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doi:10.2204/iodp.proc.346.103.2015

Geochemistry

Site U1422 is located ~40 km from Site 795 (Shipboard Scientific Party, 1990). As such, the sedimentary record and geochemical profiles were expected to be similar. Several processes impact the pore water and solid-phase geochemistry at Site 795. Over the upper 200 meters below seafloor (mbsf), the modest degradation of organic carbon drives a series of microbially mediated reactions. Leg 127 found these reactions lead to a sulfate–methane transition (SMT) at ~80 mbsf, as well as precipitation of authigenic minerals such as sulfides and carbonates. At Site 795, the opal-A–opal-CT transition occurs between 290 and 325 mbsf, where diatoms and radiolarians recrystallize (Shipboard Scientific Party, 1990; Nobes et al., 1992). Also, basalt was encountered at 684 mbsf (Shipboard Scientific Party, 1990). Although the latter processes occur below the 200 m target depth of Site U1422, they affect overlying pore water chemistry through diffusion gradients (Murray et al., 1992).

The initial and main focus of the geochemistry program at Site U1422 was to better delineate the diagenetic environment in the upper 200 m of sediment. In particular, we desired to refine sediment and interstitial water profiles that reflect inputs and outputs from organic matter degradation. This approach included determination of the dissolved species examined at Site 795 at lower depth resolution and dissolved species not examined at Site 795 (Mn and Ba).

As became clear from rapid alkalinity and methane analyses of squeezed interstitial water (IW-Sq) and headspace (HS) samples, the diagenetic environment at Site U1422 is somewhat similar to that at Site 795, but several key reactions occur much shallower in the sedimentary column. For example, the SMT occurs at ~30 m CSF-A and gas voids become common below ~50 m CSF-A. This difference prompted modification of the basic plan for Site U1422 (see “Geochemistry” in the “Methods” chapter [Tada et al., 2015b]). Effectively, two experiments were added to the program. The first was a study to determine if the color of interstitial water changed downhole at a site with modest methane concentrations; the second was a study to examine alkalinity in samples acquired by Rhizons across the SMT.

Sample summary and analytical issues

During operations at Site U1422, the geochemistry team collected and analyzed a range of samples. These included the following (Tables T10, T11, T12, T13, T14):

• 1 mudline (ML) water sample from the inside of the core liner above the sediment of Core 346-U1422C-1H. We collected 50 mL.

• 47 interstitial water samples from whole-round squeezing (IW-Sq) from Holes U1422A (2 samples) and U1422C (45 samples) at nominal 4.5 m spacing. For most samples, we collected between 35 and 45 mL of interstitial water.

• 33 interstitial water samples from Rhizons (IW-Rh) from the bottom sections of Core 346-U1422E-3H and the top sections of Core 4H. This interval was chosen during drilling because it should span the SMT, based on alkalinity and methane data generated for Hole U1422C, basic geochemical theory, and rapid cross-hole stratigraphic correlation. We collected between 4 and 10 mL from each Rhizon. With these volumes, a limited number of analyses were conducted (Table T11).

• 47 sediment samples from the interstitial water squeeze cakes of Holes U1422A (2 samples) and U1422C (45 samples).

• 50 HS gas samples. Most of these (43) were paired with IW-Sq samples, and 7 were placed at section breaks in Cores 346-U1422E-3H and 4H where Rhizons were taken.

• 7 Vacutainer (VAC) samples recovered from gas expansion voids with air-tight syringes from Holes U1422C and U1422D.

Beyond the rapid pace of drilling at Site U1422, which meant a large fraction of time devoted to squeezing and weighing, two issues affected our geochemistry program at this location. First, the chloride titrator, the ion chromatograph, and the spectrophotometer each needed repairs in the middle of analysis. As noted below, there also appears to have been problems with our inductively coupled plasma–atomic emission spectroscopy (ICP-AES) analyses of Ba and Sr. Additionally, as discussed below, several whole rounds taken for squeezing seem to have been contaminated on recovery.

Carbonate and organic carbon

Sediment recovered from Site U1422 had only minor amounts of CaCO₃ (Fig. F23). Some samples that may have represented thin carbonate turbidite layers (as observed by the micropaleontologist group) had CaCO₃ contents as high as 20 wt% at 164 m CSF-A. However, most samples had CaCO₃ values <1 wt%. Total organic carbon (TOC) content varied between 1 and 3 wt% in the upper 50 m CSF-A (Table T10). Below 50 m CSF-A, TOC contents are consistently <1.5 wt%. Across all samples, TOC contributes the vast majority (94%) of the total carbon. Total nitrogen content is <0.40 wt% in samples from Site U1422.

Over the uppermost 50 m CSF-A, the carbonate and TOC data at Site U1422 appear similar to those at Site 795 (Shipboard Scientific Party, 1990). Between 50 and 160 m CSF-A, TOC is generally lower at Site U1422. The variance in TOC values should be considered when selecting and processing samples for certain organic geochemistry analyses (e.g., compound specific isotopes of biomarkers). Further analysis of discrete samples may yield short depth intervals with higher TOC contents, as might be indicated by darker core material.

Dissolved iron and manganese

The Fe profile (Fig. F24) displays a series of peaks with depth. Such Fe profiles have been documented at sites where organic carbon degradation leads to generation of methane and hydrogen sulfide (e.g., ODP Site 1230 at the Peru margin [Shipboard Scientific Party, 2003a] and ODP Site 1244 at the Hydrate Ridge [Shipboard Scientific Party, 2003b]).

The Mn profile (Fig. F24) is more straightforward. Dissolved Mn concentration is close to the detection limit (0.3 µM) in the mudline sample. Below, concentrations rise to a peak of 111 µM centered at ~4 m CSF-A and decrease to ~11 µM at 35 m CSF-A. From ~35 to 70 m CSF-A, dissolved Mn concentrations remain low with values between 6 and 10 µM. Values then increase to 30 µM at 103 m CSF-A. Concentrations range between 20 and 40 µM from 80 m CSF-A to the maximum depth drilled, 202 m CSF-A.

At locations with even modest organic matter input, dissolved Fe and Mn concentrations typically rise in shallow sediment because of Fe and Mn oxide dissolution (e.g., Froelich et al., 1979). For Mn and at sites with high alkalinity, a common phase is rhodochrosite (MnCO₃). Rhodochrosite has been documented between 50 and 700 mbsf at ODP Site 799 (Matsumoto, 1992). Another possibility is Mn-rich calcite, as perhaps indicated by the low dissolved Ca concentrations at this depth (Fig. F23). In any case, our results suggest precipitation of an Mn-bearing carbonate phase between 35 and 80 m CSF-A at Site U1422.

Dissolved Fe exhibits a more complicated profile with depth. In some cases, this may occur because Fe can form a series of authigenic minerals, including oxides (magnetite), sulfides (greigite and pyrite), carbonates (ankerite and ferroan dolomite), and Fe-rich clay (glauconite). Horizons of pyrite, ferroan dolomite, and glauconite have been documented in the upper few hundred meters of sediment at several sites in the marginal sea (Föllmi and von Breymann, 1992; Matsumoto, 1992). The complex Fe profile at Site U1422 may indicate that various Fe-rich minerals are precipitating and dissolving across multiple depth intervals in the sediment column. The Fe profile may also be influenced by dissolved iron reacting with HS^– during the anaerobic oxidation of methane (AOM), as discussed below.

Fe and Mn minerals are important to the paleoceanographic objectives of the expedition because they impart color and magnetism to the sediment. In principle, with very detailed pore water analyses, precise depths of Fe and Mn mineral dissolution and precipitation could be identified. Rhizon sampling at Site U1422 suggests that such Fe and Mn profiles can be generated (Fig. F24).

Alkalinity, ammonium, and phosphate

Alkalinity, NH₄⁺, and PO₄^3– exhibit similar profiles at numerous locations with significant amounts of organic matter degradation. This is the case at Site U1422 (Fig. F25).

Alkalinity increases from 2.5 mM in the mudline sample to 26 mM at 35 m CSF-A. The value for the seafloor is close to that measured for bottom water in the region (2.4 mM; Tishchenko et al., 2012). Below this rapid rise, alkalinity steadily decreases to 18 mM at 202 m CSF-A. The prominent kink in the alkalinity profile is the key to understanding much of the geochemistry at Site U1422. Notably, interstitial water samples from squeezers and Rhizons at similar depths had similar alkalinity (Fig. F25).

Ammonium increases gradually from below detection limit in the mudline sample to >3000 µM at 202 m CSF-A. The profile is slightly concave downward. Importantly, there is no inflection in the NH₄⁺ concentration profile at ~30 m CSF-A, and the molar ratio of alkalinity to NH₄⁺ is ~6:1 at ~200 m CSF-A.

Phosphate increases rapidly from 1.52 µM in the mudline sample to ~75 µM at 16 m CSF-A. The rise is greatest over the uppermost 5 m below the seafloor. Below a prominent peak at 16 m CSF-A, the PO₄^3– profile decreases gradually, with values <20 µM at 202 m CSF-A.

These profiles reflect a combination of several processes. As occurs at many locations, solid organic carbon with a generic Redfield composition of [(CH₂O)₁₀₆(NH₃)₁₆(H₃PO₄)] decomposes to release HCO₃^–, NH₄⁺, and PO₄^3– (Froelich et al., 1979; Berner, 1980; Murray et al., 1992). Alkalinity often approximates HCO₃^– concentrations in the marine environment. Therefore, there are downhole increases in the three profiles that are somewhat similar to the Redfield ratio. However, both HCO₃^– and PO₄^3– can precipitate into (or onto) authigenic minerals, so their concentrations decrease with depth. Beyond the formation of authigenic carbonate phases (Matsumoto, 1992), previous work on deep-sea sediment sequences from the marginal sea has documented significant formation of francolite, a phosphate mineral (Föllmi et al., 1992). Manganese oxides and hydroxides in the upper 1 m below the seafloor may also concentrate P, which could be released to interstitial water as they dissolve with burial (Cha et al., 2005).

The classic explanation for the C-N-P profiles, however, neglects some basic observations. The alkalinity kink at ~35 m CSF-A indicates significant net generation of HCO₃^–, HS^–, or both across a thin horizon at this depth (it is important to note here that HS^– contributes to alkalinity [Gieskes and Rogers, 1973]). The lack of a prominent NH₄⁺ inflection at ~35 m CSF-A further indicates that degradation of solid organic carbon does not produce the anomalous input of alkalinity at this depth. However, AOM is an obvious source for generating alkalinity in shallow-marine sediment (Borowski et al., 1996). This reaction can be expressed as (Reeburgh, 1976)

The significance of AOM to the geochemistry at Site U1422 becomes clearer below.

Volatile hydrocarbons

Two qualitative observations concerning gas are worth noting. We began to detect the odor of hydrogen sulfide with Core 346-U1422C-3H, and the odor disappeared deeper than Core 6H (as well as at corresponding depths in Holes U1422D and U1422E). Gas expansion cracks began forming between 50 and 60 m CSF-A, and major gas voids appeared in Core 346-U1422C-11H and several following cores.

Methane is, by far, the dominant hydrocarbon gas (>99.9%) in all HS samples from Site U1422 (Table T12). The same is true for the VAC gas samples taken (Table T13).

Methane concentrations in HS samples are nearly at the detection limit over the upper 30 m CSF-A. At approximately this depth, CH₄ values increase rapidly, reaching 3600 ppmv at 34.6 m CSF-A. Crucially, the rise in CH₄ occurs at the depth where the aforementioned inflection in pore water alkalinity occurs (Fig. F25). Deeper than 44.1 m CSF-A, CH₄ concentrations fluctuate between 16,600 and 40,300 ppmv to the bottom of the hole at 202 m CSF-A. Ethane concentrations range between 0 and 15.2 ppmv across all samples. The methane to ethane ratio generally exceeds 1200. No heavy hydrocarbons were detected. The near absence of higher molecular weight hydrocarbon gases strongly suggests that CH₄ at Site U1422 derives from microbially mediated methanogenesis, as opposed to thermogenesis, despite the very high geothermal gradient (see “Downhole measurements”).

As noted in “Geochemistry” in the “Methods” chapter (Tada et al., 2015b), HS sampling proceeded with a specific goal of quantifying CH₄ concentrations in interstitial water. A measured 3–5 cm³ volume of wet sediment was placed into a saturated NaCl solution for HS analyses. When concentrations are normalized to wet sediment volumes, the CH₄ profile changes somewhat, but the main features remain (Fig. F26).

Modest amounts of CH₄ are generated via microbial activity at Site U1422. At depth and with hydrostatic pressure, CH₄ at in situ conditions is dissolved in pore water. However, during core recovery and handling, CH₄ escapes solution, resulting in gas cracks and voids in the sediment cores (Paull et al., 2000). Methane concentrations in HS gas samples deeper than ~50 m CSF-A are not meaningful because of degassing.

Within the sediment column, high CH₄ concentrations at depth lead to a concentration gradient and upward flux of CH₄. The upper segment of this gradient is observed between ~30 and 50 m CSF-A. At 30 m CSF-A, CH₄ rising from depth reacts with SO₄^2– diffusing down from the seafloor as a result of AOM (Equation 1), which results in a relatively thin SMT (Fig. F26).

One consequence of AOM is the generation of HS^– (Equation 1). The H₂S odor in Cores 346-U1422C-3H through 6H represents degassing of this AOM product. The smell occurs over a broad but well-defined depth horizon because HS^– diffuses both upward and downward from the SMT. The ends of these diffusion paths are where HS^– reacts with dissolved Fe.

Microbial methanogenesis of organic matter generates HCO₃^– and NH₄⁺. As a consequence, alkalinity and NH₄⁺ concentrations can become very high at sites with significant methanogenesis. Although NH₄⁺ concentrations are fairly high at depth, alkalinity concentrations are not exceptional, giving rise to a low C:N ratio in pore water. This low ratio suggests removal of HCO₃^– relative to NH₄⁺ and may indicate authigenic carbonate precipitation at depth.

Yellowness

All interstitial water appeared visually clear, unlike waters at sites with significant methanogenesis in the upper few hundred meters below the seafloor (e.g., ODP Site 808 at the Nankai Trough [You et al., 1993], Site 1230 at the Peru margin [Shipboard Scientific Party, 2003a], and Site 1244 at the Hydrate Ridge [Shipboard Scientific Party, 2003b]). Thus, we thought Site U1422 might provide a good reference location to generate a background depth profile of color to compare to sites where interstitial water becomes yellow.

In contrast to expectations, interstitial water at Site U1422 exhibits an obvious color profile when examined for absorbance at 325 nm (Table T14). In a general sense, color tracks alkalinity (Fig. F27). However, the details are more complicated. First, the color increase with depth is slightly steeper and shallower than the rise in alkalinity. Second, absorbance scans of multiple interstitial water samples across a range of wavelengths shows that interstitial water “color” actually corresponds to a broad absorbance maximum with a center that can vary between ~270 and 290 nm.

Pore water at Site U1422 likely contains low but significant amounts of colored dissolved organic matter (CDOM), which is generated through the microbial decomposition of solid organic carbon. The composition and amount of this CDOM seems to change with depth.

Bromide

Although not typically examined during drilling expeditions, pore water Br^– concentration was measured at Site U1422 (Table T11). However, all values were between 0.85 and 0.98 mM, a range only slightly greater than the precision of Br^– analyses using the shipboard ion chromatograph (see “Geochemistry” in the “Methods” chapter [Tada et al., 2015b]). In any case, these values are slightly greater than the Br^– concentration of theoretical JSPW (0.84 mM) (see “Geochemistry” in the “Methods” chapter [Tada et al., 2015b]).

At sites with major methanogenesis, Br^– concentrations can exceed 2 mM (Martin et al., 1993; Egeberg and Dickens, 1999). This is because organic carbon landing on the seafloor incorporates Br^– from seawater and releases it to interstitial water upon degradation. The relatively low Br^– concentrations at Site U1422 are consistent with other indicators for only modest organic diagenesis.

Sulfate and barium

Sulfate is 29.3 mM in the mudline sample (Table T11), which is the value of reference seawater but greater than expected for theoretical JSPW (28.5 mM) (see “Geochemistry” in the “Methods” chapter [Tada et al., 2015b]). From the seafloor, SO₄^2– decreases to below detection at ~35 m CSF-A (Fig. F28A). The decrease is slightly concave downward but close to linear. Moreover, the depth where SO₄^2– drops below the detection limit is the same as that for the inflections in alkalinity and dissolved CH₄ (Fig. F26).

The Ba profile contrasts starkly with the SO₄^2– profile (Fig. F28). Interstitial water Ba concentrations are very low (<30 µM) from the seafloor to 20 m CSF-A. From this depth and for the next ~10 m downhole, Ba concentrations increase almost exponentially, surpassing 1200 µM by 35 m CSF-A. Values then increase with a concave-downward profile, such that Ba concentrations reach as high as 6000 µM near the bottom of the site.

The near-linearity of the SO₄^2– profile strongly suggests that, within the upper 35 m of sediment, net SO₄^2– removal dominantly occurs at the SMT. This suggestion is consistent with AOM being an important reaction at Site U1422. Beyond the aforementioned effects of AOM, the removal of SO₄^2– impacts the sedimentary Ba cycle (Dickens, 2001). Small grains of barite (BaSO₄) are a ubiquitous component of surface sediment deposited at intermediate to deepwater depths in the ocean (Dehairs et al., 1980). In sediment sections with even moderately depleted pore water SO₄^2–, solid barite is preserved because the barite equilibrium constant is extremely low. However, when pore water SO₄^2– concentrations approach zero, barite begins to dissolve (von Breymann et al., 1992):

Consequently, at sites with significant amounts of CH₄ and AOM, sedimentary barite begins dissolving as it is buried through the SMT. This liberates Ba²⁺ to interstitial water at concentrations related to the equilibrium constant (Fig. F28B). Upward diffusion of dissolved Ba²⁺ can lead to a solid barite front, which has been found just above the SMT at multiple locations (von Breymann et al., 1992; Dickens, 2001) and might be expected at Site U1422.

The SO₄^2– and Ba profiles (Fig. F28A) further indicate a problem in sample collection. Below the SMT, SO₄^2– concentration should be zero. However, several samples between 120 and 170 m CSF-A have detectable SO₄^2– and anomalously low Ba concentrations, at least compared to surrounding samples. The depth interval containing these samples has abundant gas voids. Such voids are drilled on the catwalk to release gas and are subsequently compressed to make a continuous sediment core. We think that some interstitial water samples were taken across intervals that were pressed together, so they have been contaminated with small amounts of drilling fluid (seawater) that induce barite precipitation.

Calcium, magnesium, and strontium

The Ca, Mg, and Sr profiles (Fig. F29) are somewhat similar to those at Site 795 (Shipboard Scientific Party, 1990) and at many locations where basalt lies within several hundred meters of the seafloor (Gieskes and Lawrence, 1981).

Calcium concentration is 10.1 mM in the mudline sample (Table T11), which is slightly lower than the value expected for theoretical JSPW (10.4 mM) (see “Geochemistry” in the “Methods” chapter [Tada et al., 2015b]). From the seafloor, values decrease to 6 mM at 35 m CSF-A. Calcium concentrations then increase steadily with depth, reaching 9.1 mM at 202 m CSF-A.

Magnesium concentration is 52.5 mM in the mudline sample, which is close to the value expected for theoretical JSPW (52.7 mM) (see “Geochemistry” in the “Methods” chapter [Tada et al., 2015b]). From the seafloor, values decrease to 41 mM at 10 m CSF-A. Magnesium concentrations then decrease steadily with depth, reaching 19 mM at 200 m CSF-A.

The Sr profile is somewhat erratic at Site U1422. Because other trace metal profiles are smooth, we suspect an issue measuring Sr on the ICP-AES. Nonetheless, Sr concentration is 93 µM in the mudline sample, which is close to the value expected for theoretical JSPW (see “Geochemistry” in the “Methods” chapter [Tada et al., 2015b]). The uppermost sample below the seafloor at 1.45 m CSF-A contains 82 µM Sr. From this depth, Sr values increase to 100 µM at 34 m CSF-A. Deeper than 34 m CSF-A, the Sr profile varies between 100 and 120 µM.

The Ca, Mg, and Sr profiles likely result from several processes. Almost certainly, the low in Ca across the SMT reflects the precipitation of carbonate:

or

Basically, HCO₃^– produced by AOM (Equation 1) reacts with Ca²⁺ diffusing from the seafloor and from deeper in the sedimentary column. At some locations, including those within the body of water between the Eurasian continent and the Japanese Islands, the loss of pore water Ca at the SMT coincides with the precipitation of calcite (Snyder et al., 2007); at other locations, the loss coincides with the occurrence of dolomite (Moore et al., 2004). Without detailed examination of the sediment and flux-based calculations, it is difficult to argue for a particular carbonate mineral, especially because the inflections in Ca and Mg do not align with respect to depth.

Although not obvious from operations at Site U1422 but clear from the geochemistry at Site 795, alteration of ash and basalt (at nominally 500 mbsf) drive the overall rise in Ca and drop in Mg with depth. Essentially, the Ca and Mg trends at 200 m CSF-A extend for another 300 m.

Salinity, chlorinity, sodium, and pH

Salinity decreases from 35 in the mudline sample to a minimum of 30 at 202 m CSF-A. The decrease in salinity is manifest in the Cl^– and Na concentration profiles, although the Cl^– values (Table T11) may be compromised at Site U1422 because of a faulty electrode during titration that was discovered after moving to Site U1423. Chloride and Na concentrations are 539 and 470 mM, respectively, in the mudline sample, which compare to 545 mM and 468 mM, respectively, in theoretical JSPW (see “Geochemistry” in the “Methods” chapter [Tada et al., 2015b]). Interstitial water concentrations of both elements drop rapidly over the upper 15 m below the seafloor and more slowly over the next 185 m.

pH varies between 7.5 and 8.0, with a mean of 7.7. This is consistent with interstitial water at many drill sites. There are no clear trends in pH, as values fluctuate over a relatively short distance. This lack of trend is true in Rhizon samples as well.

Interstitial water with relatively low salinity and dissolved Cl^– concentration was documented at Site 795. The initial drop in salinity beneath the seafloor was ascribed to loss of SO₄^2– (Shipboard Scientific Party, 1990). However, this is unlikely, as discussed more fully in “Geochemistry” in the “Site U1423” chapter (Tada et al., 2015c).

Potassium

Potassium concentration is ~10 mM in the mudline sample (Table T11). After an increase to 13 mM in the upper 10 m, K concentrations decrease steadily, reaching 8 mM at 202 m CSF-A. The rise in K immediately below the seafloor is intriguing and may reflect exchange during authigenic mineral formation. The decrease in K with depth perhaps results from further reactions with ash and basalt (Murray et al., 1992), although formation of glauconite would also remove K⁺ from interstitial water (Föllmi and von Breymann, 1992).

Lithium and boron

Lithium concentration (Fig. F30) is 28 µM in the mudline sample. Li values decrease to 20 µM at 11 m CSF-A (Table T11). Below this depth, Li concentrations increase to 108 µM at the bottom of Hole U1422C. Lithium concentrations in Rhizon samples are similar to those obtained by squeezing over the same interval.

Previous work at Site 795 suggests that the Li profile is dominated by biogenic silica recrystallization (Murray et al., 1992), although this was a speculative interpretation. Alteration of ash may also be an important process.

Boron concentration (Fig. F30) is 451 µM at the mudline and increases to 2637 µM at the bottom of Hole U1422C. Deeper than 25 m CSF-A, B concentrations increase more rapidly compared to shallower depths, suggesting that the SMT may affect B concentrations in interstitial water. We observed oscillations in concentrations from sample to sample. The samples were run on the ICP-AES in random order, and precision was calculated to be 0.4%, eliminating the idea that the oscillating trend below 25 m CSF-A is an artifact of instrumental analysis. We hypothesize that the oscillations are caused by variations in sample acquisition, such as squeezing or core flow.

Silica

Silica concentration in interstitial water at Site U1422 (Fig. F31) increases from an initial value of 84 µM in the bottom water (0 m CSF-A) to a maximum value of 1166 µM at 188 m CSF-A (Table T11). The opal-A/opal-CT boundary was not drilled at Site U1422, so the expecsted decrease in dissolved H₄SiO₄ at depth (Shipboard Scientific Party, 1990) was not observed. Notably, dissolved silica concentrations vary significantly over short distances, unlike the profiles for most other elements.

The profile of Si measured using the ICP-AES and the profile of H₄SiO₄ measured using the spectrophotometer match fairly well (Fig. F31). As the two sets of analyses were analyzed from independent aliquots, each of which involved further processing, any variability observed in the silica profiles is unlikely to have been caused by sample preparation or instrumental imprecision. The high degree of variability in the Si profile may reflect changing lithology, including diatomaceous sediment and turbidites in Unit II.

Summary

Geochemical profiles at Site U1422 are somewhat similar to those at nearby Site 795 (Shipboard Scientific Party, 1990). The major difference is that Site U1422 has much higher CH₄ concentrations. The high CH₄ concentrations at depth lead to an upward flux of CH₄ through the sediment column and AOM across a relatively shallow and thin SMT (~35 m CSF-A). In turn, this shallow zone of AOM impacts other species. The generation of HS^– leads to reactions with dissolved Fe, the generation of HCO₃^– leads to reactions with dissolved Ca, and the loss of SO₄^2– dissolves barite.

These detailed interstitial water profiles are amenable to numerical modeling during future shore-based studies. However, because the organic carbon input is highly variable, a full understanding of the system may require a nonsteady-state approach.

Furthermore, there is also an issue regarding alkalinity. As noted above, alkalinity approximates the sum of HCO₃^– and HS^–. In retrospect, had we predicted that a shallow SMT would be found at Site U1422, we could have measured dissolved inorganic carbon, HS^–, and total sulfur. This would enable the direct determination of the proportion of alkalinity caused by HCO₃^– and HS^–.

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