Although apparently contradictory, the co-occurrence of dissolution and precipitation processes was identified in some samples (e.g., #2, Figure 6). Additionally, gypsum crystals were identified in several samples, as well as inorganic calcite (Figure 2). Such processes may happen either during storage of sediments after core recovery, or during early diagenesis. All biogeochemical reactions mentioned in the following sections are numbered and listed in Table 3.
5.3.1. Transformations Occurring During Sediment Storage
 Laminated sediments from the eastern Pacific OMZ are known to contain large amounts of reduced sulfur species such as hydrogen sulfide and solid-phase iron sulfides (e.g., mackinawite FeS, pyrite FeS2) [Reimers et al., 1996]. When placed in an oxic environment (i.e., if oxygen diffuses into the sediments after core recovery), these reduced sulfur species may be oxidized stepwise to produce SO42−, elemental S and H+ (equation (8) in Table 3). The organic matter that was preserved in the sediments can also be oxidized to produce CO2 (equation (1)), which is incorporated in the dissolved carbonate pool. These oxidation processes can lead to a local decrease in pore water pH that favors the dissolution of biogenic carbonates, as already observed by Schnitker et al.  and Self-Trail and Seefelt  for sediments freshly recovered and dried in the presence of oxygen. Both studies also reported the formation of gypsum, resulting from the presence of dissolved calcium and sulfate in the pore space. In core MD02-2520, the δ34S values of gypsum range from −10 to +15‰, which is lighter than seawater sulfate (∼20‰ [Rees et al., 1978]). Isotopically light sulfate could originate from the oxidation of reduced sulfide species, which are enriched in 32S due to the preferential uptake of the light isotope during sulfate reduction [Jørgensen, 1990]. Unfortunately, the isotopic composition of dissolved sulfate and sulfide, and solid phase sulfide is not available for core MD02-2520. However, studies of similar sedimentological settings (like the laminated sediments of the Santa Barbara Basin, which have similar accumulation rates and TOC contents) demonstrate that dissolved and solid phase sulfide species tend to have negative δ34S values (∼−20‰ for H2S and FeS2 [Brüchert and Pratt, 1996]). Therefore, the gypsum crystals that were analyzed for the present study may potentially reflect post-sampling sulfide oxidation, with additional contributions from isotopically heavier (residual) seawater sulfate and isotopically lighter organic sulfur (with δ34S values ranging from −5 to −30‰ [Werne et al., 2008]).
 However, post-sampling oxidation of sulfide and organic matter generally results in the dissolution of biogenic carbonate [Schnitker et al., 1980; Self-Trail and Seefelt, 2005], which is a minor process in core MD02-2520 (section 5.2.1). This suggests that the calcite aggregates most probably formed during early diagenesis. Self-Trail and Seefelt  demonstrated that the amount of carbonaceous fossils dissolved increases with drying time. In other words, the longer sediments are put in contact with oxygen, the more biogenic carbonate will be dissolved. This is not observed in our case: the foraminiferal assemblage and quantity did not vary between the two batches processed, nor did the amount of gypsum despite a few months elapsing between the batches (Figure 2 and section 3.1). Therefore, although we cannot rule out post-sampling oxidation, this effect seems to be of minor importance in core MD02-2520.
5.3.2. Transformations Occurring During Early Diagenesis
 In marine sediments, two main processes largely control the carbonate system of pore waters, and therefore influence the dissolution of biogenic carbonates and precipitation of inorganic calcite: i) the mineralization of organic matter and ii) the anaerobic oxidation of methane (AOM) [Jørgensen and Kasten, 2006]. The dissolution of biogenic carbonates is generally induced by the mineralization of organic matter in the presence of oxygen (equation (1)), which leads to a release of CO2 into the pore waters [Jahnke et al., 1997; Pfeifer et al., 2002; Volbers and Henrich, 2002]. In laminated sediments from the eastern Pacific OMZ, the oxic zone is restricted to the upper few cm (sometimes even to the upper few mm) of sediments deposited [Reimers et al., 1996] (Figure 7). Additionally, the precipitation of iron monosulfides produces protons that might lower the pore water pH (equation (6)) and lead to dissolution of biogenic carbonates. However, in sediments containing a large amount of iron oxides (which is the case for core MD02-2520 [cf. Blanchet, 2006]), iron liberation reactions dominate in the suboxic zone and determine the pore water pH (equation (3)), which can reach particularly high values (like in the Santa Barbara Basin, where pH are higher than 8) [Reimers et al., 1996]. Therefore, we consider that carbonate dissolution was restricted to the oxic zone and is a minor diagenetic process at the site of core MD02-2520.
Figure 7. Schematic sediment column for core site MD02-2520, with (right column) geochemical zonation, (center) pore water chemical profiles (central column), with diffusion directions (arrows) and changes in isotopic composition, and (left column) transformations occurring within the solid phase.
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 Precipitation of inorganic calcite is tightly linked to pore water alkalinity, which generally increases in organic-rich marine sediments as a result of anaerobic organic matter mineralization and AOM (equations (2)–(5)). When oxygen is completely depleted in the pore waters, other oxidants such as nitrate (equation (2)), iron and manganese oxides (equation (3)) or sulfate (equation (4)) are used to mineralize the organic matter, which releases carbonate ions into the pore water and increases alkalinity [Reimers et al., 1996]. Such reactions might induce authigenic calcite precipitation in the suboxic and sulfidic zones (Figure 7). In the suboxic zone, the pore water δ13CDIC is largely determined by the δ13C signature of the organic matter (∼−20‰) [Stott et al., 2002]. Around the SMT, the process of AOM results in an increase in pore water alkalinity and a decrease in δ13CDIC, due to contribution of methane-derived carbonate ions with δ13CDIC sometimes reaching −50‰ [Torres et al., 1996]. Precipitation of inorganic calcite around the SMT has been well-documented [e.g., Bohrmann et al., 1998; Greinert et al., 2002; Nöthen and Kasten, 2011], but such calcite generally has lighter δ13C signatures than the aggregates from core MD02-2520 (e.g., ∼−15‰ in organic-rich sediments from the Chilean upwelling region [Treude et al., 2005]). Authigenic calcite formed in methane seeps is also known to overprint the fossil foraminiferal record (due to precipitation of authigenic calcite on the tests) but it also carries δ13C values much lower than those obtained for core MD02-2520 (−22 to −51‰ [Torres et al., 2003; Cook et al., 2011]). Unfortunately, pore water δ13CDIC profiles are not available for our study site, thus it is impossible to precisely determine the formation depth of the observed inorganic calcite. We hypothesize that it formed in the suboxic zone and at the top of the sulfidic zone, mostly as a result of anaerobic oxidation of organic matter with a minor contribution from carbonates produced during AOM (Figure 7).
 The formation of gypsum crystals in marine sediments has been reported and described in only a few studies [Siesser and Rogers, 1976; Briskin and Schreiber, 1978; Muza and Sherwood, 1983; Hoareau et al., 2011]. In order to characterize the sedimentary context where gypsum might precipitate, Hoareau et al.  calculated the gypsum saturation index for numerous ODP/IODP cores using the available pore water data. They reported that gypsum saturation was observed in the presence of evaporitic horizons or volcanogenic material (able to provide sufficient calcium and sulfate ions to the pore water). The sediments of core MD02-2520 are barren of evaporitic or volcanogenic material, and pore water data are unfortunately not available. It is therefore impossible to precisely pin-point the mechanism leading to gypsum formation. However, in the following we will review the potential processes that could lead to local supersaturation of pore waters with respect to gypsum. Such processes can involve either: i) oxidation of the reduced sulfide species that could provide sufficient sulfate ions to form gypsum, or ii) an increase of pore water salinity that would be sufficient to reach gypsum supersaturation.
 Sulfide oxidation in anoxic marine sediments is an important component of the biogeochemical sulfur cycle, and can significantly contribute to the dissolved pore water sulfate pool [Jørgensen and Kasten, 2006; Riedinger et al., 2010]. Fossing and Jorgensen  have for instance observed a complete re-oxidation of radio-labeled monosulfides (FeS) and elemental sulfur (S0) to sulfate in anoxic sediments. The complete re-oxidation of sulfur species (of various oxidation states) to sulfate generally involves a suite of chemical and microbially mediated chain reactions. The H2S produced during either organoclastic sulfate reduction or AOM might either: i) react with solid-phase Fe(III) oxides, dissolved (Fe2+) iron or iron monosulfides to form iron sulfide minerals of various oxidation states (FeS or FeS2, equations (6) and (7)), or ii) be re-oxidized by oxidants such as iron and manganese oxides, which gives rise to sulfide species of intermediate oxidation state such as elemental sulfur (S0) (equations (9)–(11)) [Thamdrup et al., 1994]. Schippers and Jørgensen  have shown that manganese dioxide is a powerful chemical oxidant, that is able to oxidize monosulfides to elemental sulfur (equations (10) and (11)). Oxidation of sulfur species coupled to the reduction of iron (oxyhydro)oxides has also been demonstrated to occur in sediments of the Nankai Trough [Riedinger et al., 2010]. When accompanied by microbial processes such as elemental sulfur disproportionation, monosulfides as well as pyrite can be reoxidized to sulfate (equations (13) and (14)) [Aller and Rude, 1988; Schippers and Jørgensen, 2001]. Sulfur disproportionation is a type of inorganic fermentation that occurs in anoxic sediments rich in intermediate sulfur species such as elemental sulfur or thiosulfate (S2O32−). In this way, a quarter of the elemental sulfur is oxidized to sulfate while the remaining three quarters return to the sulfide pool (equation (12)) [Jørgensen and Kasten, 2006]. In core MD02-2520, the glacial sediments are enriched in iron (oxyhydr)oxides [Blanchet, 2006], which were recently shown to favor oxidative sulfur processes [Riedinger et al., 2010; Holmkvist et al., 2011a, 2011b]. Such processes could therefore lead to the accumulation of sulfate in micro-environments in the pore water and could explain the co-existence of pyrite and gypsum in the glacial sediments of core MD02-2520.
 Finally, we speculate that local increases in pore water salinity, sufficient enough to reach gypsum saturation state, might also occur as a result of gas hydrate formation. During gas hydrate formation, ions are excluded from the hydrate structure and lead to the formation of residual brines either in the sediments surrounding the gas hydrates or in the pore space of the hydrates [Ussler and Paull, 1995]. Site MD02-2520 lies within the gas hydrate stability field and the presence of in situ gas hydrates was evidenced by sediment expansions during core recovery (i.e., section 3.1). It is therefore possible that oversaturation and thus precipitation of gypsum (and possibly also authigenic calcite) has been induced in situ during formation of gas hydrates below the SMT where iron sulfides are known to precipitate [e.g., Kasten et al., 1998; Jørgensen et al., 2004].