RESULTS AND DISCUSSION

Lithologic Units, Pore Waters, Sulfate Reduction, and Anaerobic Methane Oxidation

The depth profiles of the contents of different sedimentary carbon, sulfur, and iron fractions measured at all sites recovered during Leg 207 are presented in Figure F4. The geochemical data demonstrate the occurrence of highly OM rich (Cretaceous) black shales (Unit IV) with TOC contents as high as 18 wt%. The black shales were deposited on synrift clastic sediments (Unit V). Units IV and V are overlain by organic-poor open-marine chalk and calcareous claystones (Units I–III) (Erbacher, Mosher, Malone, et al., 2004). Downcore variations of the pore water sulfate (Fig. F2) clearly indicate that deep-seated microbial sulfate reduction at slow rates is occurring in the sediments above Unit IV. Anaerobic oxidation of upward diffusing methane (AOM), which is derived from the black shale sequence, is the process associated with the microbial sulfate reduction process (Erbacher, Mosher, Malone, et al., 2004). AOM has been found to be carried out in marine sediments by a consortium of archaea and sulfate-reducing bacteria (Hoehler et al., 1994; Hinrichs et al., 1999; Boetius et al., 2000). The mechanistic, qualitative interpretations from the pore water profiles are additionally confirmed by quantitative modeling (Arndt et al., 2006). The reaction zone is currently positioned at Sites 1257–1260 above the black shale sequences and is triggered by a flux of biogenic methane from the OM-rich shales, where it is produced by methanogenesis. At Site 1261, sulfate reduction already goes to completion within the upper 200 mcd, indicating that OM degradation takes place in the rapidly deposited Pliocene nannofossil clay of Unit I (Erbacher, Mosher, Malone, et al., 2004). Whereas OM mineralization at depth is reflected by the continuous increase in ammonium concentrations, alkalinity data are partly superimposed by carbonate precipitation. Accumulation of dissolved barium concentrations in the pore waters (Fig. F2) originates from the dissolution of biogenic barites and only takes place where pore water sulfate was completely exhausted. Enhanced dissolved iron concentrations, on the other hand, exclude significant sulfide concentrations.

Sedimentary Iron and Sulfur Species

The burial of OM in the black shale sequence is associated with an enrichment of all analyzed sulfur fractions: total, pyrite, and organic-bound sulfur (Fig. F4). This is due to the coupling of OM deposition to microbial sulfate reduction and the associated formation of sedimentary sulfur compounds. Dissimilatory sulfate reduction leads to the formation of hydrogen sulfide that may further react with reactive iron to precipitate iron sulfides (essentially pyrite, FeS2) and with OM to form organic sulfur compounds (e.g., Aizenshtat et al., 1983, 1995; Sinninghe Damsté and De Leeuw, 1990; Bein et al., 1990; Rullkötter, 2000; Werne et al., 2004). Other metal sulfides (e.g., ZnS) that may have been formed in a sulfidic paleowater column (Brumsack, 1980), although found in investigated black shale samples (A. Hetzel, unpubl. data), are quantitatively only of trace importance.

Reactive Iron Phases and Pyrite Formation

Pyrite typically occurs in marine sediments in framboidal and euhedral occurrence, depending on the physicochemical boundary conditions (Wilkin et al., 1996; Wang and Morse, 1996). As shown in Figure F5, framboidal pyrite was found in black shale samples from Site 1260. This occurrence is typical for pyrite that is formed in a euxinic water column or in sediment close to the sediment/water interface during early diagenesis (Wilkin et al., 1996). Besides sulfur, the iron contents are also enhanced in sediments of Unit IV (Fig. F4), in particular, the FeP and FeD fractions. The only exception is Site 1261. The formation of pyrite is ultimately limited by the availability of iron minerals that are able to react with dissolved sulfide (Canfield, 1989). The amount of so-called highly reactive iron (FeHR) in marine sediments consists of the sum of the iron fraction that already reacted to pyrite (FeP) and sedimentary iron that is still able to react with sulfide. This still-reactive iron fraction (FeD) is extracted with buffered Na dithionite solution (Canfield, 1989). The relationship between FeHR and FeT has been shown to be indicative for the redox conditions characterizing the sediment-forming environment, with FeHR/FeT ratios < 0.38 in normal marine environments with oxic bottom waters (Raiswell and Canfield, 1998; Anderson and Raiswell, 2004). In euxinic systems, on the other hand, the clastic and reactive iron fluxes to the sediment may be decoupled, which may lead to FeHR/FeT ratios > 0.38. This corresponds to an excess of reactive sedimentary sulfur in euxinic compared to oxic sediments. Below oxic bottom waters, pyrite formation takes place exclusively in the sediment. An additional fraction of iron sulfide may be formed in the water column of euxinic systems, as found in the modern Black Sea (Raiswell and Berner, 1985; Canfield et al., 1996). It has been shown for modern environments that the geochemical indicators in the iron-sulfur system, such as the FeHR/FeT ratio and the degree of pyritization, mostly lead to the same paleoredox interpretations (e.g., Shen et al., 2003).

The downcore variations of the relative fraction of highly reactive iron (FeHR/FeT) for the Leg 207 samples are presented in Figure F4. Reactive iron was enriched in virtually all black shale samples with FeHR/FeT values > 0.38, indicating euxinic conditions during deposition of the organic-rich sediments of Unit IV (Fig. F4). Differences between the sites, as well as downcore variations of the relative enrichment of reactive iron, indicate that environmental conditions and/or associated transport processes were not constant with time. Besides water column iron sulfide formation, the enrichment of reactive iron also requires the presence of a paleoshelf situation where an extended oxygen minimum zone (OMZ) led to the liberation of dissolved iron from shelf areas into a suboxic water column and, after further transport, to precipitation when reaching the sulfidic deeper waters (Canfield et al., 1996; Lyons, 1997; Wijsman et al., 2001; Anderson and Raiswell, 2004). Alternatively, such an enrichment may have been caused by a fluctuation of a chemocline on the shelf slope, leading to a pumping of dissolved iron into suboxic waters with subsequent fixation in areas of higher sulfide accumulation. The latter mechanism is similar to a model proposed by Lepland and Stevens (1998) for the formation of Mn(II) carbonates in an anoxic deep of the Baltic Sea. Additionally, an excess of FeHR is also found below the black shales in Unit V at Sites 1257, 1258, and 1260. In contrast, euxinic conditions were limited to Unit IV at Site 1259, based on the present sampling resolution. Besides an onset of euxinic conditions already occurring during clastic sediment deposition of Unit V, a later sulfidization of underlying sediments as described for sediments below sapropels of the Kau Basin (Middelburg, 1991), the Eastern Mediterranean (Passier and de Lange, 1998; Passier et al., 1996, 1997, 1999), the Baltic Sea (Böttcher and Lepland, 2000), or the Black Sea (Jørgensen et al., 2004) may also have caused this iron sulfide enrichment. A similar diagenetic sulfidization mechanism was identified in Mesoproterozoic marine sediments of the Belt Supergroup (Lyons et al., 2000).

FeHR/FeT ratios in Units I–III are below the threshold value for euxinia, indicating an essentially nonsulfidic water column during sediment formation. When the different sites are compared, however, it becomes obvious that FeHR is relatively enriched in the sediments at Site 1261, which is positioned closest to the paleocoastline and at the lowest paleowater depth. An import of iron from shallower shelf sediments according to the (suboxic) OMZ model is consistent with this observation.

The covariation of SP with TOC data is presented in a Berner plot (Fig. F6) and compared to the relationship proposed for "normal marine sediments" (Berner and Raiswell, 1983). Only a few data points coincide with the relation found for clastic sediments below an oxic water column. A number of data points plot above the regression line, indicating an excess of sulfur that may coincide with a euxinic depositional environment (Leventhal, 1983; Raiswell and Berner, 1985). At highest TOC contents, however, most data show a relative excess of OM. This indicates a certain degree of iron limitation during black shale deposition (Leventhal, 1983) as also found for the Black Sea and Mediterranean sapropels (Fig. F6) or Albian black shales from the North Atlantic (Hofmann et al., 2000). Iron limitation upon black shale formation is also indicated from the sedimentary sulfur speciation (Fig. F4). SP in the investigated black shale samples makes up between 30% and 100% of ST, with a decrease of the relative importance with increasing OM content (Fig. F7). This indicates the importance of the balance between OM and the syngenetic metal flux to the surface sediments. In addition to fixation of sulfide by the reaction with iron, OM acted as the second important sulfur trap during early diagenesis.

Organic Sulfur Formation

Sulfur can react with OM via a number of different pathways, where sulfide and polysulfides are the most likely reaction partners (Aizenshtat et al., 1983, 1995). This requires a decreased availability of reactive iron and leads to a modification of reactivity of the remaining OM (Sinninghe Damsté and De Leeuw, 1990). Organic sulfur incorporation is found in the high-TOC black shale samples at all sites (Fig. F4). From the nearly linear variation of organic sulfur and organic carbon contents (Fig. F6), essentially constant atomic C/S ratios are obtained. Quantitatively, the samples with TOC contents exceeding ~2 wt% have as much as 10 atom% organic sulfur. Most of the atomic S/C ratios fall in the range of 0.04 to 0.06. A similar linear relationship has been observed previously for Mediterranean sapropels by Passier et al. (1999). The fraction of SORG, however, is relatively more enriched in the Cretaceous black shale samples, probably due to a higher abundance of reactive sulfur species or a higher reactivity of the OM toward sulfurization. From a comparison with literature data it is obvious that OM in the black shales is significantly enriched in sulfur when compared to marine planktonic material (S/C of ~0.008) (Francois, 1987). This is due to the reaction of reduced sulfur species with OM upon early diagenesis (Aizenshtat et al., 1983, 1995; Bein et al., 1990; Raiswell et al., 1993; Sinninghe Damsté and de Leeuw, 1990; Passier et al., 1999; Werne et al., 2004). Stable sulfur isotope measurements have shown that the original seawater-derived sulfur in the OM was superimposed by the addition of diagenetic sulfur species (Bein et al., 1990; Passier et al., 1999). The relative importance of OM sulfurization compared to the bonding to pyrite increases with OM contents (Fig. F7). Deviations of atomic S/C ratios from the mean value of 0.056 (TOC > 2 wt%) may be caused by different extents of dissolved sulfur species availabilities and/or different sulfur sink capacities of OM. In Creaceous carbonates, Bein et al. (1990) observed maximum S/C ratios as high as 0.38. On the other hand, Jurassic black shales, anoxic Peru margin upwelling sediments, and Mediterranean sapropels had maximum S/C ratios of 0.019, 0.056, and 0.038, respectively (Raiswell et al., 1993; Mossmann et al., 1991; Passier et al., 1999).

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