Calcium carbonate and organic carbon concentrations were measured on sediment samples from Hole 1075A (Table 11). Organic matter atomic carbon/nitrogen (C/N) ratios and Rock-Eval pyrolysis analyses were employed to determine the type of organic matter contained within the sediments. High headspace and core void gas contents were encountered, and routine monitoring of the sedimentary gases was done for drilling safety.
Concentrations of carbonate carbon are low in Site 1075 sediments. They vary between 2.1 and <0.1 wt% (Table 11). The maximum carbonate carbon concentration is equivalent to 17 wt% sedimentary CaCO3, and most sediment samples contain less than a few weight percent CaCO3. These generally low concentrations agree with the paucity of coccoliths and the high abundances of opaline material and continental clastic sediments at this site (see "Biostratigraphy and Sedimentation Rates" section, this chapter). The range in concentrations reflects a varying combination of changes in biological production of calcareous material, dilution by noncalcareous components, and carbonate dissolution fueled by oxidation of or-ganic matter.
TOC determinations were done on a smaller number of Hole 1075A sediment samples than carbonate determinations because of time constraints. TOC values range from 4.42 to 0.96 wt% (Table 11) and average 2.60 wt%. The concentrations are ~10 times greater than the average of 0.3 wt% given by McIver (1975), based on Deep Sea Drilling Project (DSDP) Legs 1-33, a value that can be considered representative of typical deep-sea sediments. The high TOC concentrations at this site may be ascribed to a combination of high supply from elevated paleoproductivities and a high accumulation rate of sediments enhancing organic matter preservation.
Organic C/N ratios were calculated for Site 1075 samples using TOC and total nitrogen concentrations to help identify the origin of their organic matter. Site 1075 C/N ratios vary from 21.2 to 6.6 (Table 11). Lower C/N ratios occur in samples that contain less organic carbon; these values may be biased by the tendency of clay minerals to absorb ammonium ions generated during the degradation of organic matter (Müller, 1977; see "Inorganic Geochemistry" section, this chapter).
The C/N ratios average 12.6, a value that is intermediate between unaltered algal organic matter (5-8) and fresh land-plant material (25-35; e.g., Emerson and Hedges, 1988; Meyers, 1994). These organic carbon-rich sediments probably contain a mixture of partially degraded algal material and detrital continental organic matter. Preferential loss of nitrogen-rich, proteinaceous matter can elevate the C/N ratios of algal organic matter during settling to the seafloor (Meyers, 1997).
A Van Krevelen-type plot of the hydrogen index (HI) and oxygen index (OI) values indicates that the sediments contain a mixture of type II (algal) and type III (land-derived) organic matter (Fig. 31). This source assignment for the organic matter is consistent with the intermediate C/N ratios for these samples, which also suggest that the organic matter is a mixture of marine and continental material. A more likely possibility, however, is that the sediments principally contain algal-derived organic matter that has been altered by microbial processing during early diagenesis. Well-preserved type II organic matter has high HI values (Peters, 1986), which can be lowered by microbial oxidation (Meyers, 1997). The low HI values of fresh type III organic matter, however, cannot become elevated by postdepositional alteration. In general, Hole 1075A sediments having higher Rock-Eval TOC values also have higher HI values (Table 12). This relationship confirms that the algal organic matter has been oxidized. Variable Tmax values reflect poorly defined S2 peaks and not actual thermal maturities of organic matter. Those samples in which the geometry of S2 peaks was sharp have relatively low Tmax values (Table 12), showing that organic matter is thermally immature with respect to petroleum generation (Peters, 1986) and therefore contains little detrital organic matter derived from erosion of ancient sediments and transported to this site by the Congo River.
Sediments from Hole 1075A were very gaseous. Gas pressures became great enough in sediments below Core 175-1075A-15H (125 mbsf) to require perforating the core liner to relieve the pressure and alleviate core expansion. Natural gas analyses determined that much of this gas was CO2 (Table 13). Hydrogen sulfide could be detected by nose, but not by hydrogen sulfide-sensing instruments having a sensitivity of ~1 ppm, in Cores 175-1075A-3H through 6H (12.5-48.5 mbsf).
Methane (C1) first appears in headspace gas samples in Hole 1075A sediments at 12.5 mbsf. Concentrations gradually increase and become significant in sediments below 30 mbsf (Fig. 32). Two sources of the gas are possible. First, gas from some deeper origin may have migrated into the unit. Evidence for migration of methane into porous sediments from deeper sources has been found at Sites 762 and 763 on the Exmouth Plateau where a thermogenic source exists in underlying Jurassic rocks (Meyers and Snowdon, 1993). Mesozoic sources of thermogenic gas are known on the southwest African margin from ongoing exploration. A second possible source is in situ formation by methanogenic microorganisms. High methane/ethane (C1/C2) ratios and the absence of major contributions of higher molecular weight hydrocarbon gases (Table 13) indicate that the gas is biogenic, as opposed to thermogenic, in origin. The origin of the methane is probably from in situ microbial fermentation of the marine organic matter present in the sediments. Similar microbial production of methane from marine organic matter has been inferred from high biogenic gas concentrations in Pliocene-Pleistocene sediments from DSDP Site 532 on the Walvis Ridge (Meyers and Brassell, 1985), Sites 897 and 898 on the Iberian Abyssal Plain (Meyers and Shaw, 1996), and also in middle Miocene sediments from Site 767 in the Celebes Sea (Shipboard Scientific Party, 1990). A biogenic origin of the methane is supported by the disappearance of interstitial sulfate at approximately the same sub-bottom depth where methane concentrations begin to rise (see "Inorganic Geochemistry" section, this chapter). As noted by Claypool and Kvenvolden (1983), the presence of interstitial sulfate inhibits methanogenesis in marine sediments.