Interstitial water samples from Hole 1132B were taken at a rate of one per core for the first 15 cores and one every other core thereafter, recovery permitting. As a result of poor recovery, only three samples were taken from Hole 1132C (Cores 182-1132C-14R, 15R, and 31R). No samples were taken from Hole 1132A. Samples were analyzed according to the procedures outlined in "Inorganic Geochemistry" in the "Explanatory Notes" chapter. The data are presented in Table T5 and in Figures F15, F16, F17, F18, and F19.
Salinity values increase below the fourth core (30.3 mbsf), reaching 71 in Core 182-1132B-23X (205.1 mbsf) and slowly increasing to a maximum of 80 at 518.1 mbsf (Fig. F15). Chlorinity mirrors this pattern and increases from values near normal seawater to 1275 mM in Core 182-1132C-34R (Fig. F15). The increase in Cl- and salinity suggests the presence of a high-salinity fluid within and below the cored interval.
The concentration of Ca2+ decreases from 11.9 to 8.0 mM at 68.2 mbsf (see Fig. F15), followed by a rapid increase to 39.6 mM at 239.9 mbsf. We infer that the decrease is caused by precipitation, and the increase by dissolution of carbonate minerals. The Mg2+ concentration remains almost constant at values near that of normal seawater for the upper 49.2 mbsf. Below 49.2 mbsf, the Mg2+ concentration decreases from 52.8 to 49.8 mM at 77.7 mbsf, probably as a result of the precipitation of dolomite (Figs. F15, F17). As is the case with Mg2+, the concentration of K+ remains almost constant at values near that of normal seawater for the upper 49.2 mbsf. Below 49.2 mbsf, the K+ concentration increases to 357.6 mbsf, where it reaches a maximum value of 22.6 mM. The excess K+ concentration shows that ~2.5 mM of K+ is lost from the pore waters (Fig. F17), probably caused by ion-exchange reactions with clay minerals. The Sr2+ concentration increases slightly from 83 µM in the upper four cores to 95 µM at 30.3 mbsf. Thereafter, it increases sharply to a maximum concentration of 578 µM at 87.2 mbsf, indicating the onset of carbonate recrystallization (Fig. F15). Note, however, that the high sulfate concentration results in saturation of the pore waters with respect to celestite, thus limiting the maximum Sr2+ concentration to values similar to those observed at Site 1130. The concentration of Li+ remains almost constant at ~37 µM in the upper 30 mbsf. Below this depth are two maxima, 94 µM (106.2 mbsf) and 86 µM (239.9 mbsf), separated by a zone of slightly lower Li+ concentrations (76 mM; Fig. F16). This complex pattern reflects (1) the Li+ supply from the brine and (2) the release of Li+ from clay minerals by ion-exchange reactions in the presence of NH4+ (Shipboard Scientific Party, 1996). The concentration of H2SiO40 remains constant at 48 µM during the first three cores and increases rapidly to 712 µM at 106.2 mbsf. This first maximum is followed by a decrease to 493 µM (153.2 mbsf) and a second maximum of 926 µM at 205.1 mbsf. Below that depth, concentrations of H2SiO40 decline to 166 µM at 546.9 mbsf (Fig. F16). As a result of the dependence of H2SiO40 availability on soluble phases like opal-A, we assume that this pattern reflects the opal-A content of the sediment. In the upper part of Hole 1132B, Fe2+ concentrations are very low and generally remain below the detection limit. Below 169.7 mbsf, Fe2+ concentration increases sharply to values as high as 32.1 µM (220.9 mbsf), falling again to levels below the detection limit by 357.6 mbsf (Fig. F16).
The concentration of SO42- remains near that of normal seawater for the upper 40 mbsf, declining to 24.0 mM at 58.8 mbsf, indicating the onset of sulfate reduction. Below 58.7 mbsf, the concentration of SO42- increases slowly to 54.2 mM at 239.9 mbsf. However, calculation of the excess SO42- shows that the zone of depleted SO42- concentration extends to 357.6 mbsf. Below that depth, SO42- concentration attains a stable plateau of 60 mM (±1 mM) until the base of the cored interval (Fig. F16). With increasing sulfate reduction, alkalinity rises to a maximum of 28.01 mM at 77.7 mbsf and subsequently decreases to 2.37 at 516.7 mbsf (Fig. F16). Ammonium, as a by-product of organic matter decomposition in the sulfate reduction zone, approximately follows the alkalinity distribution, although the maximum concentration is shifted downhole (6654 µM at 125.2 mbsf; Fig. F16). The pH declines slowly from 7.52 at 4.5 mbsf to 7.28 at 39.8 mbsf, and then decreases more rapidly to 6.83 at 49.2 mbsf. Below that depth, pH attains values as low as 6.54 at 186.2 mbsf, increasing thereafter to 7.12 at 357.6 mbsf (Fig. F16).
Although the increase in alkalinity and reduction of SO42- are lower compared to Site 1130, the depletion of Mg2+ is greater at Site 1132. The greater depletion in Mg2+ may be a result of increased amounts of dolomite formation at Site 1132 compared to Site 1130, although it would intuitively seem that higher amounts of carbonate recrystallization should occur at Site 1130 because of the higher alkalinity values.
Unlike Site 1130, Site 1132 shows no exceptionally high salinity/depth gradient. A possible explanation for this is evident if salinity is related to absolute depth with respect to the sea surface (Fig. F18). The salinity values of Sites 1130 and 1132 reach a stable plateau of 82 and 80, respectively, at the same depth below the sea surface, which suggests that a brine with similar salinity is located ~520 m below the sea surface at both sites. Therefore, the difference in the salinity depth/gradient between the two sites can be explained simply by the different distances between the seafloor and brine surface (Fig. F18).
A striking feature of Site 1132 is the absence of vertical gradients in the concentrations of both conservative and nonconservative constituents in the upper 30.3 mbsf. A similar phenomenon was noted during Leg 166 on the western margin of Great Bahama Bank and was attributed to the active flushing of seawater. Possible causes for this phenomenon include thermally induced seawater intake (Eberli, Swart, Malone et al., 1997) or rapid sedimentation.
As is the case with Site 1130, the mineralogy of Site 1132 is characterized by high-frequency alternations between low-Mg calcite (LMC) and high-Mg calcite (HMC). In contrast to Site 1130, Site 1132 shows a higher degree of variability of the minor constituents aragonite and dolomite. Because of the accelerated recrystallization rates (compared to Site 1130; see "Discussion," immediately above), the complete replacement of HMC with LMC and/or dolomite occurs at a much shallower depth than at Site 1130 (Table T6, also in ASCII format). However, isolated peaks of as much as 60 wt% HMC can be observed down to a depth of 155.4 mbsf (Fig. F19). Although the amount of aragonite decreases with depth, concentrations never decrease to zero, as was the case for Site 1130.