The Leg 164 sediments have TOC values that are typical of ocean margins (1%-2%) and are enriched in organic matter, compared to deep-sea sediments. The reasons might be the additional transport of organic matter from the nearshore environments to the Blake Ridge and enhanced preservation caused by rapid burial. Sedimentation rates increase from about 70 m/m.y. (106 yr) within the Quaternary to about 300 m/m.y. within the late Miocene at Site 994 (Paull, Matsumoto, Wallace, et al., 1996). The TOC content follows this increase, at least in the upper part of the profile, down to about 200 mbsf (Fig. 1).
The mass accumulation rate of TOC, which takes into account the porosity and the sedimentation rate of the sediments (van Andel et al., 1975, Stein et al., 1989), increases 25-fold downcore (Fig. 8A) from < 0.02 g cm-2 k.y.-1 (103 yr) in latest Pleistocene to 0.5 g cm-2 k.y.-1 during the Miocene for sediments from Hole 994C (Paull, Matsumoto, Wallace, et al., 1996). This increase is paralleled by an increase in the vitrinite concentration (Fig. 8B) and by a slight increase of the odd-even predominance of the n-C27H56 (Fig. 8C). We believe therefore that the prevalence of the odd-numbered n-alkanes is caused by the presence of terrestrial organic matter and not by marine microalgae (Zegouagh et al., 1998). Both parameters suggest that terrestrial organic matter occurs more in the older sections. The microscopic investigation of the organic matter, however, reveals a practically constant composition of the particulate organic matter along the profile. A more detailed inspection of the data (Table 3) shows that the organic particles of the uppermost sample from Hole 994C consists of liptinite (85%) and inertinite (15%). The liptinite of this sample is made up of 80% spores (terrigenous) and 20% algae (marine). Therefore, altogether, 17% of the organic particles is of marine and 83% of terrestrial origin. A comparable overall composition was found for the structured organic matter of the lowermost sample from Hole 994C. However, in this case the organic particles consist of only 25% liptinite, which is made up of algae to 95%. Therefore, the marine part of the organic particles amounts to 24% and the terrestrial part to 76%. Additionally, around 20% of the terrigenous organic matter consists of vitrinite, which is not present in the uppermost sample. Obviously, the terms "terrestrial" or "terrigenous," often taken as a synonym for type III kerogen, should be used with caution. It should be kept in mind that sporinite, a liptinite classified as type II kerogen, is of terrestrial origin.
The liptinites (pollen and algae) were less degraded than the inertinites and vitrinites, perhaps because of less transport. It is well established that during transport, deposition, and early diagenesis of organic matter the more stable terrestrially derived compounds will be preserved preferentially (Cornford, 1979). Therefore, the organic matter that has survived and can be analyzed today is only a fraction of the original input, as far as amount and composition is concerned. The observed increase of terrestrially derived organic matter with depth might be due to the preferential removal of more easily degradable organic matter of the liptinite group. Therefore, the terrestrially derived organic matter that prevails nowadays in the late Miocene/early Pliocene sediments might be a concentrated residuum. Similar results were achieved by Rullkötter et al. (1986) when looking at Pliocene samples from Hole 533A at the Blake Outer Ridge.
The extractable organic matter of the sediments still contains unsaturated, labile compounds (e.g., hopenes). These compounds are mostly produced through the microbial reworking of the organic matter and do not represent the organic matter input. The increase of the relative concentrations of the microbially produced hopenes and ß,ß-hopanes (Brassell et al., 1981; Ourisson et al., 1984) downhole also may reflect the longer time span available for microbial activity. This may also be true for the increase of the relative isoprenoid concentrations, expressed as phytane/n-C18H38 or pristane/phytane ratios.
Hopenes have been found previously in ODP sediments, such as in Pliocene samples off the coast of Chile (Kvenvolden, et al., 1995); however, no clear depth trend was recognized. Kvenvolden et al. (1995) has suggested three possible mechanisms for the observation of C32-C35 hopane epimer ratios typical of mature organic matter:
However, pollution by grease used for drilling equipment cannot excluded. Further investigations by compound-specific isotope analyses might help to solve this question.
The hydrogen indices (Espitalié et al., 1985), which range between 100 and 300 mg HC/gC (Paull, Matsumoto, Wallace, et al., 1996), classify the organic matter as a mixture of type II and type III. In detail, samples from Site 997 have somewhat lower HI levels (around 100) than samples from the other two sites, which is reflected by the higher content of inertinite and vitrinite in the Site 997 samples (78%), compared to 67% in Hole 994C samples (Table 3). Selection of organic matter during transport by currents of different intensities on the crest and on the flanks of the ridge might explain this finding. It should be mentioned, however, that only the upper part of the drilled profile of Site 997 has been sampled for Rock-Eval pyrolysis.
There is a discrepancy
between the isotopic signature of the kerogen and the facies type of the organic
matter that needs further discussion and investigation. The isotope ratios of
the kerogens are all in the range of -21
for the Hole 995A samples. This value is often used as an indication for the
input of marine plankton that in numerous cases is in the range of -18
to -21 in
Miocene to Holocene sediments (Arthur et al., 1985). However, recent studies of
the variation of the 
13C
ratio of modern phytoplankton in different oceanic environments suggest a
broader distribution of the 13C
levels from -21
to -8 (Popp
et al., 1997).
Organic petrographic
investigations show the prevalence of inertinite, vitrinite, and sporinite,
which are all of terrestrial origin. There are no indications for the occurrence
of grasses (C4 plants), which could be responsible for relatively
heavy carbon isotope ratios (Fry et al., 1977). Additionally, the alkane
distribution, dominated mostly by n-C29H60 (not
by n-C31H64, which is typical for grasses;
Cranwell, 1973), suggests higher land plant (C3 plants) input,
consistent with the Rock-Eval deduced occurrence of type II/III kerogens.
Therefore, the kerogen isotope ratios should be in the range of -24
to -29,
typical for the C3 photosynthetic cycle.
The results of the organic-petrographic investigations and the carbon isotope analyses, which seem to be contradictory at the first glance, might be explained by a mixture of isotopically light terrestrial and isotopically heavy marine organic matter.
In contrast to the findings of different authors (Vetö et al., 1995; Dean and Arthur, 1989), the hydrogen indices do not correlate with the TOC values. There is also a lack of correlation between TOC and TS content (Berner and Raiswell, 1983). Both observations indicate the degradation of the organic matter that may be used for the sulfate reduction. Another reducing agent besides the sedimentary organic matter could be methane, which migrates in varying amounts from below the sulfate-reduction zone. From the linear shape of the sulfate concentration profile, Borowski et al. (1996) postulated that anaerobic methane oxidation is a major sulfate-consuming process in sediments from the Blake Ridge.
Several authors (Littke et al., 1991; Vetö and Hetényi, 1991; Lallier-Vergčs et al., 1993) attempted to quantify the loss of TOC due to microbial sulfate reduction by using the sulfur content of the sediments. This might be possible in nonbioturbated sediments, because the amounts of H2S that escaped seem to be limited. Stoichiometry shows that 1% reduced sulfur corresponds to a 0.75% loss of TOC (Vetö et al., 1995). However, Blake Ridge sediment bioturbation may result in a substantial loss of H2S. Therefore, original TOC contents would be higher than calculated by the method of Vetö et al., (1995). On the other hand, if methane is the major organic substrate for the sulfate reduction, the loss would be overestimated.
As a rough estimate and assuming that the loss of H2S may compensate the nonconsumption of organic matter in the course of sulfate reduction, the concentration of total sulfur is equivalent to the TOC content that might have been lost. Therefore, the original TOC content of the Blake Ridge sediments might have been around 2.5%, in comparison to the mean content of 1.5% today.
Further loss of TOC occurs below the sulfate-reduction zone because of the methane generation. This loss cannot be quantified, because the amount of gas generated is not known. Headspace gas measurements probably give too low of readings because of the loss of gas during core recovery.
Paull et al. (1994) have calculated that under ideal conditions a transformation of 0.5% TOC would be necessary to fill 6% of the pore space with gas hydrate. The pore space of the Blake Ridge sediments from within the gas hydrate-stability zone may be filled up to 9% with hydrates, deduced from direct methane concentration measurements (Dickens et al., 1996) and vertical seismic profiling (Holbrook et al., 1996). These figures are in the same order of magnitude. Therefore, from the quantitative aspect of the available consumable organic matter, the greater part of the methane now occurring as gas hydrate could be explained as having been generated in situ by microbial activity. The generation might be an ongoing process because viable microbes are present all along the profile (K. Goodman, pers. comm., 1997), as well as metabolizable organic matter (lipids, such as alkanes). However, the carbon isotope ratios of CH4 and CO2 of the free gases suggest their migration from depth (Paull et al., Chap. 7, this volume).
No indications of the hydrate-stability zone were found in the concentration data of ethane and propane (Fig. 6) and in the carbon isotope data (Fig. 7). Only minor changes in methane concentrations were observed close to the stability zone, and they are not considered to properly mark the exact depth position of the stability zone in the holes. However, the ratios C1/(C2+C3) plotted vs. depth in Figure 9 are mostly below 10 to 15, somewhat higher in the uppermost samples and significantly higher just above and below the hydrate-stability zone. The reason for the higher levels close to the stability zone is unknown at present, but a possible explanation might be given: gaseous hydrocarbons, of predominantly microbial origin migrate from depth to the surface. The gas hydrate occurring zone (GHOZ) may act as a cap rock, causing an elevated concentration of methane below the seal. Within the GHOZ, part of the methane is densely packed into the hydrate structure, part is filling the pore space as free gas. Thus, an anomalous concentration of methane, which cannot be measured due to the applied sampling technique described in this paper, has to be assumed (Note: the free methane collected with the PCS [Dickens, et al., 1996] has not been considered in this paper). The low methane concentration now observed within the GHOZ is thought to be an artifact, caused by melting of the gas hydrates during core recovery, which is accompanied by a total destruction of the sediment structure ("soupy texture") and a more efficient loss of methane than from over- and underlying horizons.
Higher methane yields above the GHOZ may be caused by the occurrence of hydrates in such a low concentration that their melting did neither influence the pore-water chemistry, nor destroy the sediment structure, but nevertheless contributed to the content of the combined gases. The elevated methane concentrations between ~20 and ~80 mbsf are caused by the active methane generation below the sulfate-reduction zone, which adds additional methane to the stream of migration from below.
Ethane and propane are
concentrated relative to methane, which is indicative of thermogenic gases (Fig.
9). The thermal origin of these gases is also evident from their carbon
isotope data (Fig. 7), except
perhaps for the uppermost samples. Previous investigations of sedimentary gases
have shown that data of sorbed gases are hardly influenced by degassing
processes (Faber et al., 1990; Whiticar and Suess, 1990; Faber et al., 1997).
Therefore, the carbon isotope data from ethane and propane give information on
the type and maturity of the corresponding source material. In Figure
10 the carbon isotope data from ethane are crossplotted vs. the propane
data. With the exception of the data from the upper parts of the cores, most
data points plot along the isotope maturity line shown in Figure
10. This line was constructed according to the model, published by
Berner and Faber (1996), for ethane and propane derived from an algal kerogen.
It is dependent on the original carbon isotopic composition and the maturity of
the kerogen. However, as the sediments from which the gases have been extracted
are immature, ethane, propane, and the thermal methane are presumed to have
migrated from deeper, more mature parts of the sedimentary column. Assuming the
isotopic composition of the kerogen to be 13C
= -21, as was
determined for kerogens of samples from Hole 995A, then the maturity of the
source rock for ethane and propane (Fig.
10) would be in the range between 0.6% to 1.0% vitrinite reflectance (VR).
This range corresponds to the oil window maturity stage, the beginning of which
was mentioned for sediments located 1600 m sub-bottom depth (Herbin et al.,
1983) at the Blake Bahama Basin close to the survey area.
Mesozoic and Cenozoic formations were drilled during DSDP Legs 44 and 76 within the Blake Bahama Basin. The organic matter from these sediments is mainly of terrestrial origin. However, marine intercalations, equivalents to the Barremian-Cenomanian North Atlantic black shales, have been observed down to 1600 mbsf (Summerhayes and Masran, 1983). The petroleum potential of these horizons, which contain a mixture of detrital and aquatic organic matter, is good to medium, according to Rock-Eval data (Herbin et al.,1983). If similar sediments exist below 1600 mbsf, the higher maturity of the organic matter would be responsible for thermal hydrocarbon generation. Long distance migration, though speculative, may explain the spread of the ethane and propane isotope data.
As mentioned above, the
combined methane is a mixture of thermal and microbial contributions and can be
characterized according to Figure 11.
The data is scattered, as expected for a mixture, but aligns along a mixing line
between data of the potential microbial and thermal partners. Because the bulk
data in Figure 10 is scattered
around 0.8% VR (model: kerogen 13C
= -21), the
isotope value of the corresponding methane can be calculated to 13C1
= -36
(according to the model of Berner and Faber, 1996); the corresponding C1/(C2+C3)
is 4 (Fig. 11). Assuming 13C1
= -64 of the
residual free methane and C1/(C2+C3) = 200 for
the sediment gases, the mixing line (Fig.
11) can be fitted to the data. Most of the methane of the combined
gases, between 30% and 90%, have a microbial origin. However, the results shown
in Figure 11 should not be
overinterpreted, because the combined gases comprise only a small fraction of
the originally present total amount of gases, and the concentration of the
thermal gases (Fig. 6) is lower
by some orders of magnitude than that of the free gases (Paull et al., Chap.
7, this volume). Also, the combined gases may have suffered from
microbial (methane) oxidation and desorption from the sediment grains.
