We collected interstitial water from 30 samples at Site 1122: 10 from Hole 1122A at depths ranging from 5.90 to 117.20 mbsf, and 20 from Hole 1122C at depths ranging from 67.40 to 591.80 mbsf (Table T13, also in ASCII format). One interstitial-water sample was taken from each core in the upper 100 mbsf. From 100 mbsf to the bottom of the hole, every third core was sampled. Samples from Holes 1122A and 1122C are plotted together in Figure F30. Stratigraphic correlation based on high-resolution data of magnetic susceptibility, GRAPE density, and spectral reflectance demonstrate that ~8 m of the upper part of Hole 1122A are missing (see "Composite Depths"). Comparison between the interstitial-water concentrations of the first sample of Hole 1122A and normal seawater values supports this conclusion. In general, the concentrations of the uppermost samples are close to the seawater composition because of the dominant water exchange at the seawater-sediment interface. The uppermost sample of Hole 1122A (5.90 mbsf) shows a large offset from the seawater values, suggesting that it originated from a depth greater than 5.90 mbsf (see "Composite Depths" ).
Salinities of the interstitial-water samples show a decreasing trend from 5.95 to 272.70 mbsf, and then remain virtually constant down to 460.30 mbsf (Fig. F30). The salinity decrease results from the removal of sulfate by bacterial reduction. The local maximum (35.0) was found at 487.40 mbsf. The salinity increase in the lower part of Hole 1122C reflects the enrichment of major ion concentrations including calcium, magnesium, and sodium, as described below.
The chloride (Cl-) concentration shows a small increase from 560 mM at 5.90 mbsf to the maximum (563 mM) at 34.20 mbsf (Fig. F30), then decreases to 554 mM at 147.90 mbsf and remains between 555 and 559 mM to the bottom of the hole. The chloride decrease seems to parallel the behavior of salinity, which has also been observed at Site 1119 (see "Inorganic Geochemistry" in the "Site 1119" chapter). However, the reduction of chloride may not be caused by the freshwater input, but is probably the result of charge balance control accompanied by complete sulfate utilization.
Interstitial water pH values show a highly variable pattern, ranging from 7.32 to 7.79 (Fig. F30). The highest pH (7.79) occurs at 431.40 mbsf, which coincides with the concentration anomalies of magnesium, strontium, potassium, and silica.
Concentrations of sodium (Na+) follow a similar trend to the chloride concentrations and vary from 438 to 468 mM (Fig. F30).
Alkalinity, sulfate, phosphate, and ammonium concentrations are strongly controlled by the availability of organic matter for bacterial degradation. The alkalinity of interstitial water decreases with depth down to 431.40 mbsf and remains almost constant toward the bottom of the hole (Fig. F30). The small variation at ~100 mbsf is caused by the depth offset between Hole 1122A and Hole 1122C (see "Composite Depths"). The alkalinity maximum at 15.20 mbsf represents the interval of the most intense sulfate reduction. The increase in alkalinity is the result of the production of bicarbonate during bacterial degradation of organic matter, primarily by sulfate reduction. The almost constant alkalinity between 45.10 and 147.90 mbsf coincides with complete sulfate reduction and the onset of methanogenesis. The linear decrease in alkalinity down to 431.40 mbsf presumably is caused by carbonate recrystallization and/or silicate reconstitution processes, which use the bicarbonate ion from interstitial water (Gieskes, 1974).
The sulfate (SO42-) concentration decreases from 3.3 mM at 5.90 mbsf to 0 mM between 15.20 and 122.00 mbsf, followed by an increase to ~26.5 mM at 514.80 mbsf (Fig. F30). Below this depth, sulfate remains almost constant. The sulfate profile is governed by several factors, of which the most important are the abundance (and/or availability) of organic matter and the sedimentation rate. Bacterially controlled sulfate reduction processes have completely depleted sulfate in the pore waters down to ~150 mbsf, resembling nearshore settings with high organic carbon contents and higher bulk accumulation rates such as the Gulf of California (Gieskes et al., 1982) or the Peruvian margin (Suess, von Huene, et al., 1988). The complete utilization of sulfate in the interstitial waters is consistent with the high methane level, as sulfate reduction precedes methane formation during diagenesis of organic matter (see "Organic Geochemistry"). High sulfate levels probably represent the original sulfate concentration that has not been used by bacteria, possibly because of the lack of enough metabolizable organic matter below lithostratigraphic Unit I (see "Lithostratigraphy"). In addition, the comparably low sedimentation rate (see "Age Models and Sedimentation Rates") may have allowed significant amounts of sulfate to diffuse downward and replenish the sulfate reservoir used by bacteria.
Ammonium (NH4+) concentrations remain almost constant down to 53.20 mbsf, then increase to a maximum of 5.90 mM at 147.90 mbsf (Fig. F30). Below 53.20 mbsf, ammonium values steadily decrease down to 431.40 mbsf and then remain constant to the bottom. An increase of the ammonium concentration reflects the intensive bacterial degradation of organic matter, whereas a decrease indicates the results of ion exchange reactions on the surfaces of clay minerals and/or the subsequent incorporation into interlayers of diagenetically formed clay minerals (Gieskes, 1981).
The phosphate (HPO42-) concentrations decrease with depth down to 305.90 mbsf (Fig. F30). Below 305.90 mbsf, the concentrations are relatively constant, but decrease slowly with depth. This trend indicates rather strong first-order removal, suggesting a diagenetic uptake of dissolved phosphate, most likely into sedimentary mineral phases.
Calcium (Ca2+) concentrations initially decrease from a subsurface value of 4.9 to 3.5 mM at ~100 mbsf, because of carbonate precipitation resulting from the buildup of alkalinity during sulfate reduction (Fig. F30). The absence of SO42-, high alkalinity, and high Mg2+/Ca2+ ratios may provide a favorable geochemical environment for dolomite formation (Baker and Kastner, 1981). Between ~300 and ~500 mbsf, Ca2+ concentrations increase strongly up to 18.9 mM.
The profile of magnesium (Mg2+) shows high concentrations in the upper 25 mbsf of Hole 1122A (Fig. F30). High Mg2+ values have been observed almost ubiquitously in anoxic environments and are thought to result from the desorption of Mg2+ from solid phases in rapidly accumulating sediments (Gieskes et al., 1982). From ~25 mbsf, Mg2+ decreases with depth down to 272.70 mbsf, which indicates precipitation of dolomite. Between 300 and 450 mbsf, a distinct Mg2+ anomaly occurs. Concentration variations of these elements may be influenced by diagenetic controls within this lithologic unit or diffusive transport by fluid movement within this interval.
Dissolved strontium (Sr2+) concentrations increase slightly from 75 µM at 5.90 mbsf to 173 µM at 243.70 mbsf and then increase rapidly to a maximum of 466 µM at 370.40 mbsf (Fig. F30). The Sr2+ values remain relatively constant throughout the lower part of Hole 1122C. Increasing strontium concentrations in interstitial waters may originate either from the recrystallization of biogenic carbonate (the associated strontium concentration decrease in the recrystallizing carbonate may be more than one order of magnitude) or the alteration of tephras. The maximum Sr2+ concentrations occur at the same level as the Mg2+ maximum, which indicates an additional input of these two elements into this interval. The change of the Sr2+ gradient at ~460 mbsf coincides with the lithologic boundary between Subunit IIA and IIB (see "Lithostratigraphy").
Dissolved silica (H4SiO4) concentrations increase gradually from 575 µM at 5.90 mbsf to 803 µM at 591.80 mbsf (Fig. F30). The dissolved silica increases indicate progressive diatom dissolution. However, the relatively large scatter of the silica data may reflect the fairly heterogeneous sediment composition.
Potassium (K+) concentrations decrease from 10.6 mM at the subsurface to 6.5 mM at the bottom of the core (Fig. F30). This indicates large-scale removal of K+ into clay minerals that are forming within the sediments.
The dissolved lithium (Li+) concentrations increase slightly from 18 µM at 5.90 mbsf to 87 µM at 399.50 mbsf and increase steeply to a maximum of 231 µM at the bottom of the hole (Fig. F30). Because the Li+ concentration is related to the biogenic silica content (Gieskes, 1981), the general increase indicates release of Li+ during diatom dissolution and silica transformation. The abrupt increase in Li+ concentration in the lower part of the hole, which corresponds to the boundary between Subunits IIA and IIB (see "Lithostratigraphy"), reflects an additional source of Li+ into the pore fluids, because the dissolved silica concentrations decrease in this interval.
The primary controlling factor on the interstitial-water chemistry at Site 1122 is sulfate reduction and methanogenesis, which governs alkalinity, phosphate, and ammonium concentration. In contrast to the complete utilization of sulfate in the upper part of the core, the increased sulfate levels in the middle of the section represent the original sulfate concentrations during sediment deposition, possibly preserved because of a lack of sufficient metabolizable organic matter and low sedimentation rates. Other important chemical profiles are magnesium and calcium concentrations, from which we may deduce the lateral transport of magnesium-rich fluid during the dissolution of carbonate. The general chemical zonations of interstitial waters at Site 1122 correspond to those of lithostratigraphic units and paleontological age divisions. In particular, note the sharp reduction of methane at 260 mbsf, which coincides with the base of the highly pyritized turbidites of the mud-wave sequence.