INORGANIC GEOCHEMISTRY

Forty-nine whole-round samples were taken from Hole 1168A for interstitial water (IW) analyses at the following frequency: three per core in the upper 60 m, one per core from 60 to 100 m, and one every third core to total depth. Only 38 of the 47 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 IW analyses are reported in Table T19 and Figure F32.

The conservative parameters, chloride (Cl^-), sodium (Na⁺), and salinity, exhibit little change in the upper 100 mbsf (Fig. F32; Table T19). Chloride increases downhole ~1% in the upper 25 mbsf to 564 mM, returning to seawater values in the underlying 75 mbsf; sodium varies from ~470 to 480 mM through the same interval. From 100 mbsf, Cl^- and Na⁺ gradually decrease to 267 mbsf, where a sharp decrease exists to the base of the hole. The less sensitive salinity measurement does not show the same sharp decrease. In the interval of sharply decreasing values, Cl^- and Na⁺ profiles covary and are marked by multiple distinct maxima and minima rather than smoothly decreasing values. The maximum dilution relative to seawater is 21% for Cl^- (440 mM), 17% for Na⁺ (402 mM), and 28.5% for salinity (25). Previous studies (e.g., McDuff, 1985; Schrag et al., 1996) have attributed increases in Cl^- in the upper 50 mbsf to increases in salinity during the last glacial maximum. The 1% increase in Cl^- centered at 20 mbsf is consistent with this interpretation.

Low Cl^- in marine pore waters has been observed in a variety of environments, ranging from accretionary prisms to passive continental margins. One possible external source of low-Cl^- fluids in passive continental margins is advection of meteoric waters from the continent (e.g., Austin, Christie-Blick, Malone, et al., 1998). However, Site 1168 is located in one of a series of strike-slip basins between upthrown ridges of Cretaceous rocks on the western Tasmania margin (see "Background and Objectives"). Thus, it seems unlikely that effective conduits for meteoric recharge could be achieved at Site 1168.

Possible internal sources that may provide low-Cl^- fluids are (1) gas hydrate dissociation; (2) dehydration reactions of hydrous minerals, such as clays and biogenic opal; and (3) clay-membrane ion filtration (e.g., Kastner et al., 1991; Hesse and Harrison, 1981; Paull, Matsumoto, Wallace, et al., 1996). The distinct maxima and minima in Cl^- and Na⁺ values suggest that the emplacement of "fresher" fluids occurred relatively recently or diffusional processes would have smoothed the profiles. Gas hydrate dissociation is a common cause of such profiles in continental margin settings. Previous legs that have recovered gas hydrates have noted that the Cl^- profile returns to higher concentrations below the gas hydrate stability zone (e.g., Paull, Matsumato, Wallace, et al., 1996). The Cl^- profile at Site 1168 does not return to higher values below the estimated gas hydrate stability zone (GHSZ) or anywhere in the cored interval. Resistivity and sonic logs also do not show characteristic responses typically associated with the presence of gas hydrates (see "Downhole Measurements"). In addition, crude estimates of the base of the GHSZ, using the measured geothermal gradient (60°C/km; see "Physical Properties") and assuming a pure methane and seawater system, suggest methane hydrates should not be stable below ~300 mbsf at this site. However, the depth of the hydrate stability zone is sensitive to the chemistry of the pore fluids and incorporation of other gases into the hydrate structure (Dickens and Quinby-Hunt, 1997). Therefore, the existence of hydrates cannot yet be discounted, and a more rigorous calculation of gas hydrate stability will be pursued postcruise. With the data available, we cannot completely eliminate any of the possible hypotheses or evaluate if they are occurring in some combination; however, postcruise isotopic analyses may assist in differentiating between possible processes.

Sulfate (SO₄^2-), titration alkalinity, and ammonium (NH₄⁺) exhibit rapid changes in the upper part of the drilled interval (Fig. F32). Sulfate concentrations decrease from near-seawater values at the top of Hole 1168A to complete depletion by 238 mbsf, coincident with a sharp increase in methane (see "Volatile Hydrocarbons" in "Organic Geochemistry;" Fig. F31). Alkalinity increases sharply in the upper 200 mbsf (maximum of 9.1 mM) then decreases to 3.9 mM across the lithostratigraphic Unit I/II boundary at 350 mbsf. There appears to be real alkalinity variation at depth, but the low fluid recovery over this depth interval prevented higher resolution alkalinity measurements. Ammonium increases steadily downhole over the entire drilled sequence, reaching a maximum concentration of 2.3 mM. The pH decreases from 7.5 at the top of the hole to ~7.0 in the upper 100 mbsf then increases to 7.6 at 324 mbsf, coincident with the alkalinity minima. Below 324 mbsf, the pH begins to decrease but the measurement became limited by low water content downcore. These major changes reflected in the profiles of SO₄^2-, alkalinity, and NH₄⁺ are representative of bacterially mediated degradation of organic matter characterized primarily by sulfate reduction followed by methanogenesis (see "Volatile Hydrocarbons" in "Organic Geochemistry"). The rate of sulfate reduction appears to be high for the low sedimentation rates noted for the upper 200 mbsf (~1.5 cm/k.y; see "Age Model and Sedimentation Rates" in "Palynology").

Strontium (Sr²⁺) concentrations increase with depth from the seafloor to a maximum of 909 µM centered at ~180 mbsf, below which values decrease to a minimum of 108 µM at the base of the hole. The calcium (Ca²⁺) profile is nonsteady state, showing several maxima and minima. Calcium concentrations remain near seawater values for the upper 78 mbsf. Calcium concentration maxima are located at 123 mbsf (11.7 mM) and 468 mbsf (20.1 mM; 1.9 times seawater concentration). Minima in the Ca²⁺ profile are located at 238 mbsf (9.7 mM) and the base of the hole (866 mbsf; 5.8 mM, a 45% decrease from seawater values). Lithium (Li⁺) concentrations steadily increase with depth, reaching 890 µM at 554 mbsf, below which values decrease to 277 µM at the base of the cored sequence. It is important to appreciate that the previously discussed pore-fluid freshening exhibited by the decrease in Cl^- should also affect the ions discussed in this section. Thus, Ca²⁺, Li⁺, and Sr²⁺ have been normalized to Cl^- in Figure F33 to remove any dilution effect, yet still show nonsteady state trends.

The pronounced increase in Sr²⁺ centered around 200 mbsf is consistent with the release of dissolved Sr²⁺ to pore fluids during the recrystallization of pelagic carbonates (Manheim and Sayles, 1974; Baker et al., 1982). The largest increase in Sr²⁺ is in lithostratigraphic Unit I, which is dominated by calcareous oozes (see "Lithostratigraphy"). Typically in deep-sea marine pore waters, Sr²⁺ reaches a maximum value at depth, maintaining a plateau concentration that is ultimately limited by celestite solubility (e.g., Baker and Bloomer, 1988). However, the rapid Sr²⁺ decrease with depth at Site 1168 indicates that either carbonate recrystallization is confined solely to the interval of elevated Sr²⁺ or a Sr²⁺ sink exists at depth.

The Ca²⁺ profile appears to be controlled by dissolution (the Ca²⁺ maxima) and precipitation (the Ca²⁺ minima) reactions. The Ca²⁺ maxima at 123 mbsf and minima at 238 mbsf are coincident with the zone of elevated Sr²⁺ values in the calcareous-rich sediments of lithostratigraphic Unit I and, therefore, may be the result of local carbonate dissolution and precipitation. The maximum concentration of Ca²⁺ at 468 mbsf corresponds with lithostratigraphic Unit II in a zone of low carbonate content (Fig. F34) and may be the result of intense dissolution of CaCO₃ in this interval. However, elevated Sr²⁺ concentrations, which usually accompany increases in dissolution of CaCO₃, are conspicuously absent here. The peak Ca²⁺ concentration also corresponds to the interval of increased Li⁺ concentration. Increases in Li⁺ with depth are often attributed to alteration of biogenic carbonate, biogenic opal, or other silicate phases, including ion-exchange reactions (Gieskes, 1983; DeCarlo, 1992); however with the present data set, we cannot distinguish the source of Ca²⁺ and Li⁺ input to the pore fluids centered around 500 mbsf.

Magnesium (Mg²⁺) and potassium (K⁺) concentrations gradually and consistently decrease downcore (Fig. F32). By the base of the hole, Mg²⁺ has decreased by 74.6% to 13.7 mM, whereas K⁺ has decreased 64% to 3.8 mM. Normalizing the profiles to Cl^- (Fig. F33) shows these decreases are not attributable to dilution by the low-Cl^- source. The smooth diffusional profiles indicate that Mg²⁺ and K⁺ are being consumed in reactions below the cored interval. The decrease in Mg²⁺ and K⁺ are highly correlated (r = 0.986), which suggest they are removed by a similar process, probably associated with silicate reaction (i.e., clay diagenesis or alteration of basement) (e.g., Gieskes, 1983).

Dissolved silica (H₄SiO₄⁰) concentrations range from ~250 µM to a maximum of ~800 µM, while showing distinct changes downcore. The highest concentrations of H₄SiO₄⁰ (>600 µM) are found between ~100 to 230 mbsf and ~400 to 500 mbsf; concentrations of ~400-600 µM were observed at 20 to 100 mbsf, 550 to 700 mbsf, and 750 to 800 mbsf. The rest of the hole has H₄SiO₄⁰ concentrations below ~400 µM.

Smear-slide assessments of siliceous sponge spicule abundance (see "Biostratigraphy") within the sediments at Site 1168 (>63 µM) covary with the pore fluid H₄SiO₄⁰ variations (Fig. F35). This observation suggests that most silica dissolution occurs in sediments enriched in biogenic silica. The decreases in dissolved silica may be associated with primary depositional fluctuations or to diagenetic transformation of biogenic silica to opal-CT at depth (e.g., Baker, 1986). Further studies (e.g., X-ray diffraction analyses) may provide additional insight into such "silica diagenetic fronts."