coastal inputs. However, while surface ocean carbon cycles are their primary focus there has been realization over the last 15 years that there are numerous active regions on the ocean floor, which are more biologically active than the surface ocean. The enhanced biological activity at these sites is driven by a number of processes including geothermal energy at vents, and reduced compounds seeping from cold seeps and mud volcanoes. Methane hydrates are commonly observed at these sites.

Modeling and predicting global climate is, in part, a function of modeling carbon cycling in the ocean. It may be that the contribution of the significant biological activity on the ocean floor should be incorporated into such models, allowing for greater heterotrophic activity in the water column fueled by substrates from the seafloor (Kelley et al., 1998). Certainly, the abundance of methane hydrates in coastal oceans at relatively shallow depths could significantly impact atmospheric composition and global warming. Warming of ocean waters by 0.5 to 1.0°C has the potential to result in an enormous flux of methane to the atmosphere, which would result in further global warming (Macdonald, 1990; Paull et al., 1991).

The content of this chapter will provide an overview of sedimentary processes that influence the carbon cycle within the sediment, at the sediment-water interface, and in the water column (Fig. 2). The presentation will concentrate on the microbial role in the geochemical cycles and will describe the use of carbon isotope data to assist in interpreting the cycles. Isotopic data on carbonate hard-grounds and organic rich sediments provide a record of geochemical pathways and metabolic processes that are active in the formation, stability and fate of methane hydrates.


There has been a broad effort in the development of elemental isotopic analyses. Stable isotope ratios have been used to examine carbon, nitrogen, and sulfur cycling to determine the major sources of organic matter that support food webs in a variety of ecosystems (Peterson and Fry 1987; Fry, 1986; Fogel et al., 1989). Further development of this technology has provided the ability to pursue cycling of carbon at the molecular level (Coffin et al., 1990; Silfer et al., 1991; Meier-Augenstein, 1995; Hullar et al., 1996) allowing identification of specific microbial roles in the biogeochemical cycling of carbon and nitrogen. In most ecosystems, a complex mixture of physical, chemical and biological factors control the carbon sources that are available to support bacterial production. Isotope analysis provides delineation of bacterial assimilation of substrates in complex ecosystems, and has provided a more thorough understanding of microbial processing (Coffin et al., 1989; Coffin et al., 1994; Hullar et al., 1996). Such methods have been developed to identify the substrate sources that support bacterial production in aquatic environments (Coffin et al., 1989; 1990).

Figure 2: This manuscript outlines the contribution of biogenic and thermogenic methane to methane hydrates, and the influence of this energy on biological cycles at the sediment-water column interface and the ocean water column.

These techniques have been combined with 613C analysis of dissolved and particulate organic carbon (DOC and POC), dissolved inorganic carbon (DIC) and methane to examine the roles of bacterioplankton in aquatic carbon cycles (Peterson et al., 1994; Coffin et al., 1994, 1997; Coffin and Cifuentes, 1999). The development of these new approaches in isotope biogeochemistry has provided the ability to determine carbon sources that support bacterial production from a variety of anthropogenic, autochthonous and allochthonous sources.

There is a large data base available on carbon isotope ratios for the components of carbon cycles in the ocean. For example, there is a long history of S13C measurements of oceanic plankton provide values that vary between -18%o and -30%o in the surface ocean (Gearing et al., 1977; Sackett, 1991). The variation of the stable carbon isotope signature of this pool has been found to be a function of the water temperature and CO2 (aq) concentration and the phytoplankton type (Rau et al., 1989; 1991). As a result, in polar regions algal biomass is substantially more depleted in 13C with values ranging from -23.2%o at 53.3° S and a decreasing to -30.3%o at 62°S in a transect across Drakes Passage from the coast of Chile to the Northern reaches of Antarctica (Rau et al., 1991).

In addition, carbon isotope analysis is used to delineate ocean regions where carbon cycling is controlled by terrestrial input, phytoplankton production, autotrophic fixation of carbon during chemosynthetic oxidation of reduced compounds, and assimilation of thermogenic carbon sources (Peterson and Fry, 1987; Rau et al., 1989; Sassen and MacDonald, 1997). An example key to this chapter is methane that is produced from thermal alteration of organic matter is relatively 13C enriched (varying for -39 to -45%o, while methane of biogenic origin is 13C depleted, varying from -55 to -110%o (Martens et al., 1992).

An additional development in this area has been the use of molecular assays of specific bacterial markers to determine the contribution to complex cycling. Compounds such as d-alanine provide the ability to trace carbon isotope signatures through the eubacterial component of the microbial web (Pelz et al. 1997; 1998). Further delineation of the role of specific bacterial groups in carbon cycling can be obtained with analysis of phosolipids, that select key components of the archeabacterial and eubacterial community (White et al., 1979; Zelles et al., 1992). In an additional step, with the analysis of specfic biomarkers such as saturated and unsaturated isoprenoids, bacterial and archaea roles in methane cycling is determined (Elvert et al., 1999; Thiel et al., 1999). This approach provides a thorough evaluation of bacterial roles in carbon cycling.

In addition to stable isotopes and molecular tracers, radiocarbon (S14C), is particularly useful in constraining the complex carbon cycle of the ocean. Radiocarbon is formed in the atmosphere when 14N is altered by cosmic rays. The quantity of 14C in organic material is determined from the 14C content of substrates which have been input to the material and the time since the isolation of the material from the production cycle, since radiocarbon decays with half life of 5370 years. In the ocean, A14C ranges between approximately +100%o and an undetectable value, near -1000%o (Cherrier et al. 1999). The A14C of dissolved organic carbon in the ocean varies widely and is related to thermohaline ocean circulation, among other factors. Open ocean surface waters range between -150%o to -258%o (Williams and Druffel, 1987; Bauer et al. 1992, 1998; Druffel et al., 1992). Deeper ocean water, below 1000m characteristically have A14C values that range from -393%o to -525%o (Williams and Druffel, 1987). A somewhat different range occurs for dissolved C02, as reported in results of the World Ocean Circulation Experiment (WOCE) for the Pacific Ocean (von Reden et al 1997). In that work, the surface ocean waters were reported to have a A14C range from about -50 to +150%o, while deeper waters below 1000 m have a range from about -160 to -220%o.

Radiocarbon isotope analysis is especially useful in tracing carbon cycling at the interface of petroleum, geothermal or seep carbon sources which contain little to no radioactive carbon (Brooks et al., 1987, Paull et al., 1989, Bauer et al., 1990), and contemporary carbon sources fixed from atmospheric CO2, which contain abundant radiocarbon. Analyses of A14C

have also been employed to study the cycling of aged and recent carbon sources in river, estuarine and wetland aquatic systems (Schiff et al., 1990; Chanton et al., 1995, Chasar et al., in press). Recently, methods have been developed to use A14C to differentiate between bacterial assimilation of new relative to aged carbon (Bauer et al., 1990, 1992, 1995; Cherrier et al., 1999). Cherrier et al. (1999) extended methods developed to study &13C of bacteria (Coffin et al. 1990) to radiocarbon analysis. A result of this effort was the finding that approximately 20% the carbon assimilated by bacteria in the surface waters in the mid-Pacific Ocean was older, presumably recalcitrant material. This result initiates a new understanding of ocean carbon cycling. In addition to A14C of bacteria, current methods include analysis of A14C for DIC, POC and DOC (Bauer et al. 1990, 1992, 1995; Chasar et al., in press).

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