Interstitial water samples from Hole 1130A were taken at a rate of one per core for the first 15 cores and one every other core thereafter, recovery permitting. No samples were taken from Hole 1130B, and only one sample was taken from Hole 1130C. Samples were analyzed according to the procedures outlined in "Inorganic Geochemistry" in the "Explanatory Notes" chapter. The data are presented in Table T8 and Figures F16, F17, F18, and F19.
Salinity values increase rapidly below the third core, reaching 82 in Core 182-1130A-5H (41.4 mbsf), and remain stable to the base of the cored interval (Fig. F16). Chlorinity shows the same rapid increase from values close to those of normal seawater to 1323 mM at 41.4 mbsf. Below this depth, values remain relatively constant (Fig. F16). We interpret the increase in salinity and chlorinity as a result of the presence of a saline brine within and below the cored interval (see "Discussion").
The concentration of Ca2+ increases from 10.5 to 26.0 mM within the upper 136.4 mbsf (Fig. F16). Between 174.4 and 210.6 mbsf, the rate of Ca2+ change increases, reaching a relatively stable concentration between 210.6 and 321.4 mbsf. The increase is a result of the increasing influence of the brine noted previously. The concentration of Ca2+ relative to Cl- is lower than that of seawater to a depth of 155.4 mbsf (Fig. F18). This depletion is caused by the precipitation of carbonate, triggered by increased alkalinity. Below a depth of 155.4 mbsf, excess Ca2+ indicates net dissolution of carbonate. As is the case with Ca2+, the Mg2+ concentration increases from 54.1 to 119.0 mM as a result of the increasing influence of the underlying brine. However, Mg2+ concentration decreases slightly to 105.0 mM at 321.4 mbsf (Fig. F16). This is even more apparent for the excess Mg2+ concentration shown in Figure F18 and is probably caused by the precipitation of dolomite. Note, however, that even a complete Mg2+ depletion of the interstitial water would only result in the precipitation of minor amounts of dolomite (Swart and Guzikowski, 1988). The concentration of K+ increases from 11.1 to 23.3 mM at 321.4 mbsf (Fig. F16). The increase is caused entirely by the increase in salinity, and the K+/Cl- ratio remains close to that of seawater throughout. The Sr2+ concentration reaches a maximum of 530 然 at 50.9 mbsf, gradually decreasing to 198 然 at 326.1 mbsf (Fig. F16; see "Discussion"). The Li+/Cl- ratio shows a highly variable signal, roughly coincident with changes in lithology (see "Lithostratigraphy"). Silica shows an almost fourfold increase (128 to 633 然) within the upper 27 mbsf, maintaining values between 580 and 758 然, until a marked decrease to values as low as 284 然 between 249 and 287.5 mbsf (Fig. F17). The observed variations are probably caused by variations in the silica content of the sediment. This is particularly evident for the interval between 263.6 and 302.1 mbsf where the silica values decrease by ~50%, and the sediment type changes from a bioclastic packstone to a nannofossil ooze (lithostratigraphic Unit II; see "Lithostratigraphy"). The concentration of Fe2+ increases from values below the detection limit to 7.8 然 at 13.0 mbsf. As a result of the presence of H2S, which causes Fe2+ to precipitate as iron sulfides, Fe2+ concentration decreases with depth and is lower than the detection limit below 60.4 mbsf. With the exception of some scatter in the data, possibly related to contamination, Fe2+ concentrations remain low until 174.4 mbsf. With decreasing H2S content in pore water, Fe2+ concentration increases below 174.4 mbsf and reaches a maximum of 58.81 然 at 268.3 mbsf (Fig. F17).
As a result of the influence of the underlying brine, the concentration of SO42- increases steadily with depth and reaches a maximum of 64 mM at 321.4 mbsf (Fig. F17). This increase is slower than the increase in salinity, which indicates sulfate reduction is occurring. As a by-product of organic matter decomposition within the sulfate reduction zone, the concentration of NH4+ increases from 171 然 to 3508 然 at 41.4 mbsf. Below this depth, the concentration of NH4+ steadily decreases to 257 然 at 321.4 mbsf (Fig. F17). The pH declines from 7.4 at 4.5 mbsf to 6.4 at 27.5 mbsf, remains almost constant until 132 mbsf, and increases slightly thereafter (Fig. F17).
Sulfate reduction at Site 1130 is approximately one-third that of Site 1127 (see "Inorganic Geochemistry" in the "Site 1127" chapter). Because both sites have an abundant supply of SO42- from the underlying brine, the difference is probably a result of variations in the initial organic carbon content of the sediment (see "Organic Geochemistry"). As a consequence of the lower sulfate reduction activity, alkalinity values are also lower (Fig. F17; see "Inorganic Geochemistry" in the "Site 1127" chapter) and less carbonate precipitation occurs (4.2 mM Ca2+ depletion, compared to 8.8 mM at Site 1127). Furthermore, the SO42- concentration is never lower than that of seawater, and interstitial waters are saturated with respect to SrSO4, thus limiting Sr2+ pore-water concentrations (Fig. F18).
The most striking feature of the pore-water chemistry at Site 1130 is the steep salinity gradient of 5.7/m, which compresses the interface between pore waters with normal seawater salinity and the brine into a zone between 13.0 and 31.9 mbsf. By comparison, Site 1126, which yielded the highest salinity values of Leg 182 (106 at 153.8 mbsf; see "Inorganic Geochemistry" in the "Site 1126" chapter), had a salinity/depth gradient of 1.7/m. Despite the steep gradient, salinity reaches only a maximum value of 82. Furthermore, salinity values remain constant, within the error limits, below 41.4 mbsf. The steep gradient, as well as the constant salinity concentrations below 41.4 mbsf (Fig. F16), suggests nonsteady-state conditions.
The mineralogy of the Pleistocene section at Site 1130 is characterized by variations between low-Mg calcite (LMC) and high-Mg calcite (HMC), whereas dolomite, aragonite, and quartz show only minor variations (Table T9, also in ASCII Format; Fig. F19). Although the upper part of the section is dominated by HMC, LMC becomes increasingly dominant with depth, and almost no HMC is observed below the Pliocene/Pleistocene boundary at 230 mbsf (see "Biostratigraphy"). Below the Pliocene/Pleistocene boundary, aragonite shows its first marked decrease. Variations in the abundance of dolomite may relate to dolomite formation during periods of nondeposition and/or periods of high sea level, such as at the present time, when Mg2+ is supplied through diffusion.