A total of 41 whole-round samples were collected at Site 1170 (30 from Hole 1170A and 11 from Hole 1170D) for interstitial water (IW) analysis at the following frequency: three per core in the upper 60 m, one per core from 60-100 mbsf, and one every third core to total depth. Only 33 of the IW samples were used for shipboard analyses; the balance of samples from the upper 60 mbsf of the hole were archived for shore-based analyses. Results of the IW analyses are reported in Table T20 and Figure F29.
Chloride (Cl-), sodium (Na+), and salinity concentrations show little variation in the upper ~450 mbsf at Site 1170 (Fig. F29A). The exception is a 1.4% increase in Cl- from 556 mM in the uppermost sample to 562 mM centered at ~40 mbsf. This increase in Cl- is followed by a decrease to ~557 mM at ~80 mbsf. Below ~450 mbsf, Cl- concentration profiles are highly variable and less than seawater values. Chlorinity values reach a minimum of 461 mM at 707 mbsf, a 17% decrease from mean seawater values.
The sharp boundary in the concentration profiles between ~450 and 470 mbsf is coincident with the base of lithostratigraphic Unit III, near the Eocene-Oligocene transition (see "Biostratigraphy" and "Lithostratigraphy"). The Unit III/IV boundary is defined by the transition from glauconitic sandstones to overlying well-lithified nannofossil limestones. The apparent chemical decoupling of the Cl- and Na+ profiles above and below the boundary, as well as with other pore-water constituents, suggests a lack of diffusional communication across this boundary. Based on the location of the chemical discontinuity, which will be referred to informally as the "boundary" throughout the remainder of this section, we infer that the very hard limestones, which are also a major excursion on the velocity profile (see "Physical Properties"), are acting as a diffusional barrier.
The existence of low-Cl- waters at depth on the STR, as well as on the western Tasmania margin ~400 km to the north of Site 1170, has implications for the origin of these fluids. If the source of the low-Cl- fluid at the two sites is related, then it is unlikely that the fluid is derived externally (i.e., advection of meteoric fluids) because of the large distances between the sites and, in particular, the distal location of Site 1170 relative to any reasonable meteoric source. Instead, the presence of low-Cl- fluids at both sites suggests similar internal processes. Previously, we postulated that the following were possible internal sources for low-Cl- fluid: (1) gas hydrate dissociation, (2) dehydration reactions of hydrous minerals, such as clays and biogenic opal, and (3) clay membrane ion filtration. At both Sites 1168 and 1170, the presence of low-Cl- fluids coincides with older parts of the sedimentary section (connate fluids?) and, in particular, with the onset of methanogenesis and, hence, may be linked, perhaps by water expelled during the organic matter maturation process. However, at present we cannot eliminate any of the aforementioned hypotheses.
Sulfate decreases gradually from seawater values to 17.4 mM, an ~40% reduction. Below the boundary, sulfate has been completely consumed, whereas methane displays a corresponding increase (Fig. F29B). Alkalinity increases to a maximum of 5.6 mM in the upper 450 mbsf and is highly variable below the boundary, varying from 2.2 to 5.7 mM. The pH decreases downcore in the upper 300 mbsf but increases slightly to the boundary. Below the boundary, pH varies conversely with alkalinity, reaching a maximum of 8.0 at 707 mbsf. Ammonium steadily increases downcore to 450 mbsf, below which the trend continues to increase to a maximum of ~1600 然 but is more variable than above 450 mbsf.
The major changes in the SO42-, NH4+, and alkalinity profiles are largely a function of microbially mediated organic matter degradation. The higher TOC content of the Eocene sediments below the boundary (see "Sedimentary Geochemistry"), combined with the diffusional barrier above, leads to complete sulfate reduction, methanogenesis, and continued production of NH4+ in this interval. The dramatically sharp break in the sulfate and methane profiles between 450 and 470 mbsf provides excellent evidence for a diffusional barrier near the Eocene-Oligocene transition.
Dissolved silica concentrations range from 561 to 1235 然 in the upper 450 mbsf with distinct shifts at 120 (negative) and 295 mbsf (positive) (Fig. F29B). Below the boundary, H4SiO40 drops precipitously to <200 然, decreasing to a minimum of 23 然 at the base of the hole.
The distinct shifts in H4SiO40 in the pore fluids correspond to shifts in biogenic silica abundance in the sediments (see "Biostratigraphy" and "Lithostratigraphy"). The rapid decrease across the boundary is likely a response to the overall decrease in biogenic silica associated with the dramatic shift in depositional environments across the boundary. In addition, opal-CT was noted in many of the XRD patterns of samples below 450 mbsf. Therefore, the lower H4SiO40 values in the lower 470 mbsf may also be the result of recrystallization of opal-A to opal-CT as previously observed (e.g., Baker, 1986), although Gieskes (1981) noted that a significant decrease in H4SiO40 content usually does not occur until the opal-CT to quartz transition.
Calcium and Sr2+ increase with depth in the upper 400 mbsf, reaching a maximum of 15.5 mM and 969 然, respectively, between 300 and 400 mbsf (Fig. F29C). Below the boundary at ~450 mbsf, Sr2+ sharply decreases to a uniform value of ~325 然. Conversely, Ca2+ concentrations remain elevated relative to standard seawater below the boundary, varying between 13.2 and 14.6 mM.
The increase in Sr2+ and Ca2+ values with depth in the upper part of the sedimentary section, where the sediments are composed principally of nannofossil ooze, is likely the result of dissolution and/or recrystallization of calcite. If the profiles are solely the result of calcite dissolution, then we would not expect the observed increase in Sr2+/Ca2+ ratios because the proportion of Ca2+ input would be much greater (Fig. F30) (Baker et al., 1982). Unlike most of the other IW constituents, Ca2+ does not show a shift across the diffusional barrier, which suggests dissolution of a Ca-rich phase(s) or exchange reactions occurring both above and below the boundary. Because there is very little carbonate in the sediments and pore-water Sr2+ is low below the boundary (Fig. F29C), it seems unlikely that the elevated Ca2+ in the lower part of the cored interval is derived from calcite dissolution. Elevated pore-water Ca2+ content (and decreasing Mg2+ and K+; see "Magnesium, Potassium, and Lithium",) has been previously attributed to volcanic matter alteration (e.g., Lawrence et al., 1975). Smear-slide analysis noted the presence of volcanic glass in various stages of preservation in Hole 1170D, (see "Lithostratigraphy"), and alteration of this glass may be the source of Ca2+ in these carbonate-poor sediments.
Magnesium (Mg2+) and potassium (K+) concentrations decrease steadily from the topmost sample to the base of lithostratigraphic Unit III at ~460 mbsf (Fig. F29C). At the boundary, the Mg2+ and K+ profiles are marked by an abrupt decrease (36.5 to 22.4 mM for Mg2+ and 8.8 to 5.0 mM for K+). Below the boundary, Mg2+ and K+ continue to decrease to the base of the hole reaching minimums that are ~73% less than normal seawater. Lithium (Li+) concentrations are uniformly low (<60 然) in the upper 450 mbsf, decreasing to <10 然 between ~50 and 200 mbsf. Below the boundary, Li+ concentrations sharply increase to ~400 然 and remain uniform at this concentration to the base of the cored interval.
As at the previous sites, the decreases in Mg2+ and K+ are highly correlated (r = 0.995). We interpreted the decreasing concentration profiles at Sites 1168 and 1169 to indicate that Mg2+ and K+ were being consumed at depth, likely associated with silicate reactions. The sharply decreasing profiles across the boundary generally support that hypothesis. However, Mg2+ and K+ do increase significantly (28% and 20%, respectively) above the boundary within the largely pelagic sediments. Because there appears to be virtually no diffusional communication across the boundary, the decrease in Mg2+ and K+ suggests there is a sink, presently unrecognized, for these elements in the upper 450 mbsf. A typical Mg2+ sink in carbonate-rich sediments is dolomitization; however, no dolomite has been observed in smear-slide or XRD analyses. In addition, we do not expect significant dolomitization in pore fluids with appreciable SO42- and low alkalinities (Baker and Kastner, 1981; Compton, 1988).
The presence of a diffusion barrier provides insights into the location of sources and sinks for Li+ that were not apparent in the profiles from the previous sites. The decrease in Li+ concentration to values lower than seawater between ~50-200 mbsf indicates that an active Li+ sink exists in this depth interval, although the nature of the sink is unknown. In contrast to the smooth Li+ profile observed at Site 1168, the much higher Li+ contents below the diffusional boundary at Site 1170 suggest that the major source of Li+ to the IW is likely from silicate phases and not from biogenic carbonates. This observation is also compatible with silicate reactions consuming Mg2+ and K+ below the boundary, which also may be related to volcanic matter alteration.
The fortuitous location of the diffusional boundary near the transition between the siliciclastic and pelagic sediments provides important insights into the location of sources and sinks for many of the pore-water constituents that were not apparent at the previous sites. In essence, the diffusional barrier isolates reactions occurring in the siliciclastic sediments and the section below the cored interval vs. the younger pelagic sediments. The Eocene organic-rich, clayey glauconitic siltstones and sandstones are the major source for Li+ and probably the low-Cl- fluids. In addition, Mg2+ and K+ are being consumed by reactions in these same sediments and/or the section below the cored interval. The bulk of sulfate reduction in the sedimentary sequence occurred in these Eocene organic-rich siliciclastics. Not surprisingly, the primary source for dissolved silica and Sr2+ are the pelagic sediments. A Li+ sink is also present within the nannofossil ooze and chalks. Calcium is being released to the waters from the pelagic sediments and, surprisingly, from the siliciclastics as well. Another surprise provided by the presence of the boundary is the considerable sink for Mg2+, K+, and Li+ in the pelagic sediments. The origin of these unexpected sources and sinks is unknown at present.