Sediments recovered during Leg 188, like those from most drilling legs, have been altered by diagenetic processes resulting from fluid circulation, geothermally induced changes, and other factors, some of which may be related to paleoclimate and paleoceanographic changes (e.g., changes in biogenic productivity, water temperature, and ocean chemistry). Leg 188 cores show commonly observed diagenetic effects (e.g., sediment compaction and microbial generation of methane) as well as diagenetic effects found only in sections with high biogenic productivity (e.g., silica dissolution and diagenesis, dissolution of magnetite, and precipitation of carbonates). Diagenesis in the cores is a secondary effect not unique to the high-latitude environment, but in the absence of a direct record (now altered), the diagenetic record gives further evidence of changes in paleoenvironments (e.g., Quilty, this volume) that at times had relatively large silica biogenic productivity common to cold-water settings.
Primary silica is noted at all Leg 188 drill sites in the form of biogenic material (e.g., diatoms, radiolarians, and sponge spicules) and magmatic products (e.g., quartz sand). Here, we focus on the biogenic component because it is more sensitive to dissolution and reprecipitation than the magmatic component.
Continental rise sediments at Site 1165 contain relatively large concentrations of biogenic silica, as much as 30% diatoms in smear slides, whereas sediments from the shelf and slope Sites 1166 and 1167 contain relatively low concentrations or none, reflecting differences in productivity across the continental margin (Shipboard Scientific Party, 2001a). Diagenesis of silica (i.e., dissolution and reprecipitation) is best documented at Site 1165 by the downhole transition from sediments with diatoms (opal-A) above 606 mbsf to sediments with silicified horizons (inferred opal-CT and/or chert) and calcified layers below this depth (Shipboard Scientific Party, 2001a). From studies in other parts of the world where similar downsection silica phase transitions are documented (e.g., Bering Sea, Sea of Japan, and offshore California [USA]) (Hein and Obradovic, 1989), the depth to the transition depends on subsurface temperatures and relative concentrations of biogenic and terrigeneous material, with higher temperatures and larger biogenic silica concentrations facilitating the transition.
At Site 1165, the full silica transition may take place over a 150-m-thick depth zone, which is marked by the first appearance of chert at 492 mbsf, loss of all diatoms below 606 mbsf, abrupt shifts in downhole logging measurements (e.g., velocity, density, resistivity, and porosity) at 610 mbsf, and a zone of low recovery that extended to ~650 mbsf with alternating soft (clay/silt) and hard sediment layers. (Shipboard Scientific Party, 2001b; Williams et al., 2002). The hard layers are likely a combination of chert, based on common chert fragments in sediment residues from below ~598 mbsf (Quilty, this volume), calcified horizons, based on rock samples and downhole logging (Shipboard Scientific Party, 2001b), and opal-CT, inferred from silica transitions in other areas noted above.
In seismic reflection data across Site 1165, the full silica transition zone is denoted by a distinct 150-ms-thick band of high-amplitude reflections (Fig. F13). The band can be traced across the Prydz Bay continental rise. Within the band is a strong reflection that results from the abrupt change at 606 mbsf and appears to be a strong bottom-simulating reflection (BSR) like diagenetic BSRs (i.e., different from gas hydrate BSRs) seen elsewhere, such as off the Antarctic Peninsula, where the BSR and inferred diagenetic boundary is the décollement surface for slumps and slides (Volpi et al., 2003). The above diagenetic evidence points to silica microfossil concentrations (likely mostly diatoms) being sufficiently large to result in creation of the thin hard silicious layers by dissolution and reprecipitation. Early Miocene depositional environments were likely marked by periods of high biogenic silica production/deposition during cooler periods and/or periods of lower terrigeneous supply. The record of younger paleoenvironments comes from the unaltered cores above the silica transition.
The diagenetic effects of high silica concentrations in sediments and pore waters at Site 1165 are recorded also by the dissolution of magnetite and reduction of magnetic susceptibility within the upper part of the sedimentary section (94–362 mbsf) (Fig. F14A). This was noted initially by the Shipboard Scientific Party (2001b) and explained by Florindo et al. (2003b). The effect, although not restricted to polar environments, is yet another indicator of the silica-enhanced and "open circulation" sedimentary section with a nearby magnetite-rich sediment source that characterizes the continental rise at this location. The combination indicates adequate nearby sources of both terrigenous and biogenic source materials to further explain the cyclic facies patterns thought to be due to ice sheet fluctuations (Shipboard Scientific Party, 2001a).
Silica diagenesis, although important at Site 1165, is not recognized as a factor at either Site 1166 (shelf) or Site 1167 (slope). Silicious microfossil concentrations are relatively small and localized to fine-grained marine sections at Site 1166, and silicious microfossils are found only in the upper 5 mbsf at Site 1167 (Shipboard Scientific Party, 2001c, 2001d). The low concentrations and small values for silica in pore waters point to silica dissolution or nondeposition on the slope, possibly due to extended sea-ice cover (Shipboard Scientific Party, 2001d).
Carbon-related diagenesis in Leg 188 cores is seen in recycled carbon (e.g., coal) and in secondary features resulting from in situ diagenesis (e.g., authigenic carbonate nodules). The diagenesis results from processes of different types characteristic of the varied paleoenvironments of the Prydz Bay margin, ranging from relatively cold deep-ocean to relatively warm lagoonal.
Organic carbon concentrations in Leg 188 rocks vary from <0.1 wt% in hemipelagic deposits on the continental rise to a maximum of 9.2 wt% in lagoonal deposits on the continental shelf; there is evidence of biogenic and recycled terrestrial carbon at all drill sites (Shipboard Scientific Party, 2001b, 2001c, 2001d). The carbon concentrations and subsurface temperatures directly affect the diagenetic processes related to microbial degradation, methanogenesis, and, in places, diagenetic formation of authigenic carbonates (Claypool et al., this volume). The carbon and oxygen isotopic compositions of the authigenic carbonates indicate the subsurface conditions when this carbonate formed. Claypool et al. (this volume) report such stable isotopic measurements for authigenic carbonate nodules from Sites 1165 and 1166.
Comparison of the stable isotopic measurements with pore water isotopic values for the cores at these sites gives evidence of prior depositional environments that are quite different. The siderite nodules from lagoonal carbonaceous shales at Site 1166 (shelf) have carbon and oxygen isotopic compositions consistent with nodule growth during early stages of methanogenesis. The authigenic siderite from Site 1165 (rise) apparently grew in or just below the sulfate reduction zone that now extends to 150 mbsf but in the early Miocene, when sedimentation rates were 8–10 times greater, likely extended only to 10–20 mbsf. Similarities in 
13C depth profiles for dissolved inorganic carbon and authigenic carbonates (Fig. F14B) and the volume percent of carbonate in these carbonates are consistent with authigenic carbonate formation within the upper part of a methane-charged sediment drift, from dissolved carbonate derived in part from anaerobic methane oxidation.
Gas hydrates were anticipated at Site 1165 but not found, likely because of currently unsuitable conditions resulting from changes in paleodepositional environments that altered sedimentation rates. Claypool et al. (this volume) suggest that as sedimentation rates decreased since the early Miocene the sulfate reduction zone thickened, thereby destabilizing methane hydrate due to lowered concentrations of dissolved methane. They also note that some authigenic carbonate nodules have high 18O values that could result from high 18O concentrations in pore waters due to decomposition of methane hydrate at the time of nodule formation. Hence, hydrates may at one time have existed at Site 1165, but are now not found here.
