A primary objective of Leg 188 drilling was to decipher the glacial history of the East Antarctic Ice Sheet in the Prydz Bay region, which is considered a drainage area for the earliest Cenozoic ice sheets. Leg 119 drilling in Prydz Bay (Barron, Larsen, et al., 1989, 1991) established that a grounded glacier covered at least part of Prydz Bay (Sites 739 and 742) during early Oligocene. Leg 188 reconfirmed the Leg 119 assertion and acquired the data to show that the transition from nonglacial to glacial paleoenvironments occurred in the late Eocene or earlier. Further, Leg 188 drilling recovered lithologic and biostratigraphic data to help document the morphologic evolution of the Prydz Bay continental margin, which resulted from Cenozoic glaciation.
Until the early Oligocene, the Prydz Bay mid-continental shelf around Sites 1166 and 742 was part of a subaerial alluvial plain that filled the Lambert rift system including the Prydz Bay Basin. Seismic facies and links to core data support this concept (Erohina et al., this volume; Handwerger et al., this volume). Cretaceous sediments were recovered at Sites 741 and 1166 in Prydz Bay (Fig. F5). Site 741 yielded alluvial sediments consisting of channel sandstone beds, floodplain siltstone and claystone, and sporadic coal beds (Turner and Padley, 1991). Palynomorphs indicate an Early Cretaceous (?Albian) age for these units (Truswell, 1991). Black to gray carbonaceous claystone and siltstone were recovered from lithostratigraphic Units IV and V at Site 1166, giving a Turonian–?Santonian microflora (Phyllocladidites mawsonii Zone) (Macphail and Truswell, this volume a) (Fig. F5). Geophysical logs suggest some interbedded sandstone units. The combination of lithologies, palynomorphs, and organic matter (Claypool et al., this volume) indicates deposition in an alluvial plain to marginal marine lagoonal setting. Mesozoic sediments intersected at Site 1166 contain Cretaceous (Turonian) palynofloras dominated by gymnosperms with rare dinoflagellate cysts, indicating a coastal plain environment with a conifer woodland and rare freshwater swamps analogous to vegetation near the northern limit of tree growth in the modern Arctic (Macphail and Truswell, this volume a). The fossil assemblage implies year-round moderate humidity and temperatures. There is no evidence of Cretaceous glaciation from Leg 188 cores.
The only other late Mesozoic to early Paleogene rocks recovered offshore in the region are core samples containing recycled palynofloras from the outer continental shelf near Mac.Robertson Land. These samples indicate nonmarine sedimentation on this part of the margin from the Jurassic to the Early Cretaceous (Aptian?) (Truswell et al., 1999), predating Prydz Bay ODP sections.
At Site 1166, an unconformity separates the Cretaceous and middle–late Eocene units (Fig. F5). On seismic data, this unconformity shows almost no relief and only occasional truncation (Erohina et al., this volume) but nonetheless represents a break of at least 47 m.y. (top of the P. mawsonii Zone to the base of the Nothofagidites asperus Zone) (Young and Laurie 1996). Sediments deposited during this hiatus may be preserved elsewhere in Prydz Bay and the Lambert Graben because Quilty (this volume) and Quilty et al. (1999) report that fragments of the Cretaceous bivalve Inoceramus are reworked into Cenozoic sediments. Stilwell et al. (2002) reported whole Cretaceous bivalves in boulders within Cenozoic glaciomarine sediments near Beaver Lake in the northern Prince Charles Mountains. These fragments indicate marine conditions in the Lambert Graben at some stage during Cretaceous.
In the middle to late Eocene, alluvial or deltaic sediments were again deposited (Fig. F5) (Site 1166 Unit III of Shipboard Scientific Party, 2001c; Unit PS2A2 of Erohina et al., this volume). Their texture, downhole logging signature, and distribution suggests the upper part of the unit was deposited mostly in fluvial channels and also in flood or tidal basins or lagoons (Fig. F5). The lowest part of the alluvial-deltaic unit contains contorted fragments of carbonaceous sands that have yielded Cretaceous palynomorphs. These fragments are likely remnants of fluvial bank collapse blocks. The unit contains a middle to late Eocene palynoflora that suggests the Prydz Bay coastal plain was covered by a low-growing scrub of gymnosperms and angiosperms (Nothofagus) analogous to stunted rainforest scrub vegetation presently growing near the alpine limits of similar floras in Tasmania and Patagonia (Macphail and Truswell, this volume a). The flora implies humid conditions and probably microthermal (cool–cold) temperatures at sea level. More precise temperature estimates are not possible because the major components of this flora can tolerate a wide range of temperatures. With the shift of vegetation to cooler-climate species in middle to late Eocene, the first lithologic evidence for nearby glaciation in the rift-flank mountains to the south is seen in sand grain surface textures in late Eocene massive sand units (Strand et al., 2003).
In the late Eocene to early Oligocene the central part of Prydz Bay (e.g., Sites 739, 742, and 1166) permanently shifted from a fluvio-deltaic complex with rainforest scrub to an exclusively marine continental shelf environment. The massive fluvial sands of Unit III at Site 1166 grade upward into alternating sand–clay layers in Unit II (Fig. F5) with increasing marine dinoflagellates (Macphail and Truswell, this volume a) and signal a change to tidal influence and a relative subsidence of the area to near sea level. The alluvial or delta plain sediments (Site 1166; Unit III) are cut by an erosion surface with relief of ~50 m in places (Erohina et al., this volume). Glaciomarine mudstone and sandstone containing diatoms, dinoflagellates, and lonestones onlap the erosion surface in the area around Site 1166. The glaciomarine unit (Unit II of Shipboard Scientific Party, 2001c; Unit PS.2A1 of Erohina et al., this volume) is onlapped by glaciomarine diamictites encountered at Site 742 (Erohina et al., this volume; Hambrey et al., 1991).
The succession of Paleogene events in this part of Prydz Bay are as follows:
The available paleontological evidence points to these events occurring around the Eocene/Oligocene boundary. The palynomorphs within the Paleogene alluvial and glaciomarine units at Site 1166 (Units III and II) (Fig. F5) bracket the ages of 33.9 and 39.1 Ma, based on age estimates of the N. asperus Zone in eastern Australia (Macphail and Truswell, this volume a; Young and Laurie, 1996). Diatoms yield age estimates of 33–37 Ma (Shipboard Scientific Party, 2001c). Paleontological evidence at Sites 739 and 742 yield similar ages; however, Sr isotope ages on shell fragments from Site 739 yield ages of 22.7–29.3 Ma (Thierstein et al., 1991). Lavelle (2000) reported revised Sr dates for these shells of ~33 Ma. Such an age would resolve the discrepancy; however, his brief report has not been followed by full publication.
Rocks from times between the early Oligocene and early Miocene were not recovered during ODP Prydz Bay drilling but were cored during ODP Legs 119 and 120 on the 1000-km distant Kerguelen Plateau. There, principally pelagic marine units were sampled (Barron, Larsen, et al., 1989; Wise, Schlich, et al., 1989). Seismic reflection data (e.g., Fig. F3) and modeling provide the only evidence that the early Oligocene continental shelf continued to prograde and gradually deepen to normal continental shelf water depths as a result of shelf erosion, increasing ice loading onshore, and sediment loading of the continental slope (e.g., ten Brink et al., 1995). During this period, sea level changes may have resulted in periodic subaerial erosion of the shelf with sediments being carried down canyons on the slope into widespread channel-levy systems and drift deposits of the continental rise (e.g., Kuvaas and Leitchenkov, 1992; Cooper et al., 2001).
The complete section of drift deposits to the basal regional reflector P3 (which is the same as reflector P2 of Kuvaas and Leitchenkov, 1992) was drilled at Site 1165 (Handwerger et al., this volume). The Site 1165 cores establish that the early to middle Miocene was a time of abundant terrigenous sedimentation on the continental rise, with sediments rapidly building the Wild Drift on the flank of the long-established Wild Canyon system, which originates from near the western side of Prydz Bay (Figs. F1, F6). The relative amounts of downslope and along-slope sediment transport cannot be established unequivocally from drill cores. But regional sediment distribution patterns from seismic reflection data (e.g., Fig. F3) show that the continental slope and rise are underlain by thick sediment bodies and large and extensive canyon systems that have been in the same place throughout most of the Cenozoic; these observations indicate active nearby sediment source(s). The Lambert Graben is a large, long-lived Paleozoic and younger rift valley that is now, and likely has been, a conduit for a large drainage basin—further indication that Site 1165 sediments were likely derived principally from onshore Prydz Bay.
Studies in the Lambert rift-flank Prince Charles Mountains show that polythermal glaciers existed in at least late Oligocene to early Miocene and likely provided abundant suspended sediments to the shelf areas and beyond (Hambrey and McKelvey, 2000a, 2000b). The biogenic–terrigenous cyclicity in the continental rise sediments at Site 1165 in the early to middle Miocene may indeed record the variable terrigenous sediment supply from onshore glacier fluctuations (Shipboard Scientific Party, 2001a; Rebesco, this volume). The palynological record for Site 1165 indicates that southern beech vegetation (Nothofagus) may have existed nearby into early Miocene (Macphail and Truswell, this volume b), suggesting that climates and glacial conditions were more temperate than today. Pospichal (this volume) reports scattered warmer-water nannoplankton and thin chalk horizons through the Miocene at Site 1165, suggesting the intrusion of warmer waters into the area.
The continental shelf, which before the middle Miocene prograded principally by sediment bypass and erosion of the shelf's Paleogene topset strata, began to aggrade by deposition of topset strata that formed glacial banks. The aggrading initiated as the shelf subsided more rapidly than before under the increasing inner-shelf erosion, greater flexural loading by widespread onshore ice build up, and glacially derived slope sediments (e.g., ten Brink et al., 1995) (Fig. F3). The bypass sediments are those on the continental rise in lithostratigraphic Unit III (999–308 meters below sea floor [mbsf]) at Site 1165. This is based on identification of Permian, Jurassic, Cretaceous, and Paleogene recycled pollen and spores from samples below 807 mbsf (Macphail and Truswell, this volume b) and on the rare occurrence of recycled phytoliths (291–601 mbsf) that are comparable to those in modern and equivalent older trees/shrubs, grass, and ferns (Thorn, this volume).
On the rise the early to middle Miocene sediments of Site 1165 Unit III are predominantly fine-grained mudstones with a high proportion of quartz and detrital feldspar, probably in the silt size fraction (Shipboard Scientific Party, 2001b). The main clay minerals in samples are illite and kaolinite with minor chlorite. Mechanical weathering of igneous and metamorphic rocks typically delivers illite to sedimentary basins (Ehrmann et al., 2003), so is not unexpected in sediments adjacent to a glaciated landmass. Kaolinite, however, is typically a product of chemical weathering. Its presence in continental rise sediments off Prydz Bay that postdate the initial early Oligocene glaciation suggests that either kaolinite was recycled from older sedimentary basins on the shelf or that chemical weathering continued onshore through the early Miocene.
The input from shelf sedimentary basins to the continental rise in response to climatic change, ice buildup, and increased erosion becomes clear in the middle Miocene section at the distinct boundary between lithostratigraphic Units III and II at 307 mbsf at Site 1165 (Fig. F6). Total clay minerals increase as detrital plagioclase decreases, diatoms become more common, and the number of lonestones also increases (O'Brien, Cooper, Richter, et al., 2001). Also, glauconite grains, likely from the shelf and reworked pseudomorphs and tests of Paleogene benthic foraminifers, appear in the sediments (Quilty, this volume). There is further evidence of climatic/depositional change at this time from the abrupt uphole disappearance of sponge spicules and orosphaerid radiolarians that occur from 290 to 520 mbsf (Quilty, this volume). The above transitions in conjunction with multiple strong unconformities in seismic profiles point to erosion of the outer shelf in the Prydz Bay and Mac.Robertson areas and, thus, a major expansion of ice.
The depositional systems of the continental margin began to change dramatically at about the middle Miocene (i.e., 14–16 Ma; ~300–400 mbsf), coincident with progressively declining sedimentation rates and other lithologic changes at Site 1165 (Fig. F6). The morphologic inferences come principally from regional seismic data (O'Brien et al., this volume) and model studies (e.g., ten Brink et al., 1995) and are supported by the new drilling data, especially those from Site 1165. At about middle Miocene time, depocenters began to shift landward from the continental rise and beyond to the continental slope and base of slope (Fig. F3). The shift coincided with increases in clays, diatom content, and ice-rafted debris (IRD) at Site 1165 at ~300 mbsf (Shipboard Scientific Party, 2001b)—a shift that points to coarser sediment components being deposited elsewhere (e.g., landward of the rise) and to more icebergs crossing the site.
In the middle–late Miocene, shelf and slope progradation increased and shifted toward the middle of Prydz Bay rather than the earlier even distribution along the shelf edge (Fig. F3) (Cooper et al., 2001). These shifts in the geometry of the margin coincide on the rise at Site 1165 with the first occurrence of ?recycled glauconite at 213 mbsf (Quilty, this volume), notable decreases in grain density, and porosity at ~200 mbsf and with additional uphole increases in IRD concentrations above ~150 mbsf (Shipboard Scientific Party, 2001b). Other more subtle changes are noted in the Site 1165 cores over this interval (e.g., 300–150 mbsf) and include increased organic carbon, increased color reflectance, higher silica in pore waters, and greatly diminished magnetite. Collectively, these variations point to changes in erosion processes, sediment-source locations, and lithologic controls on secondary circulation processes and diagenesis.
In this middle–late Miocene transition, the continental shelf was initially overdeepened by glacier erosion (e.g., La Macchia and De Santis, 2000), and many slope canyons were filled and buried, resulting in more distal continental rise areas such as Site 1165 receiving diminishing sediment (Figs. F3, F6) (O'Brien et al., this volume). Prior to the change, temperate glaciers eroded onshore and intermittently on the shelf, resulting in widespread erosion of a normal water depth shelf during lowered sea levels.
We envision that after the change, larger and (?)colder glaciers episodically extended far onto the continental shelf. Onshore, the Lambert Glacier system oscillated between warm and cold states. In the warm state, the Lambert Graben was a huge fjord with glaciers flowing from the edges and a mixture of open water, sea ice, and icebergs in the basin proper (Hambrey and McKelvey, 2000a, 2000b; Whitehead and Bohaty, this volume). In the cold state, glacial maximum glaciers in the graben merged into a major axial ice stream. The ice stream eroded troughs beneath fast-moving areas and deposited banks beneath slow-moving ice on parts of the shelf. This resulted in localized erosion in deep trough areas of the inner shelf, Mac.Robertson shelf, and Prydz Trough during times of glacier expansions in colder paleoclimates.
In the late Miocene and early Pliocene, shelf morphology was further modified strongly by focused erosion troughs and deposition banks as a direct result of glacier fluctuations. The widespread shelf erosion is documented by the early Oligocene to late Miocene hiatus at Site 739 and the early Oligocene to Pliocene unconformities documented at Sites 1166 and 742 (Barron, Larsen, et al., 1989; Shipboard Scientific Party, 2001c). Massive and overcompacted glacial diamictons that form Four Ladies Bank lie above these unconformities. The shelf banks have been sampled at Sites 739, 742, and 1166 and their subglacial origin deciphered from subglacial sand grain textures in the diamictons (Strand et al., 2003) and broad bank/trough morphologies on the shelf (O'Brien et al., 1999). Periods of reduced ice cover and open water on the Prydz Bay shelf since the late Miocene are documented at all ODP shelf drill sites by the presence of diatomaceous muddy units (e.g., Site 1166) (Whitehead and Bohaty, this volume), but such units are only a small fraction of the late Neogene recovered core.
On the rise the late Miocene and Pliocene are marked by decreasing sedimentation rates and a number of hiatuses in the section at Site 1165 (Fig. F6) (Warnke et al., this volume). The early Pliocene section (~34–50 mbsf) has low IRD concentrations and relatively low kaolinite levels (Warnke et al., this volume; Grützner, this volume). Upsection increases in kaolinite and IRD at 32 mbsf (~3.5 Ma) (Florindo et al., 2003a) are accompanied by an increase in the grain size of magnetic minerals; these increases point to enhanced detritus eroded from sedimentary basins on the shelf and basement outcrops on the shelf and onshore.
The Neogene record of terrestrial plants in the Prydz Bay region is sparse, with most palynomorphs at Sites 1165 and 1167 likely recycled (Macphail and Truswell, this volume b; Thorn, this volume). There are a few occurrences of species of the genus Coptospora plus Phyllocladites and Nothofagidities that may not be recycled and may be indicators of Miocene-age tundra-style vegetation similar to that suggested for the Ross Sea region (Raine, 1998). These pollen have no modern counterparts, so climatic interpretation is uncertain.
The last major episode of late Neogene glacial erosion and sedimentation in the Prydz Bay region created the Prydz and Svenner Channels on the shelf and the Prydz Channel Fan on the upper slope. Widespread erosion of the shelf and slope is marked by seismic reflection surface A of Mizukoshi et al. (1986), which is the same as PP-12 of O'Brien et al. (this volume). O'Brien et al. (1995) tie surface PP-12 to ODP Site 739 (105.9–130 mbsf) to infer that the surface is of early Pliocene age. Prior to the erosion of PP-12, the shelf prograded and aggraded nearly evenly across Prydz Bay, but after, sedimentation concentrated in the Prydz Channel Fan. On the shelf, post-PP-12 sediment sections are flat-lying and aggrade beneath the Four Ladies Bank, whereas in Prydz Channel coeval sediments are thin and strongly progradational (Fig. F7). Changes in the reflection geometries in Prydz Channel point to the development of an ice stream beneath which most debris was transported in a mobile, basal layer to the shelf edge. On the bank, slower-moving ice deposited more basal till.
Site 1167 is the first deep drill hole in an Antarctic trough-mouth fan, and the cores give an expanded view of glacier fluctuation history for the past 1–2 m.y. The majority of the fan comprises poorly sorted pebbly, clayey sands and diamictons deposited as debris flows; however, the upper 5 m of the fan is principally hemipelagic muds and occasional turbidites (Golding, 2000; Shipboard Scientific Party, 2001d; Passchier et al., 2003). The succession of facies is consistent with predictions based on shallow cores from tough-mouth fans found on other glaciated margins (Vorren and Laberg, 1997). The debris flow intervals are up to tens of meters thick and are separated by thin mud units and rare sand and gravel beds. The debris flows are thought to be derived from slumping of subglacial debris that melted out at the shelf edge when the Lambert Glacier grounded there, and the muds were formed by hemipelagic settling during periods when the ice had retreated from the shelf edge (Shipboard Scientific Party, 2001a).
During periods when ice retreated from the shelf edge, the fan surface was also probably eroded and reworked followed by deposition of some mud units. Such erosion can be seen by the development of erosion surfaces within the fan (O'Brien et al., this volume). Some of these surfaces can be mapped throughout the fan in seismic reflection data. Several of these surfaces intersect Site 1167, where they are close to mud intervals or changes in sediment composition, reflected in geophysical logs, mineralogy, and magnetic properties (O'Brien et al., this volume). It is not possible to definitively tie the drill core records from the slope (Site 1167) to the rise (Site 1165) because the sedimentary section thins between the sites and there is a hiatus at Site 1165 that corresponds to the lower half of Site 1167 (Warnke et al., this volume). But, as described below, cyclicities are observed in cores from the slope and rise to further suggest advances and retreats of the ice sheet.
Fan construction occurred principally between the early Pliocene and late Pleistocene, with only 34 m of sediment (<5%) younger than 780 k.y. (i.e., younger than the Brunhes/Matuyama boundary, noted in the cores) (Shipboard Scientific Party, 2001d). The upper meter of the fan has an age of 36.9 ± 3.3 ka at 0.45 mbsf (Theissen et al., 2003). The oldest sediments that could be dated using nannoplankton and strontium isotope dating are from a depth near 217 mbsf and have an age of ~1.1 Ma (M. Lavelle, pers. comm., 2001). Drilling reached 447 mbsf and did not reach the bottom of the fan (i.e., reflector PP-12) (O'Brien et al., this volume), which is thought to be early Pliocene. The fan stratigraphy shows a decreasing number of inferred muddy horizons upsection, indicating a reduction in the number of advances to the shelf edge in the late Pleistocene. In fact, there may have been as few as three advances of the ice sheet to the shelf edge in the past 780 k.y. (O'Brien et al., this volume). On the shelf, the Last Glacial Maximum grounding zone wedges of the Amery Ice Shelf are arranged around the flanks of Prydz Channel no more than 120 km north of the present edge of the Amery Ice Shelf and far from the shelf edge (Domack et al., 1998; O'Brien et al., 1999).
Site 1167 cores exhibit systematic changes in pebble composition, clay mineralogy, magnetic properties, and geophysical log readings that may also be related to ice sheet variations. At 217 mbsf, there is a pronounced uphole change from sandstone- to granite-dominated clasts (Fig. F8). This change corresponds to an abrupt decrease in smectite abundance, an increase in magnetic susceptibility and spectrophotometer values, and an increase in gamma log values (indicating more potassium-bearing minerals). These changes suggest a shift from erosion of sedimentary basins on the shelf by sedimentary pebbles, recycled clays, and organic matter, finer-grained magnetites typical of redbeds in Prydz Bay Basin (Domack et al., 1998) to erosion of offshore and onshore basement rocks. The changes occur at ~1.1 Ma at a boundary equivalent with regional seismic unconformity PP-4.
Further evidence of prior ice fluctuations comes from above 217 mbsf at Site 1167. Magnetic susceptibility values form an uncommon "sawtooth" pattern whereby values start high then fall linearly before abruptly jumping to higher values again (Fig. F9) (Scientific Party, 2001d). Though susceptibility values do not fall to the "sediment-rich" values seen below 217 mbsf, they still may be explained by progressive increases in sedimentary detritus during cycles of ice volume increase. These systematic cyclic variations, however, do not correspond exactly to depositional cycles indicated by the interbedding of muds and debris flow deposits. Rather, susceptibility cyclic variations typically occur over several phases of debris flow and mud deposition, implying a longer-term fluctuation of ice volume than that of the advance and retreat cycles indicated by the sediment facies. Another explanation is that the sawtooth trends may result from secondary alteration of magnetic minerals below unconformities that represent relatively long breaks in sedimentation. Although the sawtooth patterns cannot now be fully explained, we believe it is linked to changes in ice volume that have acted throughout late Neogene.