The S isotope record in pyrite

The downcore distribution of δ34S values in pyrite (CRS; Fig. 1B; Table S1) shows values between −30‰ and +5‰. In the top few metres of the sedimentary sequence, pyrite shows δ34S values around −20‰, and similar values are found in the middle part, within the present-day methanogenic zone. Higher values, reaching 0‰ or more occur near 30 m and 100 m below seafloor (mbsf), and one value near 180 mbsf. The isotope offset of about 70‰ between pyrite and sulphate in porewater of the uppermost few metres (data from Böttcher et al.22) is at the upper limit of fractionation by microbial sulphate reduction commonly observed in marine sediments (e.g. Wortmann et al.23; Jørgensen et al.24; Pellerin et al.20). The offset increases to 80‰ around 100 mbsf but decreases to about 30‰ near the bottom of the sequence, hence showing a strong variation in the apparent fractionation. The highly variable offset between the δ34S signatures of pyrite and sulphate is not consistent with pyrite precipitation from present porewater sulphide, which would expectedly have a more constant offset from porewater sulphate (cf. Jørgensen et al.24). Several factors could have influenced the profile of δ34S in pyrite, such as (1) the depth of pyrite formation, (2) variations of ε34S, (3) changes in microbial activity, or (4) sedimentation rate. Here we discuss which factors are relevant and whether δ34S in pyrite may indeed record overall dissimilatory activity in ancient sediments.

The depth of pyrite formation

Ratios of CRS-related Fe to CRS-Fe + total sequentially extracted Fe (using the five-step extraction scheme of Poulton and Canfield25; see method description below) increase from the sediment water interface to ~0.6 within the top 2 mbsf and remain near this level through the profile (Fig. 2A,B; Table S1). These ratios approximately correspond to the degree of pyritization (DOP; ratio of pyritic Fe to pyritic Fe + acid-soluble Fe; Berner26), indicating that most pyrite forms within the uppermost 2 mbsf. The increase is even higher, to ~0.8, if only fractions I through IV of the extraction scheme are considered for calculation. According to Poulton and Canfield25 fractions I through IV essentially include all Fe-phases but the poorly reactive silicates. As shown by leaching experiments (Kasina et al.27) fraction III includes reactive sheet silicates, such as smectites, from which Fe could potentially react with sulphide.

Figure 2 Abundance of pyrite-bound iron relative to different solid-phase iron fractions in sediment of ODP Site 1229. (A) Degree of pyritization with respect to total Fe (orange triangles), extracted Fe (light-blue squares), and extracted Fe without the poorly reactive silicate fraction (black diamonds) plotted vs. depth. (B) Cross-plot of CRS-bound Fe (pyrite) vs. total Fe, extracted Fe + CRS-Fe, and extracted Fe + CRS-Fe without the poorly reactive silicate fraction. (C) Plot of reactive Fe vs. different Fe-fractions to show extraction efficiency. Lines shown in the plots are regression lines. Full size image

Fe K-edge XANES spectroscopy showed that Fe occurs in significant proportions as pyrite (17–26%), while the rest of the Fe is present in the structure of sheet silicates, mostly from the illite and the chlorite groups (Fig. 3). Only in the uppermost sample at 2.55 mbsf, some Fe was present in smectite. As XANES spectra at the Fe K-edge are similar for different smectites (Fig. 3B; Rennert et al.28), the contribution of the standard spectra we had available (beidellite, saponite, and nontronite) may also represent montmorillonite, which commonly occurs in marine sediments. Overall, the pyritizable Fe-fractions make up a minor amount of the total Fe in the sediment (triangles in Fig. 2C).

Figure 3 Fe K-edge XANES spectroscopy of sediments. (A) XANES spectra of sediment samples collected at ODP Site 1229 at 2.55, 16.49, and 56.35 mbsf. (B) XANES spectra of reference minerals of relevance for this study: sheet silicate minerals of the chlorite group, illite group and smectite group (beidellite, saponite, nontronite), and pyrite. (C) Example of fitting analysis for the sample at 2.55 mbsf that contains Fe in all three sheet silicate groups and in pyrite. Full size image

More abundant is the Fe bound in poorly reactive sheet silicates, which are extracted with boiling HCl (extraction step V). This fraction contributes up to 30% of the total Fe as visualized by the slope of the regression line of the total extracted Fe (fractions I–V) vs. total Fe in Fig. 2C (dashed line). Canfield29 has shown that the poorly reactive sheet silicates, such as chlorite and illite do not significantly react if exposed to sulphidic conditions over millions of years. Indeed, XANES analyses showed chlorite and illite as the most abundant Fe-mineral phases. The presence of these unreactive silicate phases explains why the ratio of CRS-Fe to total extracted Fe is only around 0.6 (blue regression line in Fig. 2A). An overestimation of CRS-Fe due to organically bound S is unlikely, as organic S has been shown to be largely non-extractable by Cr(II)-solution in Peru Margin sediments (Mossmann et al.30; cf. also Fossing and Jørgensen31), and would cause a perturbation of the linear trend of extracted vs. total Fe (Fig. 2D; cf. Böttcher et al.32). Because of the approximate extraction efficiency of 80% for sequential leaching of Fe (cf. Fig. 2C), the ratio of CRS-Fe to the total Fe-content, measured by X-ray fluorescence (orange regression line in Fig. 2A), is somewhat lower than the ratio of CRS-Fe to extracted Fe. Nevertheless, the same pattern is observed as for the ratio of CRS-Fe to extracted Fe (both with and without unreactive sheet silicates) and the regression lines of all three ratios in Fig. 2A show the same slope. The pool of reactive iron is exhausted in the uppermost few metres (~5 m) below the seafloor, and pyritizable Fe-fractions only make up a minor proportion of the total Fe in the rest of the sediment column.

The precipitation of Fe-sulphides is the result of reductive dissolution of reactive Fe(III) phases by free sulphide. HS− is produced by microbial sulphate reduction and reacts abiotically with different reactive Fe-oxides and oxyhydroxides (Berner26; Rickard and Luther III33). The actual electron transfer from HS− to the Fe(III) has been demonstrated to control rates of iron reduction, whereas binding of HS− on the Fe-mineral surface is not rate limiting (Afonso and Stumm34). The reductive dissolution of Fe minerals is therefore not dependent on the HS− concentration in the porewater, as long as HS− is available. While Fe-sulphide (FeS) can be directly precipitated from the porefluid under sulphidic conditions, the formation of pyrite (FeS 2 ) requires an additional oxidation step. According to Wächtershäuser35 and Thiel et al.36, the oxidation step may be coupled to the reduction of water to H 2 (eq. 1), which, however, may be readily consumed if sulphate is present.

$${\rm{FeS}}+{{\rm{HS}}}^{-}+{{\rm{H}}}^{+}\to {{\rm{FeS}}}_{2}+{{\rm{H}}}_{2}$$ (1)

Consistent with this reaction, Riedinger et al.37 showed that acid volatile sulphide is rapidly converted to CRS if measurable quantities of free sulphide are present in the porewater. Within the sulphidic zone, the depth at which pyrite forms is, thus, limited and controlled by the amount and reactivity of solid-phase Fe(III). Based on the present activity at depth, where sulphide is still produced, it is reasonable to assume that sulphide was always available during the deposition of the 200-m-thick sediment sequence at ODP Site 1229, except, perhaps interrupted by episodes of expanded methanogenic zones, which however, did not cause a significant re-distribution of Fe. This is supported by the observation that the DOP increases rapidly in the uppermost sediment and remains rather constant through the sequence, suggesting that the reactive Fe-phases are rapidly pyritized in the uppermost few metres, independent of how intense the sulphate reduction rates are.

Therefore, δ34S in diagenetic pyrite at ODP Site 1229 reflects the conditions in the near surface porefluid, unless sediment was affected by a permanently shallow SMTZ (i.e. shallower than the 2.5 m that may have been episodically reached; Contreras et al.7) or the sediment was largely exposed to suboxic, i.e. sulphide-free conditions. This means that the pyrite has recorded the isotopic composition of dissolved sulphide in the top few metres of the sediment through time, and the sulphide produced at depth today is not preserved in the pyrite record. Assuming that the sedimentation rate was constant, the S isotope profile could thus be used as a record of average microbial sulphate-reducing activity in the near-surface sediment interval.

Sulphur isotope fractionation between sulphate and sulphide

A high sulphate reduction activity in the top interval of sediment leads to a steep depth-gradient in δ34S (i.e. sulphide is more strongly enriched in 34S in the same depth interval), whereas a low sulphate reduction activity leads to a low depth-gradient in δ34S. Thus, if Fe-sulphide formation occurs over a constant depth range, one would observe less negative/more positive values preserved in the Fe-sulphides for situations in which the sulphate reduction activity was intense, i.e. higher δ34S values indicate higher sulphate reduction activity.

The situation is complicated by the fact that the sedimentation rate may change (Hartmann and Nielsen38; Pasquier et al.39). An increase in sedimentation rate would result in a larger depth range over which the Fe-sulphides form. At the same time, a higher sedimentation rate would lead to a more rapid burial of reactive organic matter, higher sulphate-reducing activity and accordingly steeper sulphate gradients (Meister et al.40). Indeed, an upward shift of the SMTZ could be shown for the Peru margin ODP Site 1229, where the age model in Contreras et al.7 suggests variations in sedimentation rate between 0.01 and 0.8 m/ka. The steeper gradient in δ34S would then also lead to more positive values in the Fe-sulphides.

A further uncertainty could be the isotope enrichment factor. However, large fluctuations in the ε34S are not expected for marine sediments that receive organic matter that – with regard to its composition and content – does not vary substantially over time. In laboratory experiments with sulphate reducing bacteria, ε34S increases when cell-specific sulphate reduction rates become smaller, but this effect is tied to the substrate, respectively the energy yield rather than the total availability of the substrate (e.g., Kaplan & Rittenberg12; Chambers & Trudinger41; Sim et al.17; Wing and Halevy42). One exception to that rule is sulphate reduction coupled to the anaerobic oxidation of methane (AOM), where ε34S depends on methane partial pressures and low methane availability at the SMTZ results in a large ε34S near 70‰ (Deusner et al.43). Such large ε34S values also are common for organoclastic sulphate reduction in marine sediments (Rudnicki et al.44; Wortmann et al.23; Claypool45; Sim et al.17; Pellerin et al.20), and in very shallow sediments (a few 10’s of cm) the observed ε34S can be amplified by disproportionation of S compounds tied to oxidative S-cycling (Canfield and Thamdrup46; Cypionka et al.47; Habicht et al.48). Thus, the isotopic difference between sulphate and sulphide of ~70‰ in the top few mbsf at Site 1229 could reflect a long-term average ε34S between sulphate and sulphide observed in many porewaters of marine sediments (e.g. Claypool45; Jørgensen et al.24; Böttcher et al.22). This offset should be representative for the true separation factor near the zone of pyrite formation. Thus, changes by more than 20‰ in the δ34S record of pyrite at Site 1229 are likely reflecting long-term shifts in the microbial activity in the top metres of the sub-seafloor.

The diagenetic history at Peru Margin Site 1229

For the reconstruction of the diagenetic history, the δ34S record of pyrite can be compared to several other parameters (Fig. 4). The two intervals showing higher δ34S-values also show somewhat elevated concentrations of total organic carbon (TOC). While the TOC data show a large scatter, and the measurements are not available at high resolution, shipboard core scans of the chromaticity value a* (red-green value; D’Hondt et al.2) have been shown to correlate with the abundance of diatom ooze (Meister et al.49; Aiello and Bekins6). Diatom ooze contains high TOC, and a* can therefore be used as a proxy for TOC. Both TOC and a* show two maxima near 30 mbsf and 100 mbsf, respectively. A third maximum in a* occurs near 180 mbsf, where TOC data are lacking. Even though high TOC contents occur in the top few metres, one should take into account that this part of organic matter decays with a power-law with depth and its impact on the future record is not exactly known. Also, this part of the section still lies in the zone of ongoing pyrite formation.

Figure 4 Correlation of δ34S in CRS with color reflectance values a*, TOC, and δ13C of diagenetic dolomites from ODP Site 1229. Color reflectance data are reproduced from (D’Hondt, et al.2); TOC data are compiled from Meister et al.42 and new data); δ13C of dissolved inorganic carbon and dolomite are from Meister et al.10. Full size image

Based on the observed correlation with TOC it is conceivable that intervals of elevated δ34S indeed reflect past episodes of enhanced rates of sulphate-reducing microbial activity. In addition to elevated TOC it is very likely that also the sedimentation rates were larger during these two intervals. Although an age model at sufficiently high resolution is currently not available for these intervals, based on the age model of the last 100 ka at Site 1229 (Contreras et al.7), higher sedimentation rates co-occur with times of enhanced upwelling on the Peru Margin and accordingly also higher primary productivity, higher organic sedimentation rate and higher TOC in the sediment. Apparently, however, the 100 ka glacial-interglacial cyclicity at Site 1229 has not resulted in elevated δ34S throughout the 100-m-thick Pleistocene interval. It is only the large-scale intervals (at 30 and 100 mbsf), which may reflect episodes of enhanced upwelling that lasted long enough to have left a visible imprint in the diagenetic δ34S record.

The intervals with elevated δ34S also correlate with elevated δ13C DIC values in diagenetic dolomite (Fig. 4). There is general consensus that dolomites form in carbonate-free sediment as a result of alkalinity production due to anaerobic metabolic respiration, in particular, sulphate reduction and anaerobic methane oxidation (e.g. Kelts and McKenzie9; Baker and Burns50; Moore et al.51, Meister et al.10). While these processes produce negative δ13C DIC values from the decomposition of organic matter, more positive values in the inorganic carbon result from microbial production of isotopically light methane (Claypool and Kaplan13). The δ13C values in dolomite in the two intervals at 30 and 100 mbsf (Fig. 4, solid diamonds) are far higher than δ13C in DIC of the porewater (line), even within the present methanogenic zones. These values can only be explained by strongly enhanced methanogenic activity in the past. This would be consistent with the interpretation of the S isotope record, which also indicates higher microbial activity during these two intervals.

In conclusion, two episodes of enhanced sub-seafloor microbial activity during early and mid Pleistocene are documented independently in the diagenetic S and C isotope records. Apparently the short-term 100 ka cycles described by Contreras et al.7 are not recorded, but the records are useful for long-term variations in sub-seafloor biosphere activity. In continuous records, a constant ratio of CRS-bound Fe to total Fe provides a useful indicator to exclude that substantial amounts of Fe were re-distributed upon later diagenetic processes and, thus, that the δ34S in pyrite represents robust record of microbial activity. Our study thus demonstrates that diagenetic S and C isotope records bear the potential to trace million-year-scale variations in sub-seafloor microbial activity in the deep-time rock record.