Sedimentation on the continental rise west of the Antarctic Peninsula over the last three glacial cycles

Marine Geology 167 (2000) 313–338 www.elsevier.nl/locate/margeo

Sedimentation on the continental rise west of the Antarctic Peninsula over the last three glacial cycles C.J. Pudsey* British Antarctic Survey, High Cross, Madingley Road, Cambridge CB3 0ET, UK Received 6 July 1999; accepted 18 March 2000

Abstract The continental rise west of the Antarctic Peninsula includes a number of large sediment mounds interpreted as contourite drifts. Cores from six sediment drifts spanning some 650 km of the margin and 4⬚ of latitude have been dated using chemical and isotopic tracers of palaeoproductivity and diatom biostratigraphy. Interglacial sedimentation rates range from 1.1 to 4.3 cm/ ka. Glacial sedimentation rates range from 1.8 to 13.5 cm/ka, and decrease from proximal to distal sites on each drift. Late Quaternary sedimentation was cyclic, with brown, biogenic, burrowed mud containing ice-rafted debris (IRD) in interglacials and grey, barren, laminated mud in glacials. Foraminiferal intervals occur in interglacial stages 5 and 7 but not in the Holocene. Processes of terrigenous sediment supply during glacial stages differed; meltwater plumes were more important in stages 2–4, turbidity currents and ice-rafting in stage 6. The terrigenous component shows compositional changes along the margin, more marked in glacials. The major oxides Al2O3 and K2O are higher in the southwest, and CaO and TiO2 higher in the northeast. There is more smectite among the clay minerals in the northeast. Magnetic susceptibility varies along and between drifts. These changes reflect source variations along the margin. Interglacial sediments show less clear trends, and their IRD was derived from a wider area. Downslope processes were dominant in glacials, but alongslope processes may have attained equal importance in interglacials. The area contrasts with the East Antarctic continental slope in the SE Weddell Sea, where ice-rafting is the dominant process and where interglacial sedimentation rates are much higher than glacial. The differences in glacial setting and margin physiography can account for these contrasts. 䉷 2000 Elsevier Science B.V. All rights reserved. Keywords: Antarctic Peninsula; Biogenic barium; Continental rise; Glacial–interglacial; Sediment drifts; Sediment geochemistry

1. Introduction The continental rise west of the Antarctic Peninsula consists of a thick clastic sedimentary succession deposited on ocean floor of Cenozoic age. Recent marine geological and geophysical studies, including Leg 178 of the Ocean Drilling Program (ODP), have focused on the area as a valuable high-resolution record of Neogene Antarctic glaciation. The Antarctic * Tel.: ⫹ 44-1223-251400; fax: ⫹ 44-1223-362616. E-mail address: [email protected] (C.J. Pudsey).

Peninsula was a magmatic arc lying above a southeast-dipping subduction zone throughout the Mesozoic and Early Cenozoic (Pankhurst, 1982; Storey and Garrett, 1985). Subduction ceased with the arrival of segments of a spreading ridge at the trench; this occurred progressively later to the northeast along the arc (Barker, 1982; Larter and Barker, 1991), so that ocean floor near the margin ranges in age from Eocene in the southwest to Pliocene in the northeast. The history of research in this area, from early surveys related to Deep Sea Drilling Project Leg 35 through to digital multichannel seismic reflection

0025-3227/00/$ - see front matter 䉷 2000 Elsevier Science B.V. All rights reserved. PII: S0025-322 7(00)00039-6

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Fig. 1. Map of the Pacific margin of the Antarctic Peninsula, showing positions of sediment drifts 1 to 8 and the cores discussed in this paper. ODP sites 1095, 1096 and 1101 also shown. Bathymetry and drift outlines from Rebesco et al. (1998) and Pudsey et al. (2000). Present-day position of Polar Front from Orsi and Whitworth, 1995. Inset shows the Polar Front and Antarctic Circumpolar Current axis (dashed line with open arrows) and the Weddell Gyre with water of Weddell Sea origin flowing southwest, west of the Antarctic Peninsula (grey arrows). Asterisk marks the study area of Grobe and Mackensen (1992), and triangle marks that of Frank et al. (1995), who also studied one core from the SE Weddell Sea.

profiles acquired mainly by Italian, UK and US groups, was summarised by Pudsey and Camerlenghi (1998). On the upper continental rise is a series of large mounds, separated from the base of the continental slope and from each other by channels carved by turbidity currents (Fig. 1; Tomlinson et al., 1992). The mounds are 100–300 km long (perpendicular to the continental margin), 50–100 km wide and are elevated several hundred metres above the surrounding seabed. They were interpreted as sediment drifts by Rebesco et al. (Rebesco et al., 1996, 1997) and the largest drifts in the southwest were divided into six depositional units on the basis of seismic stratigraphy.

The upper three units were thought to be fully glacial sediments of Upper Miocene to Pleistocene age, with accumulation rates of 14–17 cm/ka (Rebesco et al., 1997). It was inferred that the main “drift-growth” stage reflected high terrigenous supply from the glaciated continent, with fine-grained material transported downslope as turbidity currents and redistributed along the continental rise by southwest-flowing bottom currents. An alternative view by McGinnis and Hayes (1995) and McGinnis et al. (1997) considered the mounds to consist largely of overbank deposits, i.e. as giant levees to the turbidity current channels. In this paper we favour the interpretation of

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Rebesco et al. (1997) and refer to the mounds as sediment drifts (see also Pudsey and Camerlenghi, 1998). Recently, ODP Leg 178 drilled at three sites on the continental rise in early 1998 (Barker et al., 1999). ODP Sites 1095 and 1096 were located on the distal and proximal parts of Drift 7 respectively, and Site 1101 in the centre of Drift 4 (Fig. 1). In this paper I describe piston cores from six of the drifts, spanning some 4⬚ of latitude and a distance of 650 km along the margin from northeast to southwest. These cores recovered sediment from the “drift-maintenance” stage of Rebesco et al. (1997). All cores contain at least one glacial–interglacial cycle, and three cores recovered sediment dating back to isotope stage 8, on the basis of biostratigraphy, chemical and isotope stratigraphy. The distribution and timing of biogenic and terrigenous hemipelagic/contourite deposition on the continental rise provide constraints on the extent and timing of glaciation on the shelf. Bulk geochemistry is used in addition to clay mineralogy to evaluate the importance of local or distant source areas. This work on Late Quaternary environments provides a baseline for assessing glacial–interglacial cyclicity for the Quaternary to Late Miocene, from sediments cored on ODP Leg 178. Comparison with the depositional model of Grobe and Mackensen (1992) for the southeastern Weddell Sea emphasises the effects of different glacial settings and margin physiography on continental slope and rise sedimentation. 1.1. The source area: Antarctic Peninsula geology Geological maps have been published by the British Antarctic Survey (Fleming and Thomson, 1979; Thomson, 1981; Thomson and Harris, 1981, 1982; Tectonic Map of the Scotia Arc, 1985; Moyes et al., 1994). The Antarctic Peninsula consists largely of volcanic and plutonic rocks of the Mesozoic–Cenozoic magmatic arc, with fore-arc basin sedimentary rocks on Alexander Island in the south (Fig. 1). The Carboniferous to Triassic metasedimentary basement to the arc (Trinity Peninsula Group) is exposed on the northern tip of the Peninsula and part of the east coast. The age and origin of the basement of the arc is less certain in the south, but it locally includes crystalline rocks of early Palaeozoic age (Milne and Millar, 1989). Jurassic to Cretaceous accretionary prism metasediments (Le May Group) occur on Alexander

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Island. There are small areas of Miocene to Recent alkaline volcanics (Smellie, 1999) and Cretaceous to Eocene back-arc basin sediments occur on islands east of the northern tip of the Antarctic Peninsula. The bulk of the rocks exposed belong to the Antarctic Peninsula Volcanic Group or to plutons of the Antarctic Peninsula batholith (Leat et al., 1995; equivalent to the Andean Intrusive Suite of Moyes et al., 1994), and are of Middle Jurassic to Cretaceous age. The volcanics are mainly basaltic on the west coast and andesitic to rhyolitic in the east, and include a wide variety of lavas, pyroclastic and volcaniclastic rocks. The plutons also have a wide range of compositions from mafic to felsic, many individual plutons are heterogeneous, and dykes are ubiquitous (Leat et al., 1995). Petrographic and chemical variations across the arc from northwest to southeast are generally greater than along it, which hampers the search for diagnostic detrital petrographic or chemical signatures along the continental margin.

1.2. Glaciology The Antarctic Peninsula ice sheet drains both east and west from the spine of the peninsula, with flow concentrated in ice streams; it averages 500 m thick (Drewry, 1983). Much of the east coast is occupied by the Larsen Ice Shelf. On the west side the ice sheet generally terminates at the coast as ice walls or valley glacier tongues (Keys, 1990), with ice drainage basins only a few tens of km long. Floating ice shelves draining somewhat larger areas are present in the southwest around Alexander Island (Fig. 1). Rapid retreat of some of these ice shelves has occurred within the last few thousand years, continuing into historical times (last 30 years; Kellogg and Kellogg, 1987; Doake and Vaughan, 1991; Vaughan and Doake, 1996). The exposed continental shelf is some 150 km wide, offshore of the islands. Icebergs, once calved, drift in coastal currents; in the Amundsen and Bellingshausen Seas they tend to drift clockwise, i.e. south-westwards near the coast, then north (frequently grounding on shallow areas of the continental shelf) and finally east within the Antarctic Circumpolar Current (Keys, 1990). The area of the drifts is covered by seasonal sea ice for a period ranging from 9 months per year in the southwest to 5 months in the northeast (Gloersen et

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Table 1 Core sites Core numbers

Location

Water depth (m)

Recovery (m)

Comments

PC 055 PC 101 PC 102 PC 103 PC 106 PC 107 PC 108 PC 109 PC 110 PC 111 PC 113

63⬚50.0 0 S 67⬚47.9 0 W 63⬚05.7 0 S 65⬚30.8 0 W 64⬚35.2 0 S 69⬚24.8 0 W 66⬚48.4 0 S 75⬚56.8 0 W 66⬚18.8 0 S 76⬚58.7 0 W 65⬚54.0 0 S 72⬚39.9 0 W 65⬚42 0 S 73⬚38 0 W 65⬚14.5 0 S 70⬚20.0 0 W 65⬚08.8 0 S 70⬚35.3 0 W 64⬚19.0 0 S 70⬚26.2 0 W 63⬚27.3 0 S 68⬚58.0 0 W

3155 2930 2787 2941 3662 3080 3601 2729 3025 3357 3552

7.39 6.75 9.67 10.12 9.27 8.77 9.15 11.07 7.55 10.93 10.80

Drift 3, proximal Drift 1, proximal Drift 4, proximal Drift 6, proximal Drift 6, distal Drift 5, proximal Drift 5, distal Drift 4A, proximal Drift 4A, distal Drift 4, distal Drift 3, distal

al., 1992). In winter, coastal polynyas may be present (Zwally et al., 1985). 1.3. Oceanography The drifts lie some 150–300 km south of the Polar Front (Fig. 1). Summer sea surface temperature is 0⬚ to ⫹1⬚C (Olbers et al. 1992); cold Antarctic Surface Water overlies the warmer but more saline Circumpolar Deep Water (CPDW), which occupies most of the water column. The bottom water (lowest few hundred metres) in this area and in southernmost Drake Passage is colder, fresher and more oxygenated than CPDW and is thought to originate in the Weddell Sea, flowing into the southern Scotia Sea through deep gaps in the South Scotia Ridge and thence westwards close to the South Shetland Islands (Sievers and Nowlin, 1984; Nowlin and Zenk, 1988; Camerlenghi et al., 1997; Alessandro Crise, pers. comm.). Slow and steady contour-following currents were measured at two sites on Drift 7 by Camerlenghi et al. (1997). Average speeds 8 m above the seabed were about 6 cm s ⫺1, with speed rarely exceeding 14 cm s ⫺1. Tidal currents were very weak. These measured currents are too slow to erode continental rise sediment, though they can maintain fine silt and clay in suspension. 2. Methods 2.1. Site survey and coring Cruise JR19 of RRS James Clark Ross collected

3770 km of bathymetric and 3.5 kHz acoustic profiler data in the study area in March 1997. GPS-navigated tracks were designed to complement earlier surveys (Cunningham and Vanneste, 1995; Rebesco et al., 1997) and to map the edges of the drifts. Most lines were parallel or perpendicular to the continental margin (i.e. across or along the long axes of the drifts). Core sites (Table 1) were chosen in areas of smooth seabed with acoustic sub-bottom penetration of at least 30 m. One shallow (proximal, southeastern) site and one deeper, more distal site on each drift allowed “ground-truthing” of observed sediment thickness changes. The steep sides of the drifts were avoided, as hiatuses are present there (Pudsey and Camerlenghi, 1998). A Driscoll-type piston corer was used, assembled to either 9 m or 12 m length, with a 1.2 m trigger corer. 2.2. Sedimentology Visual descriptions of the cores were supplemented by preliminary estimates of grain size made from smear slides. Magnetic susceptibility was measured on the split surface of the archive half of each core at 2 cm intervals, using a Bartington susceptibility meter with MS2F probe. Magnetic susceptibility data and careful visual examination were used to splice together the record from each trigger core and the top of the corresponding piston core; it was found that 0.2–0.6 m of sediment was either missing from, or very disturbed in, some of the piston cores. In core 109, approximately 2 m of sediment was lost from one core barrel during recovery. Biogenic silica (in the

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form of diatoms, radiolarians and silicoflagellates) was measured by point-counting smear slides. This method tends to overestimate silica compared with the true weight % (Pudsey, 1993). Total organic carbon (TOC) and total inorganic carbon (TIC) were measured in a LECO CS125 induction furnace at Robertson Laboratories, North Wales; calcium carbonate % was calculated from TIC. Sand % was measured by wet-sieving at 63 mm (4f ), and selected sand fractions coarser than 500 mm were examined in reflected light. Major and trace element X-ray fluorescence analyses were carried out at Geoscience Analytical Services, University of Keele. Sample preparation included a large-volume dilution to reduce the salt content of the sediment, followed by removal of grains larger than 500 mm, drying and crushing to pass a 120 mm sieve. X-ray diffraction analysis of clay minerals (finer than 2 mm) was also performed at Geoscience Analytical Services (water-settled mounts, CuKa radiation, scan range 3–32⬚ 2u , with a slow scan from 24–28⬚ 2u to check for the presence of kaolinite). Peak areas on the diffractograms were compared to give semiquantitative abundances of the main clay minerals (Biscaye, 1965; Petschick et al., 1996). 2.3. Age Diatoms were examined in smear slides, and this proved sufficient for an initial zonation. Calcareous nannofossils were also noted in smear slides, but even under scanning electron microscopy they were identifiable only to genus level because of dissolution. Foraminifera and radiolarians are present only in certain intervals of the cores, so unfortunately it was not possible to derive either a continuous oxygen isotope curve, or data on Cycladophora davisiana for radiolarian abundance stratigraphy (Morley and Hays, 1979). The trace element barium was used as a palaeoproductivity indicator to delineate warm climatic stages, following Shimmield et al. (1994), Frank et al. (1995), Nu¨rnberg et al. (1997), Bonn et al. (1998), and Pudsey and Howe (1998). These authors demonstrated that, in the Weddell Sea, Scotia Sea, Ross Sea and southeast Atlantic, biogenic barium was consistently high in warm isotope stages 1, 5e, 7 and 9, and low during

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cold stages 2 to 4, 6, 8 and 10. A correction must be made for lithogenic barium (Dymond et al. 1992), and the magnitude of this correction depends upon source rock composition. Uranium and thorium isotope analysis was performed at the Heidelberger Akademie der Wissenschaften. Following chemical preparation techniques given by Frank et al. (1994), Th and U activities were measured by alpha-spectrometry. Excess 230Th activity (originating from the water column) was calculated by subtracting the amount of 230Th which is in equilibrium with 234U for each sample (Frank et al., 1995). 3. Results 3.1. Acoustic profiles on the drifts The 3.5 kHz profiles on all drifts show generally deep penetration and parallel, continuous sub-bottom reflectors. An acoustic facies map, with examples of different echo types, was presented by Pudsey et al. (2000). Along the drifts (proximal to distal, southeast to northwest) sediment layers between reflectors become thinner offshore (Fig. 2). Across the drifts from northeast to southwest, there are no consistent thickness changes except within a few km of the channels. Earlier work by Camerlenghi et al. (1997) and Pudsey and Camerlenghi (1998) revealed little sedimentological variation from southwest to northeast on Drift 7 except on its steep margins. Both these results agree with observations from multichannel seismic profiles (at much lower resolution) on Drifts 6 and 7 (Rebesco et al., 1997), suggesting similar processes have affected all the drifts. Drift 1 is less elevated above the surrounding seafloor than the other large drifts, and sediments within it are partially banked around seamounts and have a more undulating geometry than in the drifts farther south. This is a consequence of the proximity of Drift 1 to the Antarctic Circumpolar Current (Fig. 1, inset). 3.2. Lithostratigraphy The cores are mainly fine-grained and show a clear cyclicity in colour, composition, occurrence of icerafted debris (IRD) and sedimentary structures. Thick, grey, terrigenous laminated units alternate

318 C.J. Pudsey / Marine Geology 167 (2000) 313–338 Fig. 2. Examples of 3.5 kHz profiles; positions in Fig. 1. (A) Dip line on Drift 4A (cores 109 and 110), showing offshore thinning of sediment layers between reflectors. (B) Strike line through core 109. This shows no significant thickness change across the drift. The channels between the drifts are a few km wide and 80–130 m deep.

C.J. Pudsey / Marine Geology 167 (2000) 313–338 Fig. 3. Core logs showing cyclicity in sedimentary structures (lamination or bioturbation) and occurrence of diatoms and foraminifera. In the column for sediment colour, brown includes greyish brown (Munsell colour 2.5Y 5/2), dark greyish brown (2.5Y 4/2), and olive grey (5Y 5/2, 5Y 4/2); grey includes greenish grey (5GY 5/1), dark greenish grey (5Y 4/1, dark grey (5Y 4/1) and very dark grey (5Y 3/1), with a few thin beds of very dark grey (2.5Y 4/1) which appear reddish by comparison. Lithological units A, B, C, D and E are also shown. 319

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110, 111 and 113 contain thin, dark brown, bioturbated layers of manganese micronodules with small Chondrites burrows. Away from these layers the burrows are of Planolites type, with a few halo burrows up to 2 cm in diameter. There is no systematic thickness change in unit A, either from proximal to distal on each drift or from northeast to southwest along the margin.

Fig. 4. Magnetic susceptibility profiles for cores 106 (distal) and 103 (proximal) on Drift 6. Values were measured every 2 cm and plotted as a 5-point moving average. The raw data exhibit single-point spikes corresponding to silt laminae or dropstones. These lines mark the suggested correlation between these two cores. Not only do these cores correlate clearly with each other, but they also match cores on Drift 7 (Pudsey and Camerlenghi, 1998). Units A–E from Fig. 3.

with thinner, brown, diatom- and foraminifer-bearing units which are bioturbated and contain dispersed sand grains and granules. There is no evidence for significant hiatuses in the form of sharp lithological changes or increases in sediment firmness (in contrast to cores SED-01, 05 and 09 from the steep sides of Drift 7; Pudsey and Camerlenghi, 1998). The succession is similar in each core, so the lithostratigraphy for the whole area will be described from the top down. Core logs are shown in Fig. 3. The units A to D corresponds to those described from Drift 7 by Pudsey and Camerlenghi (1998), emphasising the uniformity of sedimentation along the margin. 3.2.1. Unit A At the top of each core there is 0.1–0.2 m of dark greyish brown diatom-bearing clayey mud which appears to be structureless, probably because of complete bioturbation. This overlies, with a gradational burrowed contact, 0.3–1.0 m of variegated, very thin-bedded to laminated clayey mud with a few % diatoms, showing minor bioturbation. Colours include greyish brown, olive grey and dark grey and the lamination tends to be indistinct. Cores 107, 109,

3.2.2. Unit B The top of unit B is a sharp colour change to dark greenish grey or dark grey clayey mud. In most cores this coincides with the downward disappearance of bioturbation (Fig. 3) except for rare, tiny beddingparallel burrows. Unit B is 2.3–7.7 m thick, uniform in colour and faintly but pervasively laminated in most cores, though the lamination may be visible as only very faint colour contrasts. Laminae to very thin beds of a slightly redder hue occur in cores 103 and 106 on Drift 6, but not farther north. Silt laminae 1– 10 mm thick (rarely up to 30 mm) are present in unit B in cores on the three northern drifts. They have sharp bases, sharp or gradational tops and the thicker silts are normally graded. Normal microfaults occur in some cores, especially core 110. IRD is rare in unit B, though there are thin sandy lags in cores 101, 106 and 109 and isolated dropstones in some of the other cores. The lower 1–2 m of the unit contains no visible lamination in cores 103 and 106. On each drift, unit B thins offshore; this is most marked on Drift 4A where it thins from 7.7 m (core 109) to 2.3 m (core 110, only 16 km to the northwest). 3.2.3. Unit C The top of unit C is sharp or gradational, designated at the first marked downcore increase of biogenic material, either silica or carbonate. In most cores this is coincident with a colour change from dark grey to brown, though in cores 102, 111 and 113 the colour change occurs slightly lower down. Unit C is rather different in the proximal and distal cores. It comprises 0.5–2.0 m of diatom-bearing clayey or silty mud, dark greyish brown to olive brown in colour and homogenised by bioturbation. In the distal cores the upper part of the unit is foraminifer-bearing and contains up to 1% calcareous nannofossils, over a thickness of about 0.7 m (see below); in the proximal cores this upper part is reduced to a few cm of

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Fig. 5. Biogenic silica (open symbols; measured as area % in smear slides) and calcium carbonate percentages (crosses, measured as weight %). For cores 103, 107, 109 and 113, inspection of smear slides for those intervals not measured shows that CaCO3% is very low.

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foraminiferal sand, and diatom content in the unit is relatively low. Unit C is rich in dispersed ice-rafted sand grains and granules (sand content up to 10%, including 2–3% coarser than 500 m); towards the base, radiolarians are conspicuous in the sand fraction. Near the base of unit C is a fresh vitric ash up to 4 cm thick, of alkaline composition (Pudsey and Camerlenghi, 1998; Moreton, 1999). 3.2.4. Unit D The top of unit D is placed at an abrupt decrease in biogenic silica; note that this does not generally correspond to a colour change to grey. In core 101 only the upper 0.6 m of unit C is brown, and in many of the other cores olive grey mud extends 0.5–1.0 m below the diatom-bearing unit. Total thickness is 2.1–3.0 m where completely recovered, and at least 3.8 m in core 111; the base was not recovered at any of the proximal sites. Unit D consists of laminated clayey or silty mud with a very low biogenic content, with minor bioturbation in the upper part in cores 101, 106, 110 and 111. Graded, sharp-based silt laminae and very thin fine sand beds are common, exhibiting parallel and cross-lamination. In cores 101 and 113 unit D contains dispersed sand grains throughout, and sand-granule lags occur in core 106 (Fig. 3). 3.2.5. Unit E Unit E was recovered in four cores (106, 108, 110, 113) and its top is at a downcore increase in biogenic silica and carbonate. Dark greyish brown diatombearing or foraminifer-bearing mud is 0.7–1.6 m thick and contains 5–15% sand. Bioturbation is extensive, though in core 113 foraminifera occur in small patches and laminae rather than dispersed uniformly. Below unit E in cores 108 and 110, diatom content decreases in the lowest 0.2–0.3 m of the core; in core 113 the lowest 0.8 m shows further alternations of

diatom-bearing and barren mud. No more lithological units are defined on this limited evidence. 3.3. Magnetic susceptibility Magnetic susceptibility data provide a means of detailed correlation between cores on each drift (Fig. 4; Pudsey and Camerlenghi, 1998). Cyclicity corresponding to the lithological units described above can be seen on all drifts, with units A, C and E having low susceptibility. Unit B shows a weak along-margin trend of lower values to the southwest. While the low susceptibility values in units A, C and E may be ascribed to biogenic dilution of magnetic minerals, the high and variable values in units B and D are likely to reflect variations in mineralogy and grain size, which have not yet been studied in detail. 3.4. Biogenic silica, carbonate and TOC The distribution of biogenic silica (mainly diatoms, with some radiolarians and rare silicoflagellates) is markedly cyclic, with peaks in units A, C and E and low values in units B and D (Fig. 5). In many cores the highest values are at the core top, though in cores 101, 106, 109, 110 and 113 highest values occur in unit C. Even in the terrigenous units B and D few samples are barren, and in the more northerly cores diatom content is commonly 2–4%, though preservation is poor. In units A, C and E diatoms are moderately preserved. The flora is of open-ocean type, dominated by Fragilariopsis kerguelensis, with Thalassiosira lentiginosa, Actinocyclus actinochilus, Eucampia antarctica and Thalassiothrix sp. Resting spores of Chaetoceros spp. are notably rare, although they are abundant in cores on the adjacent continental shelf (Pudsey et al., 1994). Hemidiscus karstenii occurs rarely in unit E (see below). In units B and D diatom content is low, and the assemblage is very poorly preserved with

Fig. 6. Major element geochemistry. (A) SiO2/Al2O3 for core 106. The silica/alumina ratio is highest in Units A and C, and shows a good correspondence to intervals of high diatom % (compare Fig. 5). (B) Cumulative plot of major oxides (excluding SiO2, P2O5 and loss-on-ignition) for core 106. Note the lack of variation downcore except for high values of CaO in part of Unit C, corresponding to high biogenic CaCO3 (Fig. 5). This is an excellent illustration of the uniformity of terrigenous sedimentation during glacial intervals. (C) Trends in selected major oxides from southwest to northeast along the margin, in units A and C. TiO2 is plotted on an expanded scale (right-hand axis). Al2O3 and K2O decrease northeastwards while TiO2 and CaO (in unit A) increase. CaO in unit C is at a maximum in core 111. Note that unit A is only represented by two samples per core. (D) Trends in selected major oxides from southwest to northeast along the margin, in units B and D. Al2O3 and K2O show a clear decrease northeastward, while TiO2 and CaO increase.

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Fig. 7. Relative abundance of the clay minerals smectite, chlorite and illite in cores 108 and 113.

some reworking (e.g. the Lower Pleistocene form A. ingens and the Miocene Denticulopsis sp.). The biogenic carbonate content of these cores is mostly very low (1% or less) except in units C and E. In unit C, CaCO3 can attain 20–30% in the distal cores and 30–50% in the proximal cores, while in unit E it reaches 4–14% of the sediment (Fig. 5). Much of the carbonate comprises fragments of the thickshelled planktonic foraminifer Neogloboquadrina pachyderma sinistral. Preservation is better, i.e. there are more complete tests, in those samples with higher CaCO3. A few calcareous benthic foraminifers occur, and a small fraction of the biogenic carbonate consists of calcareous nannofossils (mostly Gephyrocapsa spp.; Diane Winter, pers. comm.). In unit C in cores 106, 108 and 111 there is a distinct double peak in CaCO3, corresponding to a brown–grey–brown colour sequence, in the upper part of the unit; maximum concentrations of foraminifers occur above the highest diatom peaks. In all the proximal cores the

unit C foraminiferal interval is thin (3–11 cm) and with a much higher concentration of complete tests. Total organic carbon was measured in all core tops, downcore in four complete cores and in the carbonatebearing intervals of the others. From surface values of 0.2–0.3% there is a gradual downcore decrease to 0.15–0.2% in unit E. There are no marked trends in TOC, either geographical or corresponding to the lithological cyclicity. The only exception is core 101 which has somewhat higher (0.3–0.5%) TOC in units A and C. 3.5. Major element chemistry Chemical composition was measured on core 101 (Drift 1), core 111 (Drift 4) and core 106 (Drift 6); here we also draw comparisons with earlier data from three cores on Drift 7 (Pudsey and Camerlenghi, 1998). The average major oxide composition is about 60% SiO2, 16% Al2O3, 7% Fe2O3(T) and

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Fig. 8. Ice-rafted debris (IRD) in core 110. (A) Weight % sand and weight % coarser than 500 mm. Sand peaks occur in units C and E. The sand fraction 63–500 mm includes biogenic grains (mostly planktonic foraminifera). (B) Relative abundance of grain types in the fraction coarser than 500 mm. In samples with very little sand, the grains were classified individually. In the larger sand fraction (coarse sand sample weighing up to 350 mg), grain types were categorised as abundant, common, rare, or single grain present. These were then weighted numerically and summed to 100%. The bars for each sample are scaled to a total of 1 or 10, to give a clearer representation of the amount of coarse sand.

2–3% each MgO, Na2O and K2O. CaO forms 3–10% and TiO2, MnO, and P2O5 less than 1% each. Fig. 6 illustrates geochemical variation downcore in core 106, and from southwest to northeast along the continental margin. Except for samples with high SiO2 and CaO in Unit C corresponding to high amounts of diatoms and foraminifera, the composition is very uniform downcore. This is also the case for the other three cores measured, which are compared in Figs. 6C and D. Fig. 6D shows small but significant trends in Al2O3, CaO, K2O and TiO2 in the terrigenous units of the cores; trends are less clear in the more biogenic units A and C, particularly in CaO, reflecting

the presence of foraminifera (Fig. 6C). The other major oxides do not vary significantly along the margin. 3.6. Clay mineralogy Cores 108 and 113, which extend down to unit E, were sampled for X-ray diffraction analysis. The minerals present in the ⬍2 m fraction are quartz, feldspars, chlorite, illite and smectite, with a small quantity of amphibole throughout, and calcite in the foraminiferal units C and E. Kaolinite was not detected. The clay mineralogy varies downcore

326 C.J. Pudsey / Marine Geology 167 (2000) 313–338 Fig. 9. Biogenic barium and 230Thexcess. (A) Biogenic barium (open symbols), lithogenic barium (filled symbols), and 230Thexcess (crosses) for core 106. Peaks in Babio and 230Thexcess correlate well with units A, C and E defined on lithology including biogenic silica content. (B) Biogenic and lithogenic barium for core 111. (C) Biogenic and lithogenic barium for core 101.

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corresponding to the lithological cyclicity already described (Fig. 7). Smectite is relatively abundant in units A, C and E with more chlorite in units B and D. The proportion of illite is constant at about 40–45%. Geographic variation is also apparent, with more smectite in core 113. 3.7. Ice-rafted debris Fig. 8 illustrates the downcore occurrence and composition of IRD in core 110. A lower grain size limit of 500 mm was chosen for microscopic examination, though it is recognised that fine sand, silt and clay-sized material can also be iceberg-rafted to the sites. Coarse sand is common only in units C and E, except for one sample near the top of unit B. The following grain types were distinguished: quartz, feldspar, acid/intermediate plutonic rock fragments, basalt, pumice shards, other volcanics, sedimentary rocks, and metamorphics. The occurrence of petrographic types is very different in units C and E from the rest of the core. In units A, B and D the few grains present are mainly volcanics, whereas units C and E contain a wider range of grain types and a higher proportion of quartz, feldspar and acid plutonic fragments. Additional cores are being studied to determine whether this pattern occurs elsewhere along the continental margin (O’Cofaigh and Pudsey, 2000). 3.8. Age In common with most sediments from south of the Polar Front, these cores do not contain biogenic carbonate except in a few intervals, and it was not therefore possible to obtain an oxygen isotope curve from foraminifera for dating purposes. The proportion of TOC is low (0.3% or less) and radiocarbon ages may consequently be unreliable because of the likely presence of reworked, older carbon. However, it is possible to identify the Quaternary warm isotope stages 1, 5 and 7 using a combination of siliceous biostratigraphy and chemical and isotope stratigraphy. 3.8.1. Biostratigraphy All the samples which contain diatoms belong to the T. lentiginosa zone (0–0.6 ka; Gersonde and Burckle, 1990), based on the presence of T. lentiginosa and the absence of A. ingens and Coscinodiscus

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elliptipora. Unit E in cores 106, 108, 110 and 113 contains H. karstenii. Although not common, the H. karstenii frustules are well preserved and are not considered to be reworked. H. karstenii ranges from the base of isotope stage 11 to the top of stage 7 (Burckle and Burak, 1988). At ODP Site 1095 (Fig. 1), it was observed to be well preserved in stage 7, but rare and poorly preserved in stages 9 and 11. In the absence of hiatuses in the cores, unit E is considered to correspond to stage 7. Radiolarians although common in the sand fraction in only a few samples (mainly in unit C) support the Late Quaternary age; Antarctissa spp. are abundant, and Stylatractus universus is absent. 3.8.2. Barium stratigraphy In cores with a significant non-biogenic component, as here, the contribution of Ba from detrital aluminosilicate minerals must be subtracted from total Ba to calculate biogenic Ba. A correction factor may be applied by assuming that all the Al is of detrital origin, and that the Ba/Al ratio in aluminosilicates is constant. Then Balith ˆ Al × ratio, and Babio ˆ Batotal ⫺ Balith (Dymond et al., 1992). The average worldwide elemental abundances for the upper continental crust given by Taylor and McLennan (1985) correspond to a Ba/Al ratio of 0.0069. In these cores, however, the ratio of 0.0069 yields negative (to ⫺230 ppm) biogenic Ba in samples from units B and D, which cannot be correct. A lower Ba/Al ratio is required to make calculated biogenic Ba zero or positive. Nu¨rnberg et al. (1997) used a ratio of 0.0067, calculated from terrigenous cores in the southern Weddell Sea. We require a still lower value of 0.0057 in core 106, 0.0048 in core 111 and 0.004 in core 101. Calculated biogenic Ba in units A and C is 400– 1100 ppm in cores 106 and 111, (Fig. 9; comparable with Drift 7; Pudsey and Camerlenghi, 1998), and somewhat lower in core 101. An even lower Ba/Al ratio is possible, and would yield higher Babio throughout. The near-constant values of calculated Balith downcore (Fig. 9) suggest that all the Al is of detrital origin, with no effect attributable to biogenic scavenging of Al (c.f. Bonn et al., 1998). The low biogenic Ba values in core 101 are not yet understood. They are unlikely to result from Ba mobilisation in reducing conditions, since TOC values, although higher than

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Fig. 10. Sedimentation rates. Stage boundaries from Martinson et al. (1987), lithological unit boundaries from Fig. 3. One symbol is used per drift, with thick lines representing proximal sites and thin lines distal sites. Where no symbol is shown at the end of a line, the stage was not fully recovered in that core, i.e. the sedimentation rate shown is a minimum.

elsewhere in this area, are comparable with Scotia Sea cores where biogenic Ba shows clear peaks in interglacial sediments (Pudsey and Howe, 1998). 3.8.3. Thexcess dating 230 Thexcess data for core 106 are shown in Fig. 9. High values occur in lithological units A and C; unit C has a double peak from 5–6 m depth, corresponding to the double peaks in carbonate and silica (Fig. 5).

4. Discussion 4.1. Age model and sedimentation rates A preliminary age model is presented in Fig. 10, based on the lithological succession in all cores supplemented by new geochemical data on three cores and 230Th data for core 106. The model depends on comparison with deep-water cores from other parts of the Antarctic margin. The assumption is made that

the succession is continuous from the seabed downwards, i.e. that hiatuses are not present at the selected sites (in contrast to cores on the steep sides of drift 7; see cores 1, 5 and 9 in Pudsey and Camerlenghi 1998). The upper part of unit A (homogenous brown diatom-bearing mud) is interpreted to be Holocene in age, with high values of biogenic silica, Babio and 230 Thex. Using the recovered thicknesses, Holocene sedimentation rates range from 0.4 to 4.2 cm/ka (Fig. 10), but these are considered unreliable because of likely loss of core-top material. Unit B is considered to represent glacial stages 2–4, with the lower part of unit A perhaps being a deglaciation sequence (see below). Low values of biogenic carbonate, silica, Babio and 230Thex reflect annual sea ice cover much more extensive than in the present interglacial (Nu¨rnberg et al., 1997; Bonn et al., 1998). Stage 3 has not been unambiguously identified using palaeoproductivity indicators in Antarctic continental margin sediments (Frank et al., 1995; Bonn et al., 1998; Ceccaroni et al., 1998). Parts of unit B with

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slightly higher diatom content may represent stage 3 (3.9–4.5 m in core 103, 1.8–2.7 m in core 108, 1.1– 1.7 m in core 110, 2.4–3.4 m in core 113; Fig. 5). Average sedimentation rates for stages 2–4 are 4.6– 13.5 cm/ka, and are higher in proximal than distal cores (Fig. 10). Stage 5 was a time of high productivity south of the present-day Polar Front, with particularly high biogenic flux in stage 5e (Frank et al., 1995; Nu¨rnberg et al., 1997). Unit C represents stage 5, the highest Babio and diatom silica values corresponding to stage 5e. In core 106 the upper peak in silica, Babio and 230 Thex may represent stages 5a–c and the lower, larger peak stage 5e. The possibility that unit B represents only stage 2 and unit C is a condensed section comprising the whole of stages 3–5e cannot be entirely ruled out. This would, however, require an unrealistically large variation in terrigenous sedimentation rates between stage 2 and stages 3–5e. Additional independent dating techniques would be required to verify this possibility. Bulk sedimentation rates for stage 5 are 1.8–4.3 cm/ka (Fig. 10); if unit C corresponds to stage 5e only, these would increase by two or three times. Unit D is interpreted as stage 6; where recovered complete its sedimentation rate is 1.8–6.4 cm/ka. Unit E has high carbonate and silica values and contains H. karstenii, and is therefore considered to represent stage 7 (1.1–1.9 cm/ka; Fig. 10). Stage 8 may just occur at the base of cores 108, 110 and 113. Because this age model is preliminary, no attempt is here made to calculate fluxes, nor to quantify lateral sediment redistribution vs. palaeoproductivity changes (c.f. Frank et al., 1995). Overall Late Quaternary deposition was slower than the accumulation rates of the drifts throughout the Neogene (estimated as 14– 17 cm/ka using seismic profiles and knowing the age of the basement; Rebesco et al., 1997) but comparable with Pleistocene–Late Pliocene accumulation rates measured on ODP Leg 178 (5–10 cm/ka; Barker et al., 1999). 4.2. Sediment supply and transport The two principal sources of sediment to the drifts are biogenic production in the surface waters (controlled by temperature, nutrients and the seasonal sea ice edge) and terrigenous supply from the conti-

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nental margin (controlled by glacial processes and deep-sea currents). Both have varied greatly over late Quaternary glacial cycles. Warm, nutrient-poor surface water to the north of the Polar Front favours the production of biogenic carbonate, while cold, nutrient-rich water to the south favours biogenic silica. Preservation of carbonate and silica in deep-sea sediment is a function of bottom water corrosiveness. Although the oceanography of the study area is not well known, comparison may be made with the Scotia Sea and Weddell Sea, which are better documented (Fig. 1; Olbers et al., 1992). In the northern Scotia Sea at 3000–4000 m depth, where the deepest water mass is Circumpolar Deep Water (Locarnini and Whitworth, 1993), foraminifera and poorly-preserved nannofossils occur in Holocene sediments from the area of the Polar Front and for about 100 km to the south (Howe et al., 1997; Pudsey and Howe, 1998). Diatoms are abundant throughout the Scotia Sea in interglacial sediments, but decrease southwards. Weddell Sea Bottom Water is corrosive to silica, and severe dissolution of diatoms in surface sediments was observed by Zielinski and Gersonde (1997), even in areas of the Weddell Sea which have several months per year of open water. The Weddell Sea is thought to be the source for the bottom water flowing SW along the Antarctic Peninsula slope (Gordon, 1966; Nowlin and Zenk, 1988) so some silica dissolution might be expected in this vicinity. Gradients in biogenic production west of the Antarctic Peninsula are likely to be perpendicular to the sea ice edge and Polar Front, i.e. approximately north–south (Fig. 1; though note that the Polar Front swings to the south near drift 4). Gradients in preservation may be related to water depth and to the tongue of cold water flowing SW along the slope. For terrigenous sediment, the relative importance of downslope and alongslope transport is of interest. It is clear from the pattern of progradation of the continental margin and the existence of long-lived northwest-flowing turbidity current channels (Tomlinson et al., 1992; Larter et al., 1997) that downslope processes have played a major role in shaping the continental shelf, slope and rise. On the upper continental rise, though, there is potential for alongslope redistribution of fine sediment (supplied by ice rafting, in turbidity currents and in meltwater plumes)

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by southwest-flowing currents (Rebesco et al., 1997; Pudsey and Camerlenghi, 1998). Differences in sediment composition between drifts, related to differences in onshore geology, may reveal the relative contribution of local and distant sources. Survey lines across the drifts (southwest–northeast) do not show significant thickness changes in the upper sediment layers, except on the very edges of the drifts within a few km of the channels. In contrast with the offshore thinning (southeast–northwest; Fig. 2), this immediately suggests that downslope transport was by far the more important process (Pudsey et al., 2000). 4.3. Glacial–interglacial depositional environments This study includes cores which recovered three glacial cycles, and it is apparent that the Late Quaternary cyclicity is by no means regular. Each warm stage and each cold stage is slightly different, and the record of the most recent deglaciation differs from the earlier two. We discuss the Quaternary record from the top downwards; palaeoenvironmental reconstructions are shown in Fig. 11. 4.3.1. Interglacials: units A, C and E (Fig. 11A) In the present interglacial, ice is grounded at the coast. There are small areas of floating ice shelf around Alexander Island, but the rest of the continental shelf area has only seasonal sea ice cover. Terrigenous sediment is released from ice at the grounding line and is largely trapped on the shelf. Icebergs are common in the area, but few carry any debris (Anderson et al., 1980). The Polar Front is 150–300 km north of the drifts, and the average annual extent of open water ranges from about 3 months in the southwest to 7 months in the northeast. A gentle southwest-flowing slope current can maintain fine silt and clay in suspension, thus redistributing sediment.

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The environment of deposition is a low-energy one dominated by hemipelagic settling; the sediments are hemipelagites or muddy contourites. Slow accumulation and an abundant food supply favour reworking of the sediment by bioturbation. Biogenic silica and TOC do not show any systematic northwards or offshore trend, though there is a southwestward increase in biogenic barium. At present no calcareous foraminifera or nannofossils are preserved in surface sediments, though they do accumulate in deep sediment traps (Harland and Pudsey, 1999). In the Scotia Sea, the southern limit of core-top carbonate preservation is about 100 km south of the Polar Front. The occurrence of planktonic foraminifera and calcareous nannofossils in part of unit C suggests that the Polar Front lay nearer to the Antarctic Peninsula margin at that time. During peak productivity represented by highest diatom abundance, high organic matter flux to the sediments resulted in dissolution of biogenic carbonate. Warmer water would have facilitated iceberg melting and hence the release of IRD. Intermittent southward transport of warmer-water plankton may have been facilitated by shedding of warm-core rings from the Polar Front (Gouretski and Danilov, 1994). As at present, slow sedimentation and an abundant food supply favoured bioturbation of the sediment. Diatom content in unit C decreases southwards, perhaps resulting from partial sea-ice cover over the sites of cores 103 and 107 and Drift 7. Within the foraminiferal interval there was a period of faster current flow which winnowed only the shallow sites, leaving a thin layer of carbonate-rich sand (Fig. 5). This could have been an intensification of the slope current or a southward excursion of the path of the Antarctic Circumpolar Current (Pudsey et al, 2000). The winnowing episode affected all the shallow sites, regardless of distance from the present-day ACC axis, which makes enhanced Weddell Sea

Fig. 11. Suggested palaeoenvironments of interglacial and glacial deposition. Downslope flow shown as open arrows, alongslope flow as solid arrows. Drift outlines and core sites from Fig. 1. (A) Interglacial; inset profiles show the differences between Holocene (Pudsey and Camerlenghi, 1998, their Fig. 13) and earlier warm stages. Alongslope transport by contour currents and ice rafting was significant. (B) Glacial; the Polar Front and maximum extent of sea ice lay north of the area, and the minimum extent of sea ice was much farther north than today. Ice streams occupied the troughs on the shelf but the upper slope acted as a line source. The relative abundance of major elements (e.g. Al, Ca, K and Ti) and of clay minerals (illite, chlorite, and smectite) varied along the margin, attesting to the importance of local sources and downslope transport. Inset profiles show the differences between stages 2–4 and 6, with more turbidites in stage 6 (particularly in the north) and more iceberg rafting.

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outflow more likely; this would also be consistent with the interglacial intensification of the Weddell Gyre inferred from Weddell Sea cores by Pudsey (1992) and Grobe and Mackensen (1992). Detailed grain-size analysis of unit C in all the cores is in progress, to reconstruct the palaeoflow history of this area. Geochemical trends in unit C are poorly defined (Fig. 6C). The clay fraction contains abundant smectite in both cores analysed (Fig. 7). Smectite is abundant in the Cenozoic volcanics of the South Shetland Islands, Anvers and Brabant islands, where its formation is thought to have been contemporaneous with eruption (John Smellie, pers. comm.). It is unlikely to be a product of present-day chemical weathering in the Antarctic, but may also be related to recent volcanism, e.g. at Deception Island (Petschick et al., 1996). Evidence for the Stage 7 palaeoenvironment comes from only four distal cores which recovered unit E, but it appears to have been comparable to Stage 5, and the same reconstruction applies (Fig. 11A). Carbonate contents are lower than in unit C, while the amount of IRD is similar. No evidence is seen for current winnowing. 4.3.2. Glacials: units B and D (Fig. 11B) Depositional models based on seismic stratigraphic analysis show the ice sheet grounded to the shelf edge at glacial maximum (Larter and Barker, 1991). Sediment transported within the basal ice or in the deforming till layer (e.g. Hart, 1998) was released during melting at the shelf break and upper slope, leading to oversteepening, frequent slope failures (Vanneste and Larter, 1995) and the initiation of small turbidites. Note that, although ice flow is believed to have been concentrated in ice streams located in the shelf troughs (Pudsey et al., 1994; Rebesco et al., 1998), there are no canyons on the slope corresponding to the ends of troughs. Furthermore, the heads of the channels at the base of the slope are located in between, as well as directly in front of, troughs. The slope appears to have behaved as a line source. Unit B is dominated by very fine-grained mud (commonly ⬎ 70% clay, finer than 4 mm, on Drift 7; Pudsey and Camerlenghi, 1998). Fine sediment could have been supplied by ice rafting, turbidity currents, meltwater plumes or bottom-current erosion farther

upstream, i.e. to the northeast. The last is considered insignificant, on account of: (i) the slow steady flow measured at the South Shetland Trench (Nowlin and Zenk, 1988); and (ii) the likelihood of even slower flow during glacials as the Weddell Gyre slowed down (Pudsey, 1992). Supply to Drift 7 by turbidity currents and meltwater plumes was discussed by Pudsey and Camerlenghi (1998), who noted the potential for high but dilute turbidity-current suspension clouds arising from a hydraulic jump at the base of the very steep continental slope. These authors also suggested that melting at the ice front in relatively warm water at the shelf break at glacial maximum (CPDW; Jacobs, 1989) could generate sedimentladen meltwater plumes. By analogy with presentday Antarctic fjords and the Svalbard coast, plumes could spread laterally at intermediate depths or on the surface (Pfirman and Solheim, 1989; Domack and Ishman, 1993). Eventually their entrained sediment settled out over the slope and upper rise. Unit B thins offshore but does not thin laterally away from the channels. This implies that a continental-margin line source (the ice front) was more significant than channelised sources (turbidity currents). The rarity of graded silt laminae except in the three northern drifts suggests a lack of turbidity currents thick enough to extend over the drifts from the floor of the adjacent channels (a height of 400–600 m at the proximal sites and 200–400 m distally). Compositional variation along the margin is quite marked in unit B. Major element composition shows clear trends (Fig. 7) and the proportion of smectite is higher northwards. Only a few lithological features are continuous between two or three adjacent drifts, such as the occurrence of “reddish” bands on Drift 6, similar to those described from Drift 7 by Pudsey and Camerlenghi (1998). The low biogenic content suggests that the annual extent of sea ice was considerably greater than today. Local polynyas along the ice front (Zwally et al., 1985) may have allowed a little diatom productivity. Unit D is comparable in thickness to unit B in the three distal cores where it was recovered complete; the acoustic evidence suggests it too thickens inshore. It is coarser than unit B, containing graded silt to fine sand layers up to 60 mm thick interpreted as turbidites. Laminae of dispersed sand grains and granules are thought to result from ice rafting rather than

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current winnowing, on account of their very poor sorting (Pudsey and Camerlenghi, 1998). Terrigenous supply by turbidity currents and ice rafting was probably more important than by meltwater plumes in unit D (Fig. 11B). The lower diatom content suggests the seasonal ice cover was greater, if anything, than unit B. The increase in IRD might reflect inhibition by sea ice of iceberg drift away from the area. Alternatively there may have been more icebergs, or they were carrying more debris, perhaps as a consequence of greater snow accumulation and throughput at that time. The same trends in terrigenous composition are observed as in unit B. Magnetic susceptibility is high and variable, responding to the silt layers. 4.3.3. Deglaciations The stage 8/7 transition (base of unit E) and the stage 6/5 transition (base of unit C) are both quite sharp in this suite of cores. Upward increases in biogenic silica content, sand content and biogenic barium occurs within a few cm and corresponds to a rapid transition from laminated to burrowed sediment. This is interpreted to reflect a rapid deglaciation of the shelf and decrease in annual sea ice cover. The smooth, landward-sloping continental shelf is close to the equilibrium grounding profile of a low-profile ice sheet, and there are few shoal areas to act as pinning points during ice retreat. As sea level rose during climate warming, the whole of this section of the Antarctic ice sheet may have lifted off the continental shelf at once to become a floating ice shelf. Sub-ice transport of debris to the shelf edge thus ceased abruptly, and during the interglacials sediment could only reach the drifts by iceberg rafting and by minor mass wasting on the slope (Fig. 11). The stage 2/1 transition (base of unit A) looks rather different. The lithological transition from unit B is gradual, over as much as 1.0 m of variegated, laminated mud showing minor bioturbation. The burrows in this transition are much larger than the tiny burrows which occur rarely within unit B. Although diatom % is high only in the top 7–27 cm of red–brown structureless mud, high values of Ba, Si/Al and smectite continue down into the variegated interval. Supply of fine terrigenous sediment to the drifts seems to have continued at the same time as the ocean was warming and annual sea ice extent

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decreasing, allowing biogenic productivity. There is evidence from the continental shelf off Anvers Island that an ice sheet retreated from the middle and outer shelf rapidly and that open marine conditions were established at about 11 kyr BP (Pudsey et al., 1994). The reduction in annual sea ice cover offshore may have occurred some time prior to this. 4.4. Chemistry, mineralogy and provenance The dominance of physical over chemical weathering in the Antarctic is exemplified by the majorelement composition of this suite of cores. On chemical variation diagrams (e.g. K2O/SiO2, MgO/ SiO2, Rb/SiO2, K2O/Na2O; Moyes et al., 1994; Leat et al., 1995) they plot within the fields of Antarctic Peninsula volcanic and plutonic rocks, although CaO tends to be higher, and K2O lower, in Peninsula igneous rocks of similar (andesitic) SiO2 content. Published studies of igneous geochemistry have emphasised compositional variation across the arc rather than along it (e.g. Saunders et al., 1982; Weaver et al., 1982; Moyes et al., 1994). Nevertheless, comparison of volcanics and plutonics shows some southwest–northeast trends in major and trace elements (Alexander Island—Peninsula at 54⬚S— South Shetland Islands; Weaver et al., 1982; Smellie et al., 1984; McCarron and Smellie, 1998; Philip Leat, unpubl. data). For equivalent SiO2 content, TiO2 increases northeastwards while K2O and Ba decrease, in agreement with chemical trends observed in the cores (Fig. 7). However, Fe2O3, MgO and MnO2 all increase northeastwards in the igneous source rock, the opposite trend from the cores. The influence of the sedimentary and metasedimentary source terrain is hard to evaluate, as there are no chemical data on the Le May Group sediments on Alexander Island. Unpublished analyses of the Trinity Peninsula Group show it to be generally silica-rich (65– ⬎ 80% SiO2; John Smellie, pers. comm.) though the most mafic compositions are very similar to the cores. The low Ba:Al ratios (0.004–0.0057; Fig. 9) are inconsistent with an “average continent” source for the sediments (upper crust Ba/Al is 0.0069; Taylor and McLennan, 1985). A contribution from the lower crust would reduce the ratio towards the bulk continental crust value of 0.0030 (Taylor and McLennan, 1985). This supports the conclusion of Leat et al.

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(1995) that the lower Cretaceous plutons were derived from partial melting of the lower crust. Also, Hamer and Moyes (1982) suggested that the Antarctic Peninsula Volcanic Group in the northernmost Peninsula was derived from melting at depths ⬎25 km; PC101, the northernmost core, has the lowest Ba/Al. Further trace element and isotopic studies are required to identify specific source areas. Of interest would be any offshore chemical evidence for the segmentation of the Peninsula arc suggested by Hawkes (1982) and Thomson et al. (1983). These results demonstrate the utility of geochemical data in tracing sediment source areas. In some circumstances, such as where chemical weathering is insignificant, chemistry may be more diagnostic than clay mineralogy. The clay mineralogy data from two cores show a strong northeastward enrichment in smectite, which is interpreted to reflect volcanic influence from the South Shetland Islands, Anvers and Brabant islands. No kaolinite was detected, and this is evidence against a contribution of fine sediment from the Weddell Sea (Petschick et al., 1996). 4.5. Comparison with the Weddell Sea margins The continental rise west of the Antarctic Peninsula, recording fluctuations of the small ice sheet on the peninsula, presents some interesting contrasts with the southeastern Weddell Sea which borders the large East Antarctic Ice Sheet (Grobe and Mackensen, 1992) and also with the northwest Weddell Sea near the South Orkney Islands (Frank et al., 1995). In the southeastern Weddell Sea the exposed continental shelf is narrow (30–80 km), with ice shelves extending up to 100 km inland. The continental ice drainage basin extends some 400 km farther inland with ice up to 2 km thick (Drewry, 1983). The South Orkney Islands have a broad shelf (150 km on the south side) but constitute a very small glacial source. The cyclonic Weddell Gyre (Fig. 1) affects the whole margin; bottom current speeds average 10–15 cm/s over the upper slope, but less than 5 cm/s over the lower slope and rise (Fahrbach et al., 1994). A comprehensive account of late Quaternary sedimentation on the Weddell Sea slope terrace at about 5–15⬚W was published by Grobe and Mackensen, 1992; Fig. 1, inset). Cores from 2000–2800 m depth contain sufficient biogenic carbonate to measure d 18O

and derive an age model. A comparable cyclicity in biogenic silica was observed, with peaks in interglacial stages 1, 5e and 7. There are two major contrasts with the Antarctic Peninsula margin. Firstly, in common with the area near the South Orkney Islands, sedimentation rates are higher in interglacial than glacial stages. Secondly, biogenic CaCO3 (as N. pachyderma) occurs throughout most of the succession in the SE Weddell Sea, being absent only in peak interglacials. It commonly forms 10–12% of the sediment in “moderate interglacials” (i.e. the upper parts of stages 5 and 7) and 1–2% in glacials. Other differences from the Antarctic Peninsula margin include the occurrence of more smectite in glacials than interglacials, stronger variations in the detrital components in glacials, and the occurrence of bioturbation of variable intensity throughout glacials (Grobe and Mackensen, 1992). In the depositional model of Grobe and Mackensen (1992), the dominant process during deglaciation and peak interglacials was considered to be ice-rafting. The relatively large ice drainage basins could supply large volumes of unsorted sediment directly to the slope terrace. Enhanced bottom currents winnowed most of the clay from peak interglacial sediments. In the northwestern Weddell Sea, Frank et al. (1995) attributed the high interglacial accumulation rates to sediment focusing as well as to high biogenic flux; about 50% of the interglacial flux of radionuclides was thought to be supplied by boundary scavenging. Grobe and Mackensen (1992) suggested a combination of continued ice rafting with weak bottom currents to explain the fine sediments deposited during “moderate interglacial” to glacial transitions; at glacial maxima, sediment was transported downslope in submarine canyons, bypassing the slope terrace en route to the deep sea. Frank et al. (1995) invoked redeposition of older margin sediments during glacial maxima, to explain very low radionuclide fluxes in the northwestern Weddell Sea. Higher glacial carbonate content in the southeastern Weddell Sea reflects the generally deeper CCD in this area (about 4000 m at the present day; Mackensen et al., 1990). Circumpolar Deep Water flowing south into the eastern end of the Weddell Gyre is considerably less corrosive to carbonate than the Weddell Sea Deep Water flowing north at the western end of the gyre, which is a mixture of CPDW with well-oxygenated

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Table 2 The main differences between glacial–interglacial deposition west of the Antarctic Peninsula and in the southeastern Weddell Sea Antarctic Peninsula, west side (this paper)

SE Weddell Sea (Grobe and Mackensen, 1992)

Poor carbonate preservation; outflow of Weddell Gyre, “young” bottom water generally corrosive to carbonate

Moderate carbonate preservation; inflow of Weddell Gyre, “old” deep water allows preservation of carbonate except during peak productivity Sedimentation rates higher during interglacials; narrow shelf, sediment delivered directly to upper slope even during interglacials Depocentres on slope and abyssal plain; slope terrace at middepth; Wegener Canyon was a conduit to the abyssal plain when the ice sheet was grounded to the shelf edge IRD present throughout glacials; bergs derived from large drainage basin with thick ice, also less dilution with fines

Sedimentation rates higher in glacials, especially stages 2–4; broad inward-sloping shelf acts as a sediment trap during interglacials Depocentres on upper continental rise; steep, smooth slope, so the ice margin was a line source to the continental rise when the ice sheet was grounded to the shelf edge IRD rare in glacials, as isolated laminae; bergs derived from small drainage basins with thin ice, and IRD diluted by copious fine material Detrital component very uniform in glacials; reflects downslope transport from limited area of the Antarctic Peninsula Smectite in interglacials; attributed to alongslope transport from known sources to NE

shelf water (Fahrbach et al., 1994). Hence carbonate preservation is poor in sediments beneath the Weddell Sea outflow, i.e. near the South Orkney Islands and westwards along the Antarctic Peninsula continental rise (Fig. 1). Note that while average annual sea ice extent (controlled by local air and sea temperatures, which can change significantly over years to decades) can vary rapidly in both the Antarctic Peninsula and the southeastern Weddell Sea, the response time of the East Antarctic Ice Sheet to temperature change is very much longer than that of the northern Peninsula ice sheet and glaciers. Grobe and Mackensen (1992) suggested a lag of 10–15 kyr between glacial terminations (delineated by d 18O in foraminifera and high abundance of radiolarians) and the first occurrence of IRD in East Antarctic slope sediments. On the Antarctic Peninsula sediment drifts, the onset of biogenic sedimentation and of ice rafting were synchronous at each glacial termination, within the resolution of our time scale. The differences between the Antarctic Peninsula and the southeastern Weddell Sea, summarised in Table 2, can be accounted for by the differences in glacial and oceanographic setting and in continental margin physiography between the two areas. The northwestern Weddell Sea (Frank et al., 1995) is intermediate in character, but more similar to the Antarctic Peninsula.

Detrital component variable in glacials; may reflect shelf processes or alongslope transport Smectite in glacials; attributed to long-distance current transport (“marine sediment noise”)

5. Conclusions Cores were recovered from six sediment drifts spanning some 650 km of the Antarctic Peninsula margin and 4⬚ of latitude. Biogenic barium and 230 Thex concentrations, combined with lithological correlation, allowed identification of isotope stages 1 (Holocene) and 5 in all cores. Stage 7 is indicated by the diatom H. karstenii near the base of four cores. Diatoms comprise an open-ocean assemblage. Foraminiferal intervals occur in interglacial stages 5 and 7 but not in the Holocene. Processes of terrigenous sediment supply during glacial stages differed: hemipelagic sediments derived from meltwater plumes were more important in stage 6, turbidites and IRD in stages 2–4. Glacial Termination I appear to have been more protracted than Terminations II or III. The terrigenous component shows compositional changes along strike, notably in glacials. Magnetic susceptibility varies along and between drifts. The major oxides Al2O3 and K2O are higher in the southwest, and CaO and TiO2 higher in the northeast. Low lithogenic Ba/Al ratios are consistent with some of the source rocks having a lower-crustal origin. There is more smectite among the clay minerals in the northeast. These trends reflect glacial source variations along the margin. Interglacial sediments show less clear trends, and their IRD was derived from a

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wider area. Downslope processes were dominant in glacials, but alongslope processes may have attained equal importance in interglacials. The area contrasts with the East Antarctic continental slope in the southeastern Weddell Sea, which is dominated by ice-rafting, where interglacial sedimentation rates were much higher than glacial, and where the CCD is deeper. Studies of additional Antarctic margins are required to relate the variety of glaciological regimes on the continent to the high-resolution records on the slope and rise. Acknowledgements I thank my fellow scientists and the ship’s staff on cruise JR19 for their enthusiastic support during core collection. Francis Brearley and Simone Medonos assisted with core splitting and measurement, and sample preparation. Peter Floyd, Dave Emley and Margaret Aikin at Geoscience Analytical Services provided excellent service as always. Christopher Strobl and Bjorn Schroeder made the uranium and thorium measurements; Augusto Mangini is thanked for the provision of lab facilities. John Smellie and Philip Leat advised on geochemistry, and Diane Winter identified the calcareous nannofossils. The manuscript was improved after comments from Peter Barker, Martin Frank and one anonymous referee. References Anderson, J.B., Domack, E.W., Kurtz, D.D., 1980. Observations of sediment-laden icebergs in Antarctic waters: implications to glacial erosion and transport. J. Glaciol. 25, 387–396. Barker, P.F., 1982. The Cenozoic subduction history of the Pacific margin of the Antarctic Peninsula: ridge crest-trench interactions. J. Geol. Soc. London 139, 787–801. Barker, P.F., Camerlenghi, A., et al., 1999. Initial results of the Ocean Drilling Program, Leg 178. College Station, TX (Ocean Drilling Program), 60pp ⫹ CD. Biscaye, P.E., 1965. Mineralogy and sedimentation of recent deepsea clay in the Atlantic Ocean and adjacent seas and oceans. Geol. Soc. Am. Bull. 76, 803–832. Bonn, W.J., Gingele, F.X., Grobe, H., Mackensen, A., Fu¨tterer, D.K., 1998. Palaeoproductivity at the Antarctic continental margin: opal and barium records for the last 400 ka. Palaeogeogr. Palaeoclimatol. Palaeoecol. 139, 195–211. Burckle, L.H., Burak, R.W., 1988. Fluctuations in Late Quaternarydiatom abundances: stratigraphic and paleoclimatic impli-

cations from subantarctic deep sea cores. Palaeogeogr. Palaeoclimatol. Palaeoecol. 67, 147–156. Camerlenghi, A., Crise, A., Pudsey, C.J., Accerboni, E., Laterza, R., Rebesco, M., 1997. Ten-month observation of the bottom current regime across a sediment drift of the Pacific margin of the Antarctic Peninsula. Antarct. Sci. 9, 426–433. Ceccaroni, L., Frank, M., Frignani, M., Langone, L., Ravaioli, M., Mangini, A., 1998. Late Quaternary fluctuations of biogenic component fluxes on the continental slope of the Ross Sea, Antarctica. J. Mar. Sys. 17, 515–525. Cunningham, A.P., Vanneste, L.E., ANTOSTRAT Antarctic Peninsula Regional Working Group, 1995. The ANTOSTRAT Antarctic Peninsula Regional Working Group digital navigation compilation. Am. Geophys. Union, Antarct. Res. Series 68, A297–A301. Doake, C.S.M., Vaughan, D.G., 1991. Rapid disintegration of the Wordie Ice Shelf in response to atmospheric warming. Nature 350, 328–330. Domack, E.W., Ishman, S., 1993. Oceanographic and physiographic controls on modern sedimentation within Antarctic fjords. Bull. Geol. Soc. Am. 105, 1175–1189. Drewry, D.J., 1983. Antarctica: Glaciological and Geophysical folio. Scott Polar Research Institute, Cambridge, ivpp ⫹ 9 maps. Dymond, J., Suess, E., Lyle, M., 1992. Barium in deep-sea sediment: a geochemical proxy for paleoproductivity. Paleoceanography 7, 163–181. Fahrbach, E., Rohardt, G., Schroder, M., Strass, V., 1994. Transport and structure of the Weddell Gyre. Ann. Geophysicae 12, 840– 855. Fleming, E.A., Thomson, J.W., (compilers) 1979. Northern Graham Land and South Shetland Islands. Geological map, 1: 500000, BAS 500G series Sheet 2, British Antarctic Survey, Cambridge. Frank, M., Eckhardt, J-D., Eisenhauer, A., Kubik, P.W., DittrichHannen, B., Mangini, A., 1994. Beryllium 10, thorium 230 and protactinium 231 in Galapagos microplate sediments: implications for hydrothermal activity and paleoproductivity changes during the last 100,000 years. Paleoceanography 9, 559–578. Frank, M., Eisenhauer, A., Bonn, W.J., Walter, P., Grobe, H., Kubik, P.W., Dittrich-Hannen, B., Mangini, A., 1995. Sediment redistribution versus palaeoproductivity change: Weddell Sea margin sediment stratigraphy and biogenic particle flux of the last 250,000 years deduced from 230Thex, 10Be and biogenic barium profiles. Earth Planet. Sci. Lett. 136, 559–573. Gersonde, R., Burckle, L.H., 1990. Neogene diatom biostratigraphy of ODP Leg 113, Weddell Sea. In: Barker, P.F, Kennett, J.P. (Eds.), Proceedings of the Ocean Drilling Program, Scientific Results: 113. College Station, TX (Ocean Drilling Program), pp. 761–789. Gloersen, P., Campbell, W.J., Cavalieri, D.J., Comiso, J.C., Parkinson, C.L., Zwally, H.J., 1992. Arctic and Antarctic sea ice, 1978–1987: Satellite passive microwave observations and analysis. NASA SP-511, 290pp. Gordon, A.L., 1966. Potential temperature, oxygen and circulation of bottom water in the Southern Ocean. Deep-Sea Res. 13, 1125–1138. Gouretski, V.V., Danilov, A.I., 1994. Characteristics of warm rings

C.J. Pudsey / Marine Geology 167 (2000) 313–338 in the African sector of the Antarctic Circumpolar Current. Deep-Sea Res. 41, 1131–1157. Grobe, H., Mackensen, A., 1992. Late Quaternary climatic cycles as recorded in sediments from the Antarctic continental margin. Am. Geophys. Union, Antarct. Res. Series 56, 349–376. Hamer, R.D., Moyes, A.B., 1982. Composition and origin of garnet from the Antarctic Peninsula Volcanic Group of Trinity Peninsula. J. Geol. Soc. Lond. 139, 713–720. Harland, R., Pudsey, C.J., 1999. Dinoflagellate cysts from sediment traps deployed in the Bellingshausen, Weddell and Scotia seas, Antarctica. Mar. Micropaleontol. 37, 77–99. Hart, J.K., 1998. The deforming bed/debris-rich basal ice continuum and its implications for the formation of glacial landforms (flutes) and sediments (melt-out till). Quat. Sci. Rev. 17, 737–754. Hawkes, D.D., 1982. Nature and distribution of metalliferous mineralization in the northern Antarctic Peninsula. J. Geol. Soc. Lond. 139, 803–809. Howe, J.A., Pudsey, C.J., Cunningham, A.P., 1997. Pliocene– Holocene contourite deposition under the Antarctic Circumpolar Current, western Falkland Trough, South Atlantic Ocean. Mar. Geol. 138, 27–50. Jacobs, S.S., 1989. Marine controls on modern sedimentation on the Antarctic continental shelf. Mar. Geol. 85, 121–153. Kellogg, D.E., Kellogg, T.B., 1987. Recent glacial history and rapid ice stream retreat in the Amundsen Sea. J. Geophys. Res. 92, 8859–8864. Keys, J.R., 1990. Ice. In: Glasby, G.P. (Ed.), Antarctic Sector of the Pacific, 51. Elsevier Oceanography Series, Amsterdam, pp. 95– 123. Larter, R.D., Barker, P.F., 1991. Neogene interaction of tectonic and glacial processes at the Pacific margin of the Antarctic Peninsula. In: MacDonald, D.I.M. (Ed.) Sedimentation, Tectonics and Eustasy. International Association of Sedimentologists Special Publication 12, 165–186. Larter, R.D., Rebesco, M., Vanneste, L.E., Gamboa, L.A.P., Barker, P.F., 1997. Cenozoic tectonic, sedimentary and glacial history of the continental shelf west of Graham Land, Antarctic Peninsula. Am. Geophys. Union, Antarct. Res. Series 71, 1–27. Leat, P.T., Scarrow, J.H., Millar, I.L., 1995. On the Antarctic Peninsula batholith. Geol. Magazine 132, 399–412. Locarnini, R.A., Whitworth III, T., Nowlin Jr., W.D., 1993. The importance of the Scotia Sea on the outflow of Weddell Sea Deep Water. J. Mar. Res. 51, 135–153. Mackensen, A., Grobe, H., Kuhn, G., Fu¨tterer, D.K., 1990. Benthic foraminiferal assemblages from the eastern Weddell Sea between 68⬚ and 73⬚S: distribution, ecology and fossilization potential. Mar. Micropalaeontol. 16, 241–283. Martinson, D.G., Pisias, N.G., Hays, J.D., Imbrie, J., Moore Jr., T.C., Shackleton, N.J., 1987. Age dating and the orbital theory of the ice ages: development of a high-resolution 0 to 300,000year chronostratigraphy. Quat. Res. 27, 1–29. McCarron, J.J., Smellie, J.L., 1998. Tectonic implications of forearc magmatism and generation of high-magnesian andesites: Alexander Island, Antarctica. J. Geol. Soc. Lond. 155, 269–280. McGinnis, J.P., Hayes, D.E., 1995. The roles of down-slope and

337

along-slope depositional processes: southern Antarctic Peninsula margin. Am. Geophys. Union, Antarct. Res. Series 68, 141–156. McGinnis, J.P., Hayes, D.E., Driscoll, N.W., 1997. Sedimentary processes across the continental rise of the southern Antarctic Peninsula. Mar. Geol. 141, 91–109. Milne, A.J., Millar, I.L., 1989. The significance of mid-Palaeozoic basement in Graham Land, Antarctic Peninsula. J. Geol. Soc. London 146, 207–210. Moreton, S.G., 1999. Quaternary tephrochronology of the Scotia Sea and Bellingshausen Sea, Antarctica. Unpublished PhD thesis, Cheltenham and Gloucester College, 163pp. Morley, J.J., Hays, J.D., 1979. Cycladophora davisiana: a stratigraphic tool for Pleistocene North Atlantic and interhemispheric correlation. Earth Planet. Sci. Lett. 44, 383–389. Moyes, A.B., Willan, C.F.H., Thomson, J.W. and others, 1994. Geological map of Adelaide Island to Foyn Coast. BAS GEOMAP Series, Sheet 3, 1:250000, with supplementary text, British Antarctic Survey, Cambridge, 60pp. Nowlin, W.D., Zenk, W., 1988. Westward bottom currents along the margin of the South Shetland Island Arc. Deep-Sea Res. 35, 269–301. Nu¨rnberg, C.C., Bohrmann, G., Schluter, M., Frank, M., 1997. Barium accumulation in the Atlantic sector of the Southern Ocean: results from 190,000-year records. Paleoceanography 12, 594–603. O’Cofaigh, C., Pudsey, C.J., 2000. Ice-rafted debris on the Antarctic Peninsula continental rise and in the Scotia Sea over the last 250 ka (in preparation). Olbers, D., Gouretski, V., Seiss, G., Schroter, J., 1992. Hydrographic atlas of the Southern Ocean. Alfred Wegener Institute, Bremerhaven, xviipp ⫹ 82 plates. Orsi, A.H., Whitworth III, T., Nowlin Jr., W.D., 1995. On the meridional extent and fronts of the Antarctic Circumpolar Current. Deep-Sea Res. 42, 641–673. Pankhurst, R.J., 1982. Rb–Sr geochronology of Graham Land, Antarctica. J. Geol. Soc. Lond. 139, 201–211. Petschick, R., Kuhn, G., Gingele, F., 1996. Clay mineral distribution in surface sediments of the South Atlantic: sources, transport, and relation to oceanography. Mar. Geol. 130, 203–229. Pfirman, S.L., Solheim, A., 1989. Subglacial meltwater discharge in the open-marine tidewater glacier environment: observations from Nordaustlandet, Svalbard Archipelago. Mar. Geol. 86, 265–281. Pudsey, C.J., 1992. Late Quaternary changes in Antarctic Bottom Water velocity inferred from sediment grain size in the northern Weddell Sea. Mar. Geol. 107, 9–33. Pudsey, C.J., 1993. Calibration of a point-counting technique for estimation of biogenic silica in marine sediments. J. Sediment. Petrol. 63, 760–762. Pudsey, C.J., Camerlenghi, A., 1998. Glacial-interglacial deposition on a sediment drift on the Pacific margin of the Antarctic Peninsula. Antarct. Sci. 10, 286–308. Pudsey, C.J., Barker, P.F., Larter, R.D., 1994. Ice sheet retreat from the Antarctic Peninsula shelf. Continental Shelf Res. 14, 1647– 1675. Pudsey, C.J., Howe, J.A., 1998. Quaternary history of the Antarctic

338

C.J. Pudsey / Marine Geology 167 (2000) 313–338

Circumpolar Current: evidence from the Scotia Sea. Mar. Geol. 148, 83–112. Pudsey, C.J., Howe, J.A., Morris, P., Gunn, D.E., 2000. Processes on sediment drifts from 3.5 kHz data, Antarctic Peninsula continental rise. N. Z. J. Geol. Geophys., submitted for publication. Rebesco, M., Larter, R.D., Camerlenghi, A., Barker, P.F., 1996. Giant sediment drifts on the continental rise west of the Antarctic Peninsula. Geo-mar. Lett. 16, 65–75. Rebesco, M., Larter, R.D., Barker, P.F., Camerlenghi, A., Vanneste, L.E., 1997. The history of sedimentation on the continental rise west of the Antarctic Peninsula. Am. Geophys. Union, Antarct. Res. Series 71, 29–49. Rebesco, M., Camerlenghi, A., Zanolla, C., 1998. Bathymetry and morphogenesis of the continental margin west of the Antarctic Peninsula. Terra Antartica 5, 715–725. Saunders, A.D., Weaver, S.D., Tarney, J., 1982. The pattern of Antarctic Peninsula plutonism. In: Craddock, C. (Ed.). Antarctic Geoscience, University of Wisconsin Press, Madison, pp. 305– 314. Shimmield, G.B., Derrick, S., Mackensen, A., Grobe, H., Pudsey, C.J., 1994. The history of biogenic silica, organic carbon and barium accumulation in the Weddell Sea and Antarctic Ocean over the last 150,000 years. In: R. Zahn et al. (Eds.), Carbon cycling in the Glacial Ocean: constraints on the ocean’s role in global change. Proc. NATO Advanced Research Workshop, I 17, pp. 555–574. Sievers, H.A., Nowlin, W.D., 1984. The stratification and water masses in Drake Passage. J. Geophys. Res. 83, 3010–3022. Smellie, J.L., 1999. Lithostratigraphy of Miocene-Recent, alkaline volcanic fields in the Antarctic Peninsula and eastern Ellsworth Land. Antarct. Sci. 11, 362–378. Smellie, J.L., Pankhurst, R.J., Thomson, M.R.A., Davies, R.E.S., 1984. The geology of the South Shetland Islands: VI. Stratigraphy, geochemistry and evolution. British Antarctic Survey, Scientific Reports, 87, 85pp. Storey, B.C., Garrett, S.W., 1985. Crustal growth of the Antarctic Peninsula by accretion, magmatism and extension. Geol. Magazine 122, 5–14. Taylor, S.R., McLennan, S.M., 1985. The Continental Crust: its Composition and Evolution, Blackwell, Oxford (328pp).

Tectonic Map of the Scotia Arc, 1985. Sheet BAS (Misc.) 3, 1:3,000,000. British Antarctic Survey, Cambridge. Thomson, J.W. (compiler), 1981. Alexander Island. Geological map, 1: 500000, BAS 500G series Sheet 4, British Antarctic Survey, Cambridge. Thomson, J.W., Harris, J. (compilers), 1981. Southern Graham Land. Geological map, 1: 500000, BAS 500G series Sheet 3, British Antarctic Survey, Cambridge. Thomson, J.W., Harris, J. (compilers), 1982. Northern Palmer Land. Geological map, 1: 500000, BAS 500G series Sheet 5, British Antarctic Survey, Cambridge. Thomson, M.R.A., Pankhurst, R.J., Clarkson, P.D., 1983. The Antarctic Peninsula—a late Mesozoic–Cenozoic arc (review). In: Oliver, R.L., James, P.R., Jago, J.B. (Eds.), Antarctic Earth Science. Australian Academy of Science, Canberra, pp. 289– 294. Tomlinson, J.S., Pudsey, C.J., Livermore, R.A., Larter, R.D., Barker, P.F., 1992. Long-range sidescan sonar (GLORIA) survey of the Antarctic Peninsula Pacific margin. In: Yoshida, Y., Kaminuma, K., Shiraishi, K. (Eds.), Recent Progress in Antarctic Earth Science, Terrapub, Tokyo, pp. 423–429. Vanneste, L.E., Larter, R.D., 1995. Deep-tow boomer survey on the Antarctic Peninsula Pacific margin: an investigation of the morphology and acoustic characteristics of Late Quaternary sedimentary deposits on the outer continental shelf and upper slope. Am. Geophys. Union, Antarct. Res. Series 68, 97–121. Vaughan, D.G., Doake, C.S.M., 1996. Recent atmospheric warming and retreat of ice shelves on the Antarctic Peninsula. Nature 379, 328–331. Weaver, S.D., Saunders, A.D., Tarney, J., 1982. MesozoicCenozoic volcanism in the South Shetland Islands and the Antarctic Peninsula: geochemical nature and plate tectonic significance. In: Craddock, C. (Ed.), Antarctic Geoscience, University of Wisconsin Press, Madison, pp. 263–273. Zielinski, U., Gersonde, R., 1997. Diatom distribution in Southern Ocean surface sediments (Atlantic sector): implications for paleoenvironmental reconstructions. Palaeogeogr. Palaeoclimatol. Palaeoecol. 129, 213–250. Zwally, H.J., Comiso, J.C., Gordon, A.L., 1985. Antarctic offshore leads and polynyas and oceanographic effects. American Geophysical Union, Antarct. Res. Series 43, 203–225.