Quaternary glacial and climate history of Antarctica
Published in: J. Ehlers & P.L. Gibbard (eds.) 2004:
Quaternary Glaciations - Extent and Chronology, Part III. Elsevier, pp. 3-43.
2. Pre-Quaternary glacial history of Antarctica
3. The Pleistocene in Antarctica - a brief overview
4. Late Quaternary glacial development in Antarctica
5. Methods for reconstructing Antarctic glacial history since the Last Glacial Maximum (LGM)
6. Last Glacial Maximum in Antarctica
7. Deglaciation and Holocene environmental development around Antarctica
8. Summary and conclusions on the glacial history since the LGM
The Antarctic Ice Sheet, containing 25-30 x 106 km3 of glacial ice (Drewry et al., 1982; Vaughan 2000), is the largest glacial system on Earth. The western part of the present Antarctic ice sheet (West Antactic Ice Sheet, WAIS), west of the Transantarctic Mountains (Fig. 1), is largely marine-based, currently very dynamic and discharges huge volumes of ice to floating ice shelves through large ice-stream systems. To the east of the Transantarctic Mountains the current East Antarctic Ice Sheet (EAIS) is predominantly terrestrial, relatively stable and has limited discharge into surrounding ice shelves (Fig. 2). The WAIS and EAIS contain about 3.3 x 106 km3 and 26 x 106 km3 of ice, respectively (Drewry, 1983; Denton et al., 1991). The Antarctic ice sheet, potentially containing 57-66 m of sea-level equivalent (Denton et al., 1991; Vaughan 2000), is a critical factor in regulating, modifying and forcing the global climate and oceanographic system. During most of the late Cenozoic the Antarctic ice sheet has driven global eustasy and deep-ocean circulation, and acted as a regulator of global climate (Andersson, 1999). However, the influences of Antarctic Ice Sheet fluctuations in the Quaternary history of global climate, after initiation of major glaciations in the Northern Hemisphere, are not yet well understood. More than 98% of the Antarctic continent is today covered by glacier ice, and the potential on land for obtaining high-resolution geological data pertaining to its glacial history is poor. A fundamental question is what caused the glacial fluctuations observed in the records? Antarctic glaciers respond both to global sea level fluctuations, mainly controlled during the Quaternary by developments in the Northern Hemisphere, and to Southern Hemisphere climate changes. A good understanding of the Late Quaternary glacial and climate history of Antarctica will constrain the contribution of Antarctic ice to the global sea-level- and marine oxygen-isotope records, and is important for understanding the relative timing of climate changes between the polar hemispheres (Denton et al., 1989; Clapperton & Sugden, 1990; Andrews, 1992; Colhoun et al., 1992; Moriwaki et al., 1992; Quilty, 1992; Blunier et al., 1998; Steig et al., 1998).
Fig. 1. Map of Antarctica and its continental shelf
Studies of Late Quaternary climate changes in Antarctica have been focused on ice-core (e.g. Jouzel et al., 1987; Ciais et al., 1994; Blunier et al., 1998, Thompson et al., 1998; Steig et al., 1998, 2000; Petit et al., 1999; Masson et al., 2000) and marine records (e.g. Elverhři, 1981; Anderson & Bartek, 1992; Shipp & Andersen, 1994; Licht et al., 1996; Bentley & Anderson, 1998; Anderson, 1999; Domack et al., 1991, 2001), as a consequence of the scarcity of chronologically well constrained geological data on land. However, the last two decades have seen increasingly more sophisticated evidence from ice-free areas in Antarctica, based on glacial stratigraphical and morphological investigations, studies of lake sediment and moss-bank cores, and of fossil penguin rookeries (reviews in Ingólfsson et al. (1998) and Bentley (1999)). Current knowledge primarily concerns the ‘postglacial’ developments, i.e. from Late Wisconsinan and Holocene, because those geological archives, although few and far between, are the best preserved.
The purpose of this paper is to review the current knowledge of the glacial and climate history of Antarctica. Since large-scale glaciations there are not restricted to the Quaternary Era, the paper will first present a brief summary of the pre-Quaternary history, then proceed with a general overview of the Pleistocene record, and finish with more detailed reviews of the Late Quaternary (including the Holocene) glacial and climatic development, with focus on the Antarctic Peninsula and the Ross Sea/Victoria Land regions. The focus of the review is on the terrestrial record of the glacial and climate history of Antarctica, with reference to the marine record as given in excellent reviews by e.g Anderson (1999) and Anderson et al. (2002).
Fig. 2. Drainage map of the Antarctic Ice Sheet. Modified after Anderson et al. (2002)
The Antarctic ice sheet has existed for about 35-40 million years, since the mid-Tertiary, and it is widely believed that in East Antarctica it reached continental proportions by the latest Eocene-Early Oligocene (Hambrey et al., 1989; Birkenmayer, 1987, 1991; Prentice & Mathews, 1988; Barrera & Huber, 1993; Barrett et al., 1991; Denton et al., 1991). The nature and timing of the initial glaciation of Antarctica is not well known, and the evolution and nature of the WAIS remains controversial (Wilson, 1995). Existing reconstructions of pre-Quaternary Antarctic ice volumes and ice-extent rely heavily on interpretations of Ocean Drilling Program (ODP) data from offshore cores, such as oxygen isotope (d18O) concentrations, ice-rafted detritus (IRD) concentrations, sediment type and clay mineralogy, microfossil assemblages and occurrences of hiatuses (Fig. 3). The geological literature on Antarctic pre-Quaternary glacial history and palaeoceanography is vast, and the following review relies heavily on overviews by Denton et al. (1991), Kennett & Hodell (1993) and Anderson (1999).
Fig. 3. The ODP record of Tertiary glaciations in Antarctica
2.1. Onset of Antarctic glaciation.
The onset of glaciation in Antarctica in the mid-Tertiary was probably related to the breakup of Gondwana, pole-ward drift of Antarctica and the development of ocean passages around the continent (Kennett, 1977). Plate-tectonic and palaeoceanographic isolation of the Antarctic continent successively led to cooling and glaciation. Macro- and microfossil evidence from Late Cretaceous and Paleocene deposits on James Ross Island and Vega Island (Fig. 5) suggest that climate there was warm to cool temperate and humid (Askin, 1992). ODP results from the Queen Maud Land margin suggest temperate to subtropical Late Cretaceous climate (Kennett & Barker, 1990), and ODP-data from northern Weddell Sea suggests bottom and surface waters there were relatively warm throughout most of Paleocene (Kennett & Stott, 1990; Robert & Kennett, 1994). There is no conclusive evidence for the existence of a large ice sheet on Antarctica during the Late Cretaceous or early Tertiary. Fluctuations in global sea level during late Paleocene might, however, record ice-sheet fluctuations in interior East Antarctica (Anderson, 1999).
Deep-sea oxygen isotope records show a steady increase in d18O concentrations during the Eocene, which have been interpreted to signify buildup of ice-sheet in Antarctica (Prentice & Mathews, 1988; Denton et al., 1991; Abreu & Anderson, 1998). Seismic records of a major unconformity related to glacial diamictons in ODP drill cores from the East Antarctic continental shelf provide strong evidence for an ice sheet on East Antarctica by Late Eocene to Early Miocene (Fig. 3) (Barron et al, 1991; Andersen, 1999). Birkenmajer (1988, 1991) described Eocene glacial deposits from King George Island (Fig. 5). These probably relate to a local glaciation rather than an extensive West Antarctic glaciation, since other data from that region suggest relatively warm Eocene conditions (Askin, 1992).
There is strong evidence for an ice sheet on East Antarctica during the Oligocene (Denton et al., 1991; Hambrey et al., 1991), and ice spread into the western Ross Sea by the Late Oligocene (Hambrey, 1993; Wilson et al, 1998). On West Antarctica there existed at least mountain glaciers and localised ice caps (Birkenmajer, 1998; Anderson, 1999), but the existence of an Oligocene WAIS has not been proven.
2.2 Miocene – Extensive glaciations in both East and West Antarctica
Deep-sea oxygen isotope records show stepwise increase in d18O concentrations during the Miocene and into the Pliocene, which have been interpreted to signify step-like build-up of ice volumes on Antarctica (Miller et al., 1991). Results of ODP drillings, seismic records and terrestrial stratigraphical records provide strong evidence for the presence of large ice sheets on both West and East Antarctica during the Miocene (Hambrey et al., 1991; Andersson, 1999). Miocene glacial and glaciomarine strata recorded in drill sites on the Ross Sea continental shelf and in the westernmost Ross Sea are interbedded with meltwater deposits and diatomaceous oozes, indicating shifts from temperate to sub-polar or polar climates throughout the Miocene (Andersson, 1991). The recent Cape Roberts Project provides evidence for the expansion of polythermal glaciers from the Transantarctic Mountains towards and beyond the Cape Roberts drill-site (Powell et al., 1998), as well as subsequent ice recession during the younger part of the Miocene. Larter & Barker (1989, 1991) and Bart & Anderson (1995) provided data to suggest that ice had expanded out on the Antarctic Peninsula continental shelf by Middle or Late Miocene, and there is also onshore evidence for large Miocene ice caps in the Antarctic Peninsula (Birkenmajer, 1988). Abreu & Anderson (1998) suggested that that the WAIS was probably grounded mostly above sea level, and therefore contributed to Miocene eustatic sea level changes.
2.3 Pliocene and the transition to the Pleistocene: controversies on ice volumes, ice sheet stability and continental climate
The deep-sea record from around Antarctica suggests significant Pliocene sea surface temperature fluctuations, and the oxygen isotope and eustatic records indicate that considerable ice volume fluctuations may have occurred (e.g. Prentice & Fastook, 1990; Denton et al., 1991; Dowsett et al., 1996; Flemming & Barron, 1996). The continental seismic shelf records from the Ross Sea region and Antarctic Peninsula continental shelf show a number of shelf-wide unconformities bounding till sheets, indicating high-frequency grounding events, taken to represent waxing and waning of continental glaciers during the Pliocene (Alonso et al, 1992; Bart & Anderson, 1995). There is, however, no clear evidence, in the form of delta deposits or river-valleys, showing significant meltwater discharge in connection with those oscillations (Anderson, 1999). Core data from western (near-shore) Ross Sea shows meltwater deposits interbedded with tills (Barrett et al., 1992).
The Pliocene continental record is fragmentary and difficult to interpret in terms of climate and glaciation development. Consequently it has led to considerably different opinions as to the Miocene-Pliocene-Pleistocene stability of the Antarctic Ice Sheet (Wilson, 1995), and the nature, timing, extent and evolution of the ice sheets. Reconstructions of the history of the EAIS range from drastic ice-volume fluctuations, with the ice sheet fluctuating between extensive collapse and very extensive build-up and mountain overriding (Barrett et al., 1992; Barrett, 1996), to suggestions that it has existed close to its present configuration for the past c. 14 million years (Shackleton & Kennett, 1975; Sugden et al., 1993; Denton et al., 1993). Proponents of these two incompatible views are often referred to as `dynamicists´ and ‘stabilists’, respectively (Wilson, 1995). The greatest controversy focuses around evidence for a major deglaciation in East Antarctica and different interpretations of the fossil evidence from the so-called `Sirius Group´ (Mercer, 1972, 1981; Mayewski, 1975). It consists mainly of tillites, which occur above, 2000 m a.s.l. in the Transantarctic Mountains. The tillites suggest that some time during the Pliocene the Transantarctic Mountains were covered by an ice sheet considerably thicker than the present (Mayewski, 1975; Denton et al., 1991). However, included in the Sirius sediments are palaeosols and plant remains, both leaves and in situ roots of southern beech, Nothofagus sp., at c. 1800 m a.s.l., suggesting a much-warmer-than-present situation. One major point of controversy is the occurrence of Late Pliocene-Early Pleistocene diatoms in the strata. Do they originate from interior basins of East Antarctica, implying a non-glacial situation there, or have they been reworked (windblown) into the strata? There are numerous other aspects to this discussion (for reviews see Denton et al., 1991; Wilson, 1995; Barrett, 1997; Anderson, 1999), and it has not been resolved if the ice sheet underwent such massive changes in ice volumes as for interior East Antarctica to become ice free, alternating with very extensive glacial overriding of the Transantarctic Mountains. There is also discussion on the Pliocene climate; how warm or cold conditions do the Nohtofagus remains represent? Hill & Trustwell (1993) proposed that the Nothofagus vegetation suggested summer temperatures of up to +8oC whilst Francis (1995) suggested that could be compared with the modern willow-shrub vegetation of the Arctic tundra, signifying conditions with annual mean temperatures below freezing, permafrost conditions, but seasonal moisture supply. This implies considerably warmer Pliocene interglacial conditions than at present, and raises the question whether the present polar conditions became established first in the Early Pleistocene, after 2.5 Ma (Webb & Harwood, 1991)?
The Pliocene history of the WAIS is scantily known. Reconstructions are heavily dependent on the East Antarctic records, i.e. whether or not they encompass a major deglaciation of interior East Antarctica or not (Denton et al., 1991). Based on seismic records from the Antarctic Peninsula continental shelf there were repeated fluctuations in West Antarctic ice volumes (Bart & Anderson, 1995) during the Pliocene-Pleistocene.
The combined continental and shelf evidence suggests that during Late Pliocene and perhaps Early Pleistocene glacial episodes, ice sheets advanced to the edge of the continental shelf around Antarctica, with ice extent similar to or greater than during the last (Wisconsinan) glaciation, and that those glaciations were interrupted by interglacial episodes that were warmer than the present. Denton et al. (1991) suggested that the deep-sea dO18 record could be interpreted to indicate that the overall Antarctic ice volume was never significantly less than today during the Pliocene.
3. The Pleistocene in Antarctica – a brief overview
The continental Pleistocene record of glacial and climatic events in Antarctica is fragmentary, at best. In East Antarctica, outcrops with Pleistocene deposits are confined to few localities in the Transantarctic Mountains and at some few-and-far- between coastal sites. Data from the sites in the Transantarctic Mountain have been interpreted to record a major cooling in the Early Pleistocene (Denton et al., 1986; Denton et al., 1989). Lithostratigraphical and sedimentological evidence from Pleistocene drift deposits, as well as geomorphological evidence in the form of trimlines and glacial striae at high altitudes above sea level indicate that the EAIS overrode the Transantarctic Mountains during several episodes of extensive glaciations (Denton et al., 1991). The West Antarctic continental record is equally fragmentary, but available stratigraphical and geomorphological evidence suggest that the WAIS was considerably thicker than at present on several occasions during the Early and Middle Plestocene, and probably grounded on the shelf edge (Denton et al., 1991; Anderson, 1999). The continental data corroborates with seismostratigraphic profiles and core data from the shelf areas of the Antarctic Peninsula, Weddell Sea and the Ross Sea. These data are interpreted to show numerous stacked till sheets, interbedded with thin and discontinous glaciomarine and meltwater sediments, indicating subglacial and grounding zone-proximal deposition (McKelvey, 1991; Shipp et al., 1994; Bart & Anderson, 1995).
The available data strongly suggests that the Antarctic ice sheet was considerably thicker than present on several occasions during the Pleistocene, with circum-Antarctic ice grounded out on the continental shelf edge. The debate concerning when relatively stable polar conditions were established in Antarctica does not concern the Pleistocene; most workers agree that a major cooling was at least established in Antarctica by the Pliocene-Pleistocene transition (Anderson, 1999). The controlling mechanism for the waxing and waning of the Antarctic ice sheet is not well understood: isotope records from deep-sea sediment cores off Antarctica (e.g. Mackensen et al., 1989; Grobe & Mackensen, 1992) and long Antarctic ice cores (e.g. Jouzel et al., 1987; Petit et al., 1997) show strong correlations with the global isotope record (Prell et al., 1986). This has been taken to indicate that there are linkages between Northern and Southern Hemisphere ice sheet developments during the Quaternary. Denton et al. (1991) concluded that there is an overall in-phase relationship between glacial cycles in the Northern and Southern hemispheres, perhaps driven by falling and rising global sea levels regulated by build-up and decay of Northern Hemisphere ice sheets (Hollin, 1962; Thomas & Bentley, 1978).
4. Late Quaternary glacial development in Antarctica
The late Quaternary stability and fluctuations of the marine-based ice sheet in West Antarctica have been debated during the past decades. Mercer (1968) suggested that the WAIS had repeatedly collapsed during Pleistocene interglacials, and Scherer et al. (1998) suggested, on the basis of the existence of marine sediments below present-day West Antarctic ice streams, that the WAIS had collapsed at least once during late Pleistocene. The oceanic d18O record suggests that between 150-120 ka BP there was less global ice volume than at present, and it has been suggested that collapse of the WAIS caused global sea level to be 5-6 m higher than present during Marine Isotope Substage 5e (Mercer, 1978). If this signifies the collapse or decay of the WAIS (Mercer, 1968), or if the water-source was the Greenland ice (Emiliani, 1969; Cuffey & Marshall, 2000), or it came from melting of the Northern Hemisphere ice sheets and partial melting of the Antarctic ice sheet (Denton et al., 1971; Hughes, 1987), remains unresolved.
5. Methods for reconstructing Antarctic glacial history since LGM
The terrestrial evidence for the extent of LGM ice around Antarctica comes from a number of archives:
|Extensional and directional evidence on land: evidence of more extensive ice cover than today is present on almost every ice-free lowland area around Antarctica and its outlying islands, in the form of glacial drift, erratics and striations. Directional evidence in general suggests ice flow offshore towards the shelf areas around Antarctica.|
|Evidence on ice thickness and ice movements at the glacial maximum: these include ice-abraded ridge crests, trimlines, striations, erratics and glacial drift on nunataks and coastal mountains. In general these suggest considerable thickening of LGM ice in the coastal regions compared to the present.|
|Lithostratigraphical data: good stratigraphical sections are extremely rare on Antarctica. High permafrost table prevents digging or excavating, and there is a lack of natural exposures in an environment where eroding rivers are almost totally lacking and extensive sea-ice cover and perennial ice-foot shelter most beaches from erosion. The stratigraphical data available, from natural sections and boreholes, mostly reflects environmental development subsequent to the last deglaciation.|
|Raised beaches: raised beaches around Antarctica can generally be regarded as isostatic fingerprinting of earlier expanded ice volumes compared to present. Raised marine beach deposits that post-date the last major glaciation in Antarctica have been described from many ice-free areas on the continent and nearby islands. Because of the late deglaciation of coastal areas and still extensive glaciation of the continent, the elevation of Holocene marine limit (ML) is much lower than corresponding limits generally are in the Northern Hemisphere. The highest ML described on Antarctica is from Horseshoe Island in Marguerite Bay, off the Antarctic Peninsula, at 55 m a.s.l. (Hjort & Ingólfsson, 1990). The ML on the outer coast around Antarctica is generally below 30-35 m a.s.l., whereas it usually is below 15 m a.s.l. in coves and bays more proximal to the ice sheet (Ingólfsson et al., 1998; Berkman et al., 1998). The age of the Holocene ML around Antarctica is not synchronous, but reflects differences in the timing of regional and local deglaciation. At most sites, however, it is constrained by minimum dates between 7-5 ka BP.|
5.2 Marine geological evidence for former ice-extent
Archives containing data on former ice-extent (Anderson, 1999; Anderson et al. 2002) include:
· Shelf bathymetry; glacial troughs, submarine valleys, moraine ridges, drumlins, flutes and other glacial lineations. These delineate the drainage of glaciers and show that the shelf areas have been shaped by erosion and deposition below and in front of moving outlet glaciers and ice streams.
· High-resolution seismic records; glacial unconformities, sediment cover, grounding-zone wedges. These records signify the extent of glacial erosion and subsequent deposition on the shelf.
· Sediment cores; sedimentological and petrographic analyses for identifying tills and glaciomarine sediments. The tills are first-order data on former ice extent, and 14C-dates from glaciomarine sediments provide constraining minimum dates for deglaciation of the shelf areas.
5.3 Ice core evidence on glacial history
Antarctic ice cores provide long, low-resolution proxy records of ice volume changes for the past c. 450,000 years (Petit et al., 1997), and inferred information on summer temperatures and ice thickness, for example (Lorius et al., 1985, 1993). The long Vostok record shows a strong correlation with the global isotopic record, which has been taken to indicate strong linkages between Northern and Southern Hemisphere ice sheets (Petit et al, 1997). There are, however, large discrepancies between Late Quaternary glacial and climate history as reconstructed from geological evidence from coastal terrestrial or offshore records, and ice core records (Jouzel et al., 1987; Petit et al., 1997; Ingólfsson et al., 1998, Ingólfsson & Hjort, in press; Anderson, 1999). This summary focuses on the stratigraphical terrestrial and offshore geological record as the most reliable proxies for the glacial history of Antarctica.
5.4 Chronological control
The age control on glacial and climatic events since the LGM in Antarctica is primarily through radiocarbon dating. Marine materials such as mollusc shells, marine mammal bones, penguin remains (bones, guano-debris) and snow petrel stomach oil or nest deposits are one source of datable materials. Terrestrial and lacustrine materials like mosses, lake sediment bulk samples, microbial mats, aquatic mosses and algal flakes have also been widely dated for constraining environmental changes in time. Both marine and terrestrial/lacustrine materials often yield ages that appear too old in comparison with the conventional terrestrial-based radiocarbon time-scale (cf. Björck et al., 1991a; Gordon & Harkness, 1992; Andrews et al., 1999; Domack et al., 2000). A requirement for understanding the dynamics of the Antarctic glacial system is to have a reliable chronology for both terrestrial and marine materials.
Dating terrestrial materials - Radiocarbon dates on terrestrial materials are mainly from peat deposits, primarily mosses in moss-banks on the islands off the Antarctic Peninsula (e.g. Fenton, 1980; Birkenmajer et al., 1985; Björck et al., 1991a, 1991b) and bulk sediments, microbial mats, aquatic mosses and algal flakes from Antarctic lake sediments and deltas (e.g. Stuiver et al., 1981; Pickard & Seppelt, 1984; Pickard et al., 1986; Zale & Karlén, 1989; Mäusbacher et al., 1989; Schmidt et al., 1990; Ingólfsson et al., 1992; Björck et al., 1993, 1996; Doran et al., 1999; Denton & Hughes, 2000).
Samples of moss-bank peat on Elephant Island (Fig. 5) have been found to give some of the most reliable aradiocarbon ages in Antarctica (Björck et al., 1991a, 1991b), and are thus optimal for constructing a radiocarbon chronology for environmental changes in the Antarctic Peninsula region for the past 5000 years. There is a balance between atmospheric carbon content and the intake of carbon by the mosses; old groundwater or carbon from the bedrock will not influence the carbon content of the mosses and contamination by down-growth of roots from plants living on the surface is minimal. There are difficulties in correlating the peat sequences to other archives (glacial stratigraphical sections and lake sediment sequences), but Björck et al. (1991c, 1991d) were able to correlate between moss-bank deposits and lake sediments on the Antarctic Peninsula region, using tephrostratigraphy, thus gaining an important chronological control for the lake sediment archives.
There are a number of sources of contamination when dating bulk sediments, microbial mats, aquatic mosses or algal flakes from Antarctic lake basins, often resulting in ages that are too old (e.g. Adamson & Pickard, 1986; Stuiver et al., 1981; Squyres et al., 1991; Björck et al., 1991a; Melles et al., 1994; Zale, 1994):
a) Old groundwater or an input of glacial meltwater depleted in 14C, contaminating the submerged flora, can be a serious problem. Adamson & Pickard (1986) and Stuiver et al. (1981) found that the correction needed for reservoir effects in freshwater algae is 450-700 years.
b) Reduced gas exchange with the atmosphere resulting from the long (in extreme cases perennial or decadal) duration of the ice coverage in lake- and marine environments may lead to much older radiocarbon ages. This effect was described by Weiss et al. (1979) for the Weddell Sea and by Melles et al.,(1997) for the marine basins (epishelf lakes) of Bunger Hills (Fig. 1). In the latter paper, modern reservoir effects of more than, 2000 years were estimated, considerably higher than the marine reservoir effect of 1300 years estimated for the Vestfold Hills area (Adamson & Pickard, 1986).
c) Contamination by the marine reservoir effect through input from sea mammals and birds to lake basins (Björck et al., 1991a, Zale, 1994).
d) Supply of old carbon from soils or weathered carbon-bearing rocks. Stuiver et al. (1981) described two radiocarbon dates from the same delta bed in southern Victoria Land, one from terrestrial algae, the other from a well preserved valve of the scallop Adamussium colbecki. The shell date was 5050±50 BP (corrected by 1300 radiocarbon years for marine reservoir age) and the algae 5930±200 BP, which Stuiver et al. (1981) thought could reflect contamination by carbonate from local marble bedrock. Doran et al. (1999) concluded that dates of algae that grew on proglacial lake bottoms may require corrections because of the relict dissolved inorganic carbon introduced into the lake by inflow from glacier margins, but that this was not a problem when dating near-shore microbial mats in deltas fed by meltwater streams.
e) Continuous erosion of lake bottom surface sediments resulting from bottom-freezing in winter, or oxidation of these surface sediments due to periods of desiccation, are processes which could lead to erroneous radiocarbon dates (Björck et al., 1991a).
f) Recycling of old carbon in stratified Antarctic lakes (Squyres et al., 1991).
g) Longevity of organisms may also play a role. Some Antarctic freshwater- and terrestrial algae can survive long periods of desiccation and repeated freeze-thaw cycles (Vincent et al., 1993). Cryptoendolithic algae in suitable rock types in Antarctica are thought to have very slow turnover times, on the order of 10 000 to 17 000 years (Nienow & Friedmann, 1993, Johnson & Vestal, 1991). These algae can be released to the ground when the rock erodes and then be blown or washed into other stratigraphical situations.
h) Blanketing snow and ice on perennially frozen lakes could hinder lacustrine sedimentation, causing hiatus between the deglaciation and onset of biogenic deposition, resulting in too young minimum dates for deglaciation (Gore, 1997).
Radiocarbon dates from organic remains (microbial mats, algae, water mosses) in sediment cores sampled from the present Lake Hoare in Taylor Valley, Dry Valleys, southern Victoria Land (Fig. 1), have revealed that there occur large contamination problems (Squyres et al., 1991). These were expressed as very old ages of surface sediments (varying between, 2000 and 6000 radiocarbon years) and samples obtained from deeper in the cores yielded ages similar or younger than the surface material. Squyres et al. (1991) concluded that radiocarbon dates from Lake Hoare sediments were of limited value because of the high degree of contamination, and pointed out that probably the source carbon, which the organisms fix, is old, and that relatively long-term recycling of carbon in the lake could contribute to old apparent ages of radiocarbon dates. A similar explanation is possible for radiocarbon ages of 24 000 and 35 700 BP from the base of marine inlet and lake cores in Bunger Hills, East Antarctica (Melles et al., 1997). In the fresh-water Lake Untersee (Dronning Maud Land, East Antarctica (Fig. 1), a thick, perennial lake ice cover probably leads to a reservoir effect on radiocarbon dates as high as 11 000 years (M. Schwab, personal communication, 1998, quoted in Ingólfsson et al., 1998).
Björck et al. (1991a, 1991d) concluded that the causes of erroneous ages often seem to be a combination of different contamination sources and processes and that great caution is needed when radiocarbon dates on Antarctic lacustrine samples are interpreted and evaluated. A primary control for the reliability of the dates is the stratigraphic consistency in the dated sequence. In addition, the ¶13C-value should always be measured and used to correct the 14C/12C relationship. In some cases, the reported age could be incorrect by hundreds of years without such a correction (Björck et al., 1991a). Dates on aquatic moss samples, extracted from the bulk sediments, appear to be more reliable than dates on the bulk sediments themselves, or on algae. Björck et al. (1991d) found that out of 14 radiocarbon dates from a lake basin on Livingston Island, in the South Shetlands, only three determinations could be judged reliable, after controlling the dates by tephrochronological cross-correlations. Two of these were on aquatic mosses.
Dating marine materials - Radiocarbon concentration in the Southern Ocean is dominated by the upwelling of deepwater from the Northern Hemisphere at the Antarctic Divergence. Deepwater is depleted in 14C, and although mixing with ‘younger’ surface water south of the Antarctic Convergence occurs, marine species which live in those waters have apparent radiocarbon ages that are older than 1000-1200 years (Broecker, 1963; Björck et al., 1991a; Gordon & Harkness, 1992). Other factors which influence spatial and temporal variability in the Antarctic radiocarbon reservoir are inputs of radiocarbon-depleted CO2 from melting ice, regional differences in the upwelling around Antarctica, perennial sea ice cover and local freshwater inputs into nearshore marine basins (Omoto, 1983; Domack et al., 1989; Melles et al., 1994; Melles et al., 1997).
In the geological literature on Antarctica, different authors have taken different approaches to the marine reservoir correction. For example, Sugden & John (1973), Clapperton & Sugden (1982, 1988), Payne et al. (1989), Hansom & Flint (1989) and Clapperton (1990) subtracted 750 years from their Antarctic Peninsula radiocarbon dates, while Barsch & Mäusbacher (1986), working on the South Shetland Islands (Fig. 5) used an envelope of 850-1300 years. Ingólfsson et al. (1992) and Hjort et al. (1997) applied a sea correction of 1200 years, while Pudsey et al. (1994) used a reservoir correction of 1500 years. In East Antarctica, Adamson & Pickard (1986), Colhoun & Adamson (1992a) and Fitzsimons & Colhoun (1995) used a reservoir correction of 1300 years when dealing with the Late Quaternary glacial history in the Vestfold Hills and Bunger Hills areas in East Antarctica, while Hayashi & Yoshida (1994), working in the Lützow-Holm Bay area (Fig. 1), suggested a correction of 1100 years. Verkulich & Hiller (1994) radiocarbon-dated stomach oil deposits in snow petrel colonies in Bunger Hills. They based their chronology for petrel colonisation on conventional radiocarbon dates, but stated that a reservoir correction of 1300 years probably was appropriate. Bird et al. (1991) applied a correction for reservoir effects for marine sediments of 2200 years when studying evolution of marine inlets in the Vestforld Hills oasis, East Antarctica. In the Victoria Land/Ross Sea area Stuiver et al. (1981) and Denton et al. (1989) based their chronology on uncorrected radiocarbon dates. Likewise, Baroni & Orombelli (1991) used conventional dates for their deglaciation chronology for Terra Nova Bay, but calibrated the conventional ages when bracketing a relative sea-level curve for the area. Baroni & Orombelli (1994a) presented both uncorrected conventional- and calibrated radiocarbon chronologies when dealing with the Holocene environmental history of Victoria Land, but Baroni & Orombelli (1994b) based their chronology of Holocene glacier variations in Terra Nova Bay on calibrated radiocarbon dates. Colhoun et al. (1992) used a correction of 1090 years for mollusc dates from the Ross Sea area, while Licht et al. (1996) used a reservoir correction of 1200 years for dates from the same area. Gingele et al. (1997) used a reservoir correction of 1550 years when reporting accelerator mass spectrometre (AMS) radiocarbon ages from the Lazarev Sea. Considerable efforts are being devoted to solving problems concerning radiocarbon dating of Antarctic marine sediments, including dating acid-insoluble organic matter and foraminiferal calcite (Andrews et al., 1999; Domack et al., 2000, 2001a).
A number of investigations have assessed which correction of radiocarbon ages of Antarctic marine organisms has to be applied in order to establish a coherent radiocarbon chronology for Late Wisconsinan-Holocene glacial events. Circumantarctic studies generally show the average correction for reservoir age of marine mollusc shells to be 1100-1400 radiocarbon years (Yoshida & Moriwaki, 1979; Stuiver et al., 1981; Omoto, 1983; Adamson & Pickard, 1986; Björck et al., 1991a; Gordon & Harkness, 1992; Domack, 1992; Berkman, 1994; Berkman & Forman, 1996). Studies of pre-bomb seal, whale and penguin samples have yielded greater variability than the pre-bomb marine mollusc shells, ranging between 915±75 and 1760±55 years (Curl, 1980; Mabin, 1985; Whitehouse et al., 1987;, 1989; Baroni & Orombelli, 1991; Gordon & Harkness, 1992), suggesting that longevity and ecology of different species, as well as what material (flesh, bone, feathers, guano) is dated can significantly influence the correction factor required (Mabin, 1986; Baroni & Orombelli, 1991). Berkman & Forman (1996) suggested that reservoir corrections of 1300±100 years (molluscs), 1424±200 years (seals) and 1130±130 years (penguins) should be applied to Antarctic marine organisms.
This paper adopts 1300 radiocarbon years as the best estimate for a circumantarctic correction for all radiocarbon-dated marine organisms, for the sake of comparing glacial histories of the different areas. All radiocarbon ages of fossil marine organisms given in the text have been corrected by that amount, no matter which reservoir correction or calibration was originally made by the authors cited as source of the data. All ages are in uncalibrated radiocarbon years BP, reported as kilo-annum before present (ka BP). At the same time it should be stressed that there are still large uncertainties in the Antarctic marine reservoir effect.
Fig. 4. Reconstruction of grounding line positions of the Antarctic Ice Sheet at LGM. The oldest constraining ages for initial ice retreat are given. Modified after Anderson et al. (2002)
The overall extent of ice cover in Antarctica during LGM is not well known and some existing reconstructions are controversial (reviews in Denton et al., 1991; Ingólfsson et al., 1998; Anderson et al., 2002). One maximum reconstruction suggests that the peripheral domes of the Antarctic Ice Sheet were 500-1000 m thicker than at present and that ice extended out to the shelf break around most of Antarctica (Denton, 1979; Hughes et al., 1981; Clark & Lingle, 1979; Denton et al., 1991; Zhang, 1992). Other reconstructions indicate less ice extent. Mayewski (1975) maintained that the WAIS was only somewhat, if at all, larger than it is today. Data from East Antarctica have been interpreted as indicating that ice either did not extend to the shelf edge (Colhoun & Adamson, 1992a; Goodwin, 1993) or that ice extent was insignificantly further than today in some areas (Hayashi & Yoshida, 1994). A recent study by Anderson et al. (2002) found that the East and WAISs have not advanced and retreated in concert, and whereas the LGM WAIS advanced for the most part to the outer shelf, the EAIS did not expand to the continental shelf edge during the LGM (Fig. 4).
The timing of the last maximum glacial expansion around Antarctica is likewise poorly known, and though most authors assume it to have coincided with the timing of lowest global sea levels during Marine Oxygen Isotope Stage (MIS) 2, around, 20-18 ka BP, circum-Antarctic field data suggest considerable regional differences in glacial maximum extension between >20-14 ka BP (Stuiver et al., 1981; Domack et al., 1995; Anderson et al., 1992; Kellogg et al., 1996; Ingólfsson et al., 1998; Hall & Denton, 1999, 2000; Denton & Hughes, 2000; Denton & Marchant, 2000). Maximum ages constraining ice-sheet grounding in the western Ross Sea are >24 ka BP (Denton & Hughes, 2000) and maximum extension of ice in the western Ross Sea area has been dated to c. 23-17 ka BP (Stuiver et al., 1981; Anderson et al., 1992; Kellogg et al., 1996; Licht et al., 1996; Denton & Hughes, 2000). Hall (1997) and Hall & Denton (2000) have recently dated the maximum extent of ice in Taylor Valley, southern Victoria Land, to 14.6-12.7 ka BP. In East Antarctica, on the coast of MacRobertson Land and at Prydz Bay (Fig. 1) the last maximum ice extent occurred before 17 ka BP and 10.7 ka BP, respectively (Domack et al., 1991a, Taylor & McMinn, 2001). It has been suggested that the last maximum ice extent in the Antarctic Peninsula region occurred later than 30 ka BP (Sugden & Clapperton, 1980) and prior to 14-12 ka BP (Banfield & Anderson, 1995, Domack et al., 2001a). Recent interpretations, however, suggest that the most recent Wisconsinan ice sheet maximum there may pre-date the, 20-18 ka BP LGM: Bentley & Anderson (1998) suggested that ice volumes were considerably greater before 35 ka BP compared to after 35 ka BP, and likewise, Anderson & Andrews (1999) suggest that ice extent towards the continental slope in the Weddell Sea occurred considerably earlier than 26 ka BP. Hjort et al. (1997, in press) likewise argue that the maximum ice extent since the last interglacial (MIS 5e) probably occurred before MIS 3 in the western Weddell Sea. Berkman et al. (1998) argued, on the basis of circum-Antarctic finds of fossil marine shells of ages at or beyond the limit of radiocarbon dating, that MIS 3 was an important interstadial period around Antarctica, but good stratigraphic and chronological control for a middle Wisconsin interstadial event is still lacking. A reconstruction of late Quaternary sea-ice distribution around Antarctica suggests that during LGM (MIS 2) the Antarctic winter sea-ice was extended towards the present marine polar front zone (Antarctic Convergence), whereas summer sea-ice was expanded into the area of the present winter sea-ice edge (Gersonde & Zielinski, 2000).
Fig. 5. Location map for the Antarctic Peninsula, showing minimum ages for ice retreat.
6.1 LGM-extent of ice in the Antarctic Peninsula region – terrestrial evidence
The Antarctic Peninsula (Fig. 2 and Fig. 5) ice sheet is a part of the marine-based WAIS, where sea level is a major control on ice volume. Sugden & Clapperton (1977) suggested, on the basis of bathymetric data showing glacial troughs incised into the submarine shelf, that during LGM a number of ice caps formed on the South Shetland Islands, separated from the Antarctic Peninsula ice sheet by the deep Bransfield Strait. According to their reconstruction, the main control for ice extension was sea level, and the modern -200 m shallow submarine platforms around the islands and along the Antarctic Peninsula roughly coincide with the ice extent. There is no evidence that the South Shetland Islands were overridden by ice from the Antarctic continent during the LGM (John, 1972, Sugden & Clapperton, 1977), as suggested by Hughes (1975) and Hughes et al. (1981).
Fig. 6. Crystalline erratics from the Antarctic Peninsula on northern James Ross Island
(Photograph: Ólafur Ingólfsson 1989)
Evidence of more extensive ice cover than today exists all along the Antarctic Peninsula in the form of ice-abraded ridge crests at high altitudes, striated bedrock on presently ice-free islands, erratics (Fig. 6) and thin till deposits, as well as raised beach and marine deposits (John & Sugden, 1971; Sugden & John, 1973; Curl, 1980; Clapperton & Sugden, 1982, 1988; Ingólfsson et al., 1992), whereas ancient glacial trimlines are rarely described. Systematic attempts to map altitudinal evidence of LGM ice-thickness in the Antarctic Peninsula region have been made on its southern- and central parts (Carrara, 1979, 1981; Clapperton & Sugden, 1982; Waitt, 1983; Bentley & Anderson, 1998; Bentley, 2000) and in the South Shetland Islands (John & Sugden, 1971; Sugden & John, 1973; Curl, 1980). Carrera (1979, 1981) and Waitt (1983) concluded that striations and erratics on nunataks on the southern Antarctic Peninsula showed that some time in the past they were overridden by a considerably thickened ice sheet, at least 500 m thicker than today (Waitt, 1983; Bentley & Anderson, 1998). Bentley et al. (2000) suggested that in the southern and central parts of the Antarctic Peninsula, the Wisconsinan ice sheet some time prior to 35 ka BP had expanded to form two ice domes and probably merged with an expanded WAIS in the Weddell Sea. Based on cosmogenic exposure ages from erratics, they suggested ice-sheet lowering in the southern part of the Antarctic Peninsula, indicating deglaciation, already by c. 16 ka BP. Bentley & Anderson (1998) and Bentley (1999) reconstructed the LGM configuration of an Antarctic Peninsula ice sheet, with ice flowing out towards the shelf edges from an ice-divide stretching N-S along the axis of the peninsula.
A number of additional investigations interpret former ice-thickness on the northern part of the Antarctic Peninsula region and the islands in the western Weddell Sea (e.g. Rabassa, 1983; Ingólfsson et al., 1992; Hjort et al., 1997). Hjort et al. (1997, in press) presented evidence suggesting that presently ice-free areas on James Ross, Vega and Seymour Islands in the NW Weddell Sea had, at least once, been overridden by glaciers advancing from SW-W, from the Antarctic Peninsula, depositing tills and crystalline erratics up to 300-400 m a.s.l. Hjort et al. (in press) suggested that this period of extensive glaciation was contemporaneous with Bentley & Anderson’s (1998) LGM-grounding line position some, 200 km further east, but was probably older than 30 ka BP. They suggested that the ice flowing off the peninsula during the LGM was thinner, channeled through the straits and troughs, and did not attain much greater thickness there than c. 150 m a.s.l.
Raised beaches on the Antarctic Peninsula can generally be regarded as isostatic remnants of earlier expanded ice volumes compared to present. Raised marine beach deposits, which post-date the last major glaciation in the Antarctic Peninsula region, have been described from most ice-free areas on the peninsula and nearby islands (e.g. John & Sugden, 1971; Curl, 1980; Birkenmajer, 1981; Hjort & Ingólfsson, 1990; Ingólfsson et al., 1992; Lópes-Martinez et al., 1996; Hjort et al., 1997). They imply expanded ice volumes prior to the early-mid Holocene, compared to the present. Since there are as yet no well-constrained relative sea-level curves for any part of the region, the history of relative sea-level changes cannot give more than rough constraints on possible ice volumes and deglaciation history. In the northern part of the region the postglacial marine limit (ML) generally lies at altitudes between 15-20 m a.s.l., with the highest ML on James Ross Island at 30 m a.s.l. (Hjort et al., 1997). There it is dated to c. 7.5 ka BP. The ML rises towards south, and is highest on Horseshoe Island in Marguerite Bay, at c. 55 m a.s.l. (Hjort & Ingólfsson, 1990). The Horseshoe Island marine limit is not absolutely dated, but is constrained by minimum ages for the deglaciation of Marguerite Bay as being older than 7-6 ka BP (Kennedy & Anderson, 1989; Harden et al., 1992; Pope & Anderson, 1992; Emslie, 2001). The north-south gradient in altitude of the marine limit agrees with the conclusion of Bentley and Anderson (1998) and Anderson et al. (2002) that there were greater LGM ice volumes in the central and southern part of the region than in the northern part.
Fig. 7. LGM paleo-drainage and grounding line positions off the Antarctic Peninsula
Drawn after Anderson et al. (2002)
6.2 LGM extent of ice in the Antarctic Peninsula area – marine evidence
Ice extent on the western continental shelf - Kennedy and Anderson (1989) and Anderson et al. (1991a) suggested that grounded ice had extended across the shelf areas off Marguerite Bay (Fig. 1) during the LGM, based on a prominent glacial erosion surface identified. Seismic studies by Bart and Anderson (1996) showed that this unconformity extends to the shelf break. Hiemstra (2001) recognized grounded-ice-to-recessional-ice scenario for Marguerite Bay, based on microscopic analyses of glacigenic core sediments. Anderson et al. (2002) presented swath bathymetric records that are interpreted to show drumlins, flutes and striations in the landward portions of the Marguerite Bay trough. They also described mega-scale glacial lineations in the seawards portion of the trough, extending into a prominent grounding zone wedge on the mid-shelf.
Pudsey et al. (1994) published side-scan sonar records offshore Anvers Island that show glacial flutes on the inner shelf there. Larter and Vanneste (1995) described relict subglacial till-deltas on the Antarctic Peninsula outer shelf, off Anvers Island, which they interpreted as products of progradation at former ice-stream grounding lines. They took this discovery to imply that grounded ice on the outer shelf had been low profile and fast flowing. Constraining minimum age for the till-deltas of 11.3 ka BP was provided by 14C dates from stratigraphically overlying glacial marine sediments (Pudsey et al., 1994). A study by Domack et al. (2001a) of glaciomarine sediments on top of a diamicton in the Palmer Deep also suggests that the shelf close to Anvers Island was covered by ice prior to ca. 11 ka BP.
Banfield and Anderson (1995) recognized a widespread glacial unconformity extending to the continental shelf break off the Bransfield Strait, with prominent ridges occurring above the unconformity, within troughs on the middle shelf. Bentley and Anderson (1998) suggest that these ridges are moraines, marking the LGM grounding line (Fig. 2). A recent swath bathymetric documentation by Canals et al. (2000), showed that during a past glacial maxima a major ice stream had flowed off the north-western side of the Antarctic Peninsula, depositing an >100 km long, elongated, ice-molded sediment body. They argued that it had been deposited below >1000 m thick ice, under conditions of very fast ice-flow. The age of these deposits is not yet constrained by 14C-dates, and a LGM age is still circumstantial. Canals et al. (2000) proposed that during LGM a separate ice-dome had existed over Brabant and Anvers Islands, with its eastward draining entering the Gerlache Strait and flowing northwards into the Bransfield basin.
Bentley and Anderson (1998) and Anderson et al. (2002) reviewed available data on LGM grounding line positions. They concluded that LGM glaciers flowing west from the Antarctic Peninsula had generally grounded on the middle-outer shelf, at water depths of ca. 400 m (Fig. 4 and Fig. 7), and that Antarctic Peninsula ice had met with local ice cap over the South Shetland Islands.
Ice extent in the north-western Weddell Sea - Because of the inaccessibility, the geological evidence from the western Weddell Sea is scarcer than from west of the peninsula. Sea ice has long effectively prevented the acquisition of marine geological information south of 65oS, but some is forthcoming after successful cruises to Prins Gustav Channel and Larsen Ice Shelf in, 2000 (Pudsey & Evans, 2001; Pudsey et al., 2001; Domack et al., 2001b). Shelf bathymetric data indicate submarine troughs, which have drained large outlet glaciers from the Antarctic Peninsula into the Weddell Sea (Sloan et al., 1995). Anderson et al. (1992) and Bentley and Anderson (1998) suggested, on the basis of high-resolution seismic records and sediment cores, that Prince Gustav Channel and adjoining bays contain only little sediment. They interpreted this to indicate recent glacial erosion. Multibeam records from the inner shelf show flutes and other glacial lineations (Domack et al., 2001b). Seismic profiles and sediment cores suggest a widespread erosional unconformity extending all the way to the shelf break (Anderson et al., 1992), and Anderson et al. (2002) described a prominent grounding zone sediment wedge above the unconformity on the mid-shelf. They suggested that the unconformity marked the extension of Antarctic Peninsula ice sheet, and that it had grounded on the outer continental shelf, some 200 km east of the peninsula, during recent glaciation (Fig. 4 and Fig. 7). However, the age of the unconformity is uncertain since it is not constrained by any available dates (Anderson et al., 2002) but Bentley and Anderson (1998) and Anderson et al. (2002) correlated it to the LGM.
6.3 LGM extent of ice in East Antarctica
The extent of ice during the LGM in East Antarctica is not well known and partly controversial. This is mainly due to lack of detailed marine geological surveys aimed at constraining the LGM configuration of ice on the East Antarctic continental shelf (Anderson et al., 2002) A maximum reconstruction extends ice to the shelf break (Denton, 1979; Hughes et al., 1981; Stuiver et al., 1981) while a minimalistic view suggests only modest glacial oscillations and that some coastal oases remained ice-free (Omoto, 1977; Hayashi & Yoshida, 1994; Adamson et al., 1997). Denton et al. (1991) suggested that surface elevations in the interior did not differ significantly from the present during the LGM, but that there was a considerble thickening of marginal domes in East Antarctica. Other data from East Antarctica suggest that ice extended off the present coast, but not all the way to the shelf edge (Colhoun & Adamson, 1992a; Fitzsimons & Domack, 1993; Goodwin, 1993). Anderson et al. (2002) reviewed the data used for constraining the extension of the LGM EAIS, and concluded that ice had at most advanced to positions on the middle or mid-outer continental shelf, not to the shelf break, and in some areas there had hardly been any LGM ice expansion compared to present ice configuration (Fig. 4).
Eastern Weddell Sea and Dronning Maud Land – The Weddell Sea is a major discharge centre for ice, draining both the West and East Antarctic ice sheets as well as the eastern side of the Antarctic Peninsula (Fig 1 and Fig. 2). The extent of the WAIS on the Wedell Sea continental shelf is poorly constrained, and attempts to radiocarbon date glacial marine sediments from the Weddel Sea shelf have been hampered by paucity of sediment cores and general absence of dateable carbonate materials in the cores (Anderson et al., 2002). A number of piston cores taken from the Weddell Sea continental shelf penetrate tills below few metres of glaciomarine sediments (Anderson et al., 1980; Fütterer & Melles 1990; Anderson et al., 1991b). A till provenance study by Anderson et al. (1991b) suggested that the till on the shelf had been deposited by a glacier coming from East Antarctica. High-resolution seismic profiles record a mid-shelf grounding zone wedge that rests above a glacial unconformity that extends to the outer shelf (Elverhři & Maisey, 1993). Kristoffersen et al. (2000) described two submarine ridge complexes on the shelf, parallel to the shelf edge, and reported a radiocarbon date of c., 19 ka BP constraining grounded ice extension to the mid-shelf area at the LGM.
Seismic data, sedimentological data and radiocarbon dates from the Weddell Sea have been interpreted as suggesting that a grounded ice sheet extended to the shelf break off Dronning Maud Land prior to c. 21 ka BP (Elverhři, 1981; Melles, 1991). Grobe et al. (1993) presented a sedimentary record from the eastern Weddell Sea, where they concluded that the Antarctic Ice Sheet had extended to the shelf edge during the LGM. Weber et al. (1994) studied submarine fan deposits in front of the continental slope in the south-eastern Weddell Sea. They suggested that prior to c. 20 ka BP, laminated sediments were deposited under conditions of very high sedimentation rates (250 cm ka-1) when grounded ice sheet was in the proximity at the shelf edge. Anderson & Andrews (1999) concluded that tills on the continental shelf and extending towards the shelf break in the Weddell Sea pre-dated MIS 2, and recorded a glaciation older than the the range of radiocarbon dating. Bentley & Anderson (1998), Anderson & Andrews (1999) and Anderson et al. (2002) suggested a major deglaciation event predating the LGM, occurring some time before 26-25 ka BP, and implied that ice had not extended towards the continental shelf after that (Fig. 4).
Lützow-Holm Bay - Ice-free areas fringing Lützow-Holm Bay (Fig. 1) occur on a number of islands and headlands, the largest being 61 km2. Signs of glaciation occur as discontinuous glacial drift and erratics, striated bedrock surfaces and streamlined glacial bedforms, and extensive raised beaches (Yoshikawa & Toya, 1957; Hirakawa et al., 1984; Yoshida, 1983; Igarahi et al., 1995; Maemoku et al., 1997; Sawagaki & Hirakawa, 1997). Omoto (1977) and Hayashi & Yoshida (1994) concluded that ice retreat from the Lützow-Holm Bay area occurred before 30 ka BP, and that it had not been ice-covered during the LGM. Igarashi et al. (1995) concluded that the last major deglaciation in the area dated back to the last interglacial, on the basis of the occurrence of old (42-33 ka BP) fossil shells in raised beach deposits. Similarly, Maemoku et al. (1997) also described old (46-32 ka BP) fossil shells from Lützow-Holm Bay in undisturbed raised beach deposits, and interpreted that to indicate that the area had not been overridden by LGM glaciers. The position of the ice-margin during the LGM is unknown (Igarashi et al., 1995, Maemoku et al., 1997). While offshore troughs are identified and interpreted to reflect erosion by expanded systems, no offshore marine geological or geophysical studies have been carried out (Anderson et al., 2002). Raised beaches of Holocene age, 7-2 ka BP (Igarashi et al., 1995) do, however, occur. The Holocene marine limit is at 25 m a.s.l., indicating considerable regional isostatic response to decreased ice volume. Maemoku et al. (1997) suggested the EAIS might have covered the southern part of Lützow-Holm Bay during LGM.
Taken together, the evidence can be interpreted to suggest that the EAIS did not expand significantly during the LGM in Eastern Weddell Sea and Dronning Maud Land (Anderson et al., 2002; Fig. 4).
Mac Robertson Land, Prydz Bay and Larseman Hills - The geomorphology of the shelf off Mac Robertson Land shows signs of glacial erosion during ice sheet advance, and several moraine-like features could have been formed by glaciers retreating across the shelf (Harris & O’Brien (1996). Harris & O’Brien (1998) place the grounding line of the EAIS on the outer shelf, but due to lacking dating control it remains unresolved if it was there during LGM or earlier. An uncorrected bulk-organic radiocarbon date of marine mud from the mid-outer shelf, from sediments interpreted to have been deposited below a floating ice shelf (Harris & O’Brien, 1998), gave an age of 17,1 ka BP.
It has been suggested that the small (50 km2) oasis of Larsemann Hills in Prydz Bay (Fig. 1) was either wholly or partly ice free during the LGM (Burgess et al., 1994, 1997) or completely covered by, 200-500 m thick continental ice (Gillieson, 1991). Erratic boulders are scattered throughout the Hills, but till deposits, morainic ridges and glacial striae only sparsely occur (Burgess et al., 1994). Domack et al. (1998) concluded that the most extensive glaciation of the Prydz Bay area pre-dated the Northern Hemisphere LGM. They proposed that the last maximum extension of ice there, during MIS 2, reflected an advance of outlet systems and subsystems, reaching the mid-shelf areas, and that it was not a shelf-wide glaciation or wholesale advance of the EAIS. They described uninterrupted sedimentation on the shelf for the past 30 ka.
Vestfold Hills - During LGM the whole 400 km2 of the Vestfold Hills oasis (Fig. 1) was covered by ice (Adamson & Pickard, 1983, 1986). Evidence of glacial overriding include glacial striae, erratics, till deposits and moraine ridges, as well as raised beaches at altitudes below 10 m a.s.l. (Pickard, 1985; Zhang, 1992). The orientation of glacial striae is uniform, showing ice movement towards WNW across the oasis at a time of complete ice coverage (Adamson & Pickard, 1983, 1986). Zwartz et al. (1998) estimated that the LGM ice margin had stood on the continental shelf some 30-40 km outside present margins in the Vestfold Hills oasis, and that it had thinned by 600-700 m at the present coastline since the LGM.
Bunger Hills - The Bunger Hills (Fig. 1) form the most extensive oasis of deglaciated hills and marine inlets in East Antarctica, with a total size of 952 km2 (Wisniewski, 1983). There are signs of extensive glaciation such as discontinuous but locally thick glacial drift deposits, striated bedrock surfaces, roches moutonnées and erratics, as well as extensive raised beaches. Colhoun & Adamson (1992a) and Augustinus et al. (1997) suggested that there was evidence that at some stage in the past the Antarctic ice sheet completely submerged most of the oasis, and that ice flow from south-east to north-west was independent of the local topography. Augustinus et al. (1997) proposed that this extensive glaciation may have pre-dated the LGM. Limited glacial erosion, thick glacial deposits in the north-western part of the hills and the relatively low altitude of the postglacial marine limit (9-7 m a.s.l.) suggested to Colhoun & Adamson (1992a) and Colhoun (1997) that the LGM ice sheet was not very thick and consequently did not extend far onto the continental shelf. However, Melles et al. (1997) found till in 18 sediment cores retrieved from different basins within the oasis. They concluded that probably all the oasis was inundated by glaciers during the LGM.
The Windmill Islands - The Windmill Islands (Fig. 1) at the Budd Coast in East Antarctica, show evidence of an extensive late Pleistocene ice cover including glacial polishing and striae on the gneiss bedrock, roches moutonnées and erratics, as well as raised beaches at altitudes below 32 m a.s.l. (Goodwin, 1993). A shallow veneer of unconsolidated sediments occurs on the islands. Cameron et al. (1959) interpreted this as reworked till, but Goodwin (1993) found subglacially-deposited fine sediments almost totally lacking. He argued that the best indicator for glacial overriding during the LGM were the raised beaches, bearing witness of isostatic rebound in connection with deglaciation. Goodwin (1993) calculated that the Late Pleistocene-early Holocene ice thickness over the Windmill Islands and the inner shelf had been <200 m and <400 m, respectively, and that ice had extended 8-15 km off the present coast. Marine geological surveys of the shelf areas close to the Windmill Islands reveal that the inner shelf is deeply eroded, reflecting glacial scouring, which Harris et al. (1997) suggested supported Goodwin’s reconstruction of an expanded LGM ice. Recent modelling by Goodwin & Zweck (2000), suggests that the LGM ice may have been up to 1000 m thick and extended 40 km out on the continental shelf.
Wilkes Land, Oates Land and the Pennell Shelf – Anderson et al. (1980) suggested that the LGM EAIS had grounded on the outer shelf off Wilkes Land (Fig. 1), and Goodwin & Zweck (2000) suggested that LGM thickness of ice along the margin between Wilkes Land and Oates Land had been 1000 m. The Pennell shelf (Fig. 1) has large glacial troughs extending across the shelf, and sediment cores retrieved from the area penetrate a blanketing diamicton interpreted by Anderson et al. (2002) as till. They concluded that LGM ice was grounded at least to the middle shelf (Fig. 4). A number of radiocarbon dates constrain the ice expansion onto the shelf to have occurred prior to c. 14.3 ka BP.
Fig. 8. The Ross Sea Drift in Explorers Cove, in the mouth of Taylor Valley, Southern Victoria Land, containing kenyte erratics.
(Photograph: Ólafur Ingólfsson, 1994)
6.4 The Ross Sea area and coastal Victoria Land
The Ross ice drainage system comprises about 25% of the surface of the Antarctic Ice Sheet (Fig. 2). Since the dawn of geological research in Antarctica there has been a discussion about the fluctuations of ice in the Ross Sea and Victoria Land in space and time (cf. Stuiver et al., 1981; Denton et al., 1989, 1991; Denton & Hughes, 2000). Reconstructions of LGM ice flowlines for the drainage of the East- and WAISs to the Ross Sea are conflicting (Drewry, 1979; Denton et al., 1989; Clapperton & Sugden, 1990; Kellogg et al., 1996; Denton & Hughes, 2000), but most studies suggest that the Ross Sea embayment was largely filled by a low surface profile, marine-based ice sheet. Glacial drift deposits (Ross Sea Drift), containing kenyte erratics from Ross Island (Fig. 8), show that during the LGM the WAIS thickened and grounded in the Ross Sea and McMurdo Sound. Lobes of ice to were pushed onto coastal southern Victoria Land, leving signs in form of trimlines (Fig. 9) and damming the Dry Valleys (Stuiver et al., 1981; Denton et al., 1989, 1991). Erratics on islands in south-western Ross Sea, at elevations up to 320 m a.s.l., and moraines of the Ross Sea Drift occurring at 350 m a.s.l. in the Transantarctic Mountains indicate ice sheet elevation in the Ross Sea during the LGM (Hall & Denton, 1999). Further north along the Victoria Land coast, major outlet glaciers, for example in Terra Nova Bay, drained the EAIS and coalesced with the marine-based ice in the Ross Sea (Denton et al., 1989, Orombelli et al., 1991). Although reconstructions of ice extent in the Ross Sea during the LGM agree that ice was considerably expanded compared to today (cf. Drewry, 1979; Stuiver et al., 1981; Anderson et al., 1984; Denton et al., 1989, 1991; Clapperton & Sugden, 1990; Orombelli et al., 1991; Kellogg et al., 1996; Licht et al., 1996; Conway et al., 1999; Denton & Hughes, 2000), there remains uncertainty about the maximum position of the grounded ice. Licht et al. (1996, 1999) recognised tills in the western Ross Sea, but only in areas south of Coulman Island (Fig. 1 and Fig. 16). Kellogg et al. (1996) suggested that grounded LGM ice had extended to the shelf break off Cape Adare, while Denton et al. (1989), Anderson et al. (1992), Baroni & Orombelli (1994a), Shipp & Anderson (1994), Licht et al. (1996), Conway et al. (1999) and Shipp et al. (1999) place the LGM-grounding line along the northern Victoria Land coast in the vicinity of Coulman Island (Fig. 4).
Fig. 9. Glacial trimline marked by the upper limit of the Ross Sea Drift, at ca. 350 m a.s.l. on
Hjorth Hill, Taylor Valley, Southern Victoria Land.
(Photograph: Ólafur Ingólfsson, 1994)
Although there are more than 450 radiocarbon dates reported in the literature that relate to the glacial history of the Ross Sea area (Denton & Hughes, 2000), the dating of the LGM-grounding line is not well constrained. Series of dates give maximum and minimum ages for the LGM position on the shelf: Orombelli et al. (1991), Licht et al. (1996) and Denton & Hughes (2000) constrained its age by maximum ages of >24 ka BP, where as a number of cores from sediments inside the LGM grounding line provide minimum ages in the range c. 20-11 ka BP (Anderson et al., 1992; Licht et al., 1996; Domack et al., 1999a). Anderson et al. (1992) dated it off Northern Victoria Land to 17,3 ka BP and Licht et al. (1996) dated it to c., 20 ka BP. Penguin rookeries at Cape Adare and Cape Hallet were probably occupied between >35 ka BP and 17,3 ka BP (Baroni & Orombelli, 1994a). Since they can exist only where there is access to open water in summer, they show that the coastal north-western Ross Sea was free of glacial ice during the LGM.
Marine geological studies in the central and eastern Ross Sea have revealed compelling evidences for LGM ice extending across the shelf towards the shelf edge: Alonso et al. (1992) and Shipp et al. (1999) reported the existence of a young unconformity in the stratigraphy, covered by till (Domack et al., 1999a; Shipp et al., 1999), indicating a major glacial expansion on the shelf. Swath bathymetry and side-scan sonar records show ice-contact features such as mega-scale glacial lineations and grounding zone sedimentary wedges, and the combined sedimentological and geophysical data suggests that LGM ice grounded at the shelf break in the Eastern Ross Sea (Shipp et al., 1999; Anderson et al., 2002). The dating control for the LGM expansion is problematic: Domack et al. (1999a) suggested that the LGM advance was marked by a large hiatus in radiocarbon ages between c. 25.5 and 19.5 ka BP (14C dates on bulk-organic carbon, corrected by surface age of 3200 yrs), whereas unpublished dates from K. Licht (referred in Anderson et al. 2002) indicate that the ice may have been at its LGM position as late as 13.7 ka BP.
6.5 The Amundsen Sea
6.6 Summary of LGM ice extent and age around Antarctica
Most studies suggest that the LGM ice was considerably thicker and more extensive than present during the LGM, with ice extending off the present shore around most of Antarctica (Fig. 4). Some minor coastal areas may have remained wholly or partly ice-free. LGM-ice extended onto the mid-to-outer shelf areas off western Antarctica, Antarctic Peninsula and in the Ross Sea. In East Antarctica, the LGM ice reached its maximum extent at mid-shelf positions in some areas, whereas in other areas it only expanded less than that. The age of the last circum-Antarctic maximum extent of ice is poorly constrained by radiocarbon dates that broadly range between c. 24–14 ka BP, and there is data suggesting that there might have been an earlier (pre-MIS 2) glacial event of considerably more extensive glaciation onto the continental shelf areas.
7. Deglaciation and Holocene environmental development around Antarctica
7.1 The Antarctic Peninsula region
Deglaciation of the shelf and near-shore areas - There is a striking lack of radiocarbon dates constraining the deglaciation from the Antarctic Peninsula shelf area. This is probably partly a result of the lack of datable material in transitional glaciomarine sediments deposited below floating ice shelves in proximity to the grounding lines of the retreating glaciers. The oldest radiocarbon dates (Fig. 5), constraining ice retreat from the outer and middle shelf areas west of the Antarctic Peninsula come from the Bransfield Strait (Banfield & Anderson, 1995) and the shelf areas west of Anvers Island (Pope & Anderson, 1992; Pudsey et al., 1994). These indicate ice retreat prior to 14-13 ka BP and 12-11 ka BP, respectively. Most dates from the inner shelf areas and fjords and bays on the peninsula constrain deglaciation to as late as 8-6 ka BP BP (e.g. Harden et al., 1992; Pudsey et al., 1994; Shevenell et al., 1996; Hjort et al., 2001). According to Harden et al. (1992), glaciomarine sedimentation in the central part of the Gerlache Strait began after 8 ka BP, providing a constraining minimum date for deglaciation and shelf-ice retreat in the strait. Similarly, Shevenell et al. (1996) dated the onset of glaciomarine sedimentation in Lalleman Fjord to c. 8 ka BP. The inner shelf areas around Anvers Island and in Marguerite Bay, as well as George VI Sound, were ice-free at 7-6 ka BP (Clapperton & Sugden, 1982; Kennedy & Anderson, 1989; Harden et al., 1992; Pope & Anderson, 1992; Pudsey et al., 1994; Hjort et al., 2001). Domack et al. (2001a) recently published a Late Wisconsinan-Holocene chronology for environmental changes in the Palmer Deep (Fig. 1) which showed that the inner shelf areas there might have been deglaciated as early as c. 11 ka BP.
The low resolution in the deglaciation chronologies from most of the shelf areas makes it impossible to recognise more than the general pattern, that the outer-and middle shelf areas deglaciated between 14-8 ka BP, while most inner shelf areas, fjords and bays were deglaciated before 8-6 ka BP. Improved chronologies for marine sediment cores, which are currently being developed (Domack et al., 2001a), are the key to better resolution in reconstructions of spatial and temporal patterns of the deglaciation of the Antarctic Peninsula shelf areas. Domack et al. (2001a) recognised a rapid deglacial episode in the Palmer Deep record, characterised by high primary production and iceberg rafting, between c. 11-10 ka BP. This deglacial event occurred concurrently with the Northern HemisphereYounger Dryas Stadial. In the Palmer Deep record, it was followed by a climate cooling between c. 10-8 ka BP.
There is as yet no conclusive data to suggest that ice-retreat from the LGM positions and subsequent deglaciation of the shelf areas were controlled by regional warming. Hence the ice retreat was probably controlled by the global sea-level rise that resulted from meltwater input from the Northern Hemisphere glaciers (cf. Hollin, 1962; Stuiver et al., 1981).
Fig. 10. In situ shells (Laternula elliptica) at The Naze, northern James Ross Island. They were 14C dated to 7.4 ka BP, and provide constraining minimum ages for the deglaciation as well as the age of the marine limit at c. 30 m a.s.l.
(Photograph: Ólafur Ingólfsson 1993)
Deglaciation of presently ice-free Antarctic Peninsula coastal areas - The deglaciation of coastal areas that are currently ice-free is constrained by minimum ages obtained from radiocarbon dating of mollusc shells from raised marine deposits, peat deposits in moss-banks on the islands off the Antarctic Peninsula, organics (bulk organics, microbial mats, aquatic mosses and algal flakes) from lake sediments and remains from abandoned penguin colonies. The oldest radiocarbon dates, on fossil molluscs from raised marine deposits and lake sediments, give a minimum age for initial deglaciation on King George Island as 9-8 ka BP (Mäusbacher, 1991; Sugden & John, 1973). A well-dated lithostratigraphical record from northern James Ross Island (Fig. 5), where glaciomarine and sub-littoral deposits overlie till in coastal sections (Fig. 10), constrains deglaciation there as to before c. 7.4 ka BP (Hjort et al., 1997). According to Zale (1994a), Hope Bay was deglaciated prior to ca. 6.3 ka BP. In the southern part of the peninsula, radiocarbon dated shell fragments from ice-shelf moraines at Two Step Cliffs on Alexander Island provide minimum ages for deglaciation and ice-shelf retreat in George VI Sound of 6.5-5.7 ka BP (Clapperton & Sugden, 1982; Hjort et al., 2001).
A number of studies, exploring primarily lake-sediment archives and moss-bank peats, suggest that once the glaciers were inside the present coastline, glacial retreat and ice-disintegration was slow (Barsch & Mäusbacher, 1986a; Mäusbacher et al., 1989; Zale & Karlén, 1989; Mäusbacher, 1991; Ingólfsson et al., 1992; Björck et al., 1993, 1996a, 1996b; Lopez-Martinez et al., 1996; Hjort et al., 1997): On King George Island, glaciers were at or within their present limits by c. 6 ka BP (Martinez-Macchiavello et al., 1996), on northern James Ross Island prior to 4.7 ka BP (Hjort et al., 1997), but parts of Byers Peninsula on Livingston Island were deglaciated as late as 5-3 ka BP (Björck et al., 1996a). The minimum date for deglaciation on Elephant Island is provided by the onset of moss-bank formation on the steep slopes at Walker Point by 5.5 ka BP (Björck et al., 1991b).
Fig. 11. The large penguin rookery in Hope Bay. Penguins have occupied this site since c. 5.5 ka BP. The earliest Holocene penguin occupation in the Antarctic Peninsula-South Shetland Islands region has been dated to c. 6.5 ka BP in Marguerite Bay.
(Photograph: Ólafur Ingólfsson, 1987)
Penguin colonization and location of rookeries is determined by a number of climate-dependent factors, such as availability of ice-free coastal areas suitable for nesting, absence of persistent ice-foot, access to open water during the nesting season and the availability of food (Baroni & Orombelli, 1994). The timing of penguin occupation at coastal sites is therefore a good proxy for minimum age of deglaciation and climate/sea ice situation similar to the present. Radiocarbon dated organic remains from fossil penguin rookeries in Marguerite Bay suggest initial colonization in the period 6.5-5.5 ka BP (Emslie, 2001), which confirms minimum ages for deglaciation of 7-6 ka BP provided by marine-geological studies. Zale (1994a, 1994b) studied the history of Adélie penguin occupation at the large Hope Bay rookery (Fig. 11). He dated first penguin colonization there at ca. 5.5 ka BP, about 0.8 ka subsequent to the deglaciation of Hope Bay. The oldest dated occupation of penguins at King George Island is 5.8-5.3 ka BP (Barsch & Mäusbacher, 1986b).
It can be concluded that the transition from Wisconsinan glacial to Holocene interglacial conditions in the Antarctic Peninsula region was broadly completed by c. 6 ka BP. Interglacial conditions are here defined by glacier volumes and ice configuration similar to or less than at present, lake-sediment accumulation in ice-free basins, moss-bank growth on the islands off the peninsula and penguin occupation of coastal rookeries.
Mid-Holocene glacial re-advances - There are indications from a number of sites for mid-Holocene glacial expansions: Mäusbacher (1991) found evidence of increased glacial activity between 5-4 ka BP on King George Island, supporting Sugden & John (1973) who had suggested glacial expansion here after 6 ka BP. Hansom & Flint (1989) recorded a glacial readvance on Brabant Island some time after 5.3 ka BP, and the post 5.7 ka BP expansion of the George VI Sound ice shelf is well documented (Fig. 12) (Sugden & Clapperton, 1980, 1981; Clapperton & Sugden, 1982; Hjort et al., 2001). A mid-Holocene glacial readvance on northern James Ross Island, initially described by Rabassa (1983), was constrained by further stratigraphical and chronological evidence to have culminated in an at least 7 km ice-advance around 4.6 ka BP (Hjort et al., 1997). Zale (1994) described a set of distinct moraines in Hope Bay, at the northern tip of the Antarctic Peninsula, which he suggested marked a glacial oscillation at c. 4.7 ka BP.
Fig. 12. Ice pressure ridges where George VI Sound ice shelf transgresses Alexander Island.
Before c. 6.5 ka BP the George VI Sound was ice free.
(Photograph: Ólafur Ingólfsson, 2000)
Yoon et al. (2000) identified cold waters with extensive sea-ice cover between ca. 6.2 and 4 ka BP, from a multiproxy study of gravity cores from fjord margin sediments on King George Island, which correlates well with Mausbacher´s (1991) suggestion of increased glacial activity at that time. The marine records from the Palmer Deep and Lallemand Fjord do not recognize a mid-Holocene glacial event (Shevenell et al., 1996; Taylor et al., 2000; Domack et al., 2001b; Brachfeld et al., 2002), although Shevanell et al. (1996) identified a mid-Holocene cool event in the Lallemand Fjord record. Hjort et al. (1997) suggested that the glacial advance might be a regional response to increased precipitation, due to warming and increased cyclonic activity in mid-Holocene. Domack et al. (1991) suggested a similar interpretation for mid-Holocene glacial advances in East Antarctica.
Fig. 13. The synthesis of Björck et al. (1996a) for the mid-late Holocene climate development in the Antarctic Peninsula region.
The climatic pattern for the last 5 ka is quite similar between the South Shetland Islands, in the maritime climate west of the peninsula, and James Ross Island in the colder and dryer climate in the western Weddell Sea, and this Björck et al. (1996a) suggested might indicate that the primary factor controlling the climatic variations is the strength of the high-pressure cell over the Antarctic ice sheet. A paleoenvironmental record from lake sediment cores from the maritime Signy Island, in the South Orkney Islands archipelago, is interpreted to indicate a Holocene climate optimum between c. 3,8-1,3 ka BP (Jones et al., 2000). This demonstrates that the climate optimum persisted for a longer period at Signy Island than at sites more proximal to the Antarctic continent.
The marine record from Lallemand Fjord (Fig. 5) recognises a broad climatic pattern (Shevenell et al., 1996; Domack et al., 1995; Taylor et al., 2001), with an earlier relatively warm period between ca. 7.5-5.8 ka BP and a climatic optimum, reflected by high productivity in the fjord, between 4.2 and 2.7 ka BP. Domack et al. (2001a) defined a more prolonged Holocene climatic optimum in the Palmer Deep record, between ca. 8-3 ka BP, characterized by enhanced biological productivity and minimum IRD concentrations. The difference between the terrestrial record and the Palmer Deep record with regard to the duration of the Holocene climate optimum might be a reflection of paucity of terrestrial paleoclimatic data extending beyond 6-5 ka BP, or be partly caused by oscillations in humidity, reflected in the terrestrial record but not picked up by the marine record.
The palaeoclimatic synthesis of Björck et al. (1996a) recognised that after c. 3 ka BP the climate became characterised by cold and dry conditions, which persisted until c. 1.5 ka BP. Thereafter the climate was somewhat warmer and more humid, but still cold compared to the situation during the climatic optimum. Several investigations suggest that glaciers have oscillated and expanded in the Antarctic Peninsula region during the past c. 2.5 ka (John & Sugden, 1971; John, 1972; Sugden & John, 1973; Curl, 1980; Birkenmajer, 1981; Clapperton & Sugden, 1988; Zale & Karlén, 1989; Clapperton, 1990; López-Martínez et al., 1996). For example, in the South Shetlands, Curl (1980), Birkenmajer (1981), Clapperton & Sugden (1988) and Björck et al. (1996b) found evidence for glacial expansions, in the form of readvance moraines covering Holocene raised beaches, and all suggested that this advance had coincided with the Little Ice Age (LIA) glacial advances in the Northern Hemisphere. Lichenometric dating, using Rhizocarpon geographicum thalli, date the advances to c. A.D. 1240, 1720 and 1780-1822 (Curl, 1980; Birkenmajer, 1981). Whalebone found on a moraine at the front of Rotch Dome on Livingston Island dates the advance here to after c. 0.3 ka BP, that is after c. AD 1650 (Björck et al., 1996b).
Barcena et al. (1998) studied siliceous microfossils from the Bransfield Strait. They found significant increase in sea-ice taxa after 3 ka BP. Fabres et al. (2000) presented a palaeoclimatic sedimentary record from the Bransfield Strait, extending back to almost 3 ka BP. They recognise a distinct neoglacial cooling as well as a LIA cold pulse. The marine record from Lallemand Fjord (Shevenell et al., 1996) suggests decreased productivity in the fjord after 2.7 ka BP due to cooling. Diatom abundances reflect more extensive and seasonally-persistent sea ice after c. 2.7 ka BP (Shevanell et al., 1996) and after c. 0.4 ka BP an ice shelf advance into the fjord is recorded. Levanter et al. (1996) presented a multiproxy record from the Palmer Deep, and identified both short-term (200 yr) and long-term (2500 yr) cyclicity in the variability of stratigraphic parameters, that they interpreted to signify responses to climatic fluctuations. Levanter et al. (1996) and Kirby et al. (1998) suggested, on the basis of organic productivity changes and magnetic lithostratigraphic changes, respectively, that ca. 2500 BP marked the termination of the Holocene climate optimum as reflected in the sediment core record. Domack et al. (2001a) define a Neoglacial cooling in the Palmer Deep record, starting at c. 3 ka BP and lasting until about 100 years ago. The Neoglacial conditions are characterised by greater concentrations of IRD, decreasing sediment accumulation and decreased biological productivity. The timing of the Neoglacial onset, as reflected in the Palmer Deep record, agrees well with other marine (Smith et al., 1999) and terrestrial (Björck et al., 1996a; Ingólfsson et al., 1998) evidence. Fabres et al. (2000) presented a palaeoclimatic sedimentary record from the Bransfield Strait, extending back to almost 3 ka BP. They also recognise a distinct Neoglacial cooling as well as a LIA cold pulse.
Fig. 14. Tentative syntesis for environmental changes in the Antarctic Peninsula region since the LGM.
7.2 Deglaciation and Holocene development in East Antarctica
Dronning Maud Land – There are only very few constraints on the ice-volume decrease in Dronning Mud Land. Lintinen & Nenonen (1997) reported radiocarbon dates of stomach oil deposits from nesting sites of snow petrels (Pagodroma nivea) giving 7400 BP as the minimum age for deglaciation of the southern Vestfjella nunataks and the lower altitudes of Heimefrontfjella (Fig. 1). There has been continuous avian occupation since then, but details of any glacial and climate fluctuations are lacking.
Schirmacher and Untersee Oases - The Schirmacher Oasis (Fig. 1) is a small (34 km2) ice-free area located c. 100 km south of the Lazarev Sea. The last glacial submergence of the Schirmacher Oasis is reflected in a sparse coverage of morainic material, fractured rocks, striated surfaces and roches moutonnées in many places in the oasis (Richter & Bormann, 1995). Radiocarbon dating of lake sediments suggests that deglaciation of the Schirmacher Oases began before 3,5 ka BP (M. Schwab, personal communication, 1998; cited in Ingólfsson et al., 1998). Hence, deglaciation here may have occurred significantly later than in Untersee Oasis to the south (see below), and also later that on the Lazarev Sea shelf to the north, where it was under way by 9,5 ka BP (Gingele et al., 1997).
The Untersee Oasis (Fig. 1) forms the eastern rim of the unglaciated Wohlthat Massif, surrounding Lake Untersee. At least four generations of moraines have been differentiated in the Untersee Oasis (Stackebrand, 1995). They were interpreted as glaciation and deglaciation stages, representing a succession from a total submergence by the inland ice towards the present ice-free setting. The last total ice coverage probably pre-dates the LGM (Hiller et al., 1988). This is indicated by Late Wisconsinan 14C-ages on snow petrel stomach oil deposits from high-altitude sites in the Untersee Oasis.
No indications are known from Untersee nor Schirmacher Oases for significant Holocene re-advances of ice margins. However, grain-size distribution and radiocarbon ages of sediments on the Lazarev Sea shelf could indicate that ice tongues to the east may have re-advanced some time between 8-2 ka BP (Gingele et al., 1997).
MacRobertson Land Evidence from marine cores from the Mac Robertson Land shelf suggest that deglaciation of the outer shelf occurred some time prior to 10,8-10 ka BP, and that afterwards that open-marine deposition prevailed (Harris & O´Brien, 1998; Taylor & McMinn, 2001). Harris et al. (1996) suggest open marine conditions on the mid-shelf by c. 7 ka BP and within the inner shelf basins by c. 5 ka BP. The diatom record suggests a succession of mid-late Holocene high-productivity events, associated with a climate optimum occurring broadly within the interval 5,8-0,4 ka BP. Sedwick et al. (2001) suggested that grounded ice on the outer shelf started to retreat before 9 ka BP, and that the front of a floating ice shelf had retreated to the inner shelf by c. 5 ka BP. They recognised six episodes of increased accumulation of biogenic material since 10,8 ka BP, which they interpreted as the result of enhanced diatom production over the outer shelf, possibly related to climatic warm periods.
Larseman Hills and Prydz Bay - A number of sediment cores have been retrieved from lakes in the Larsemann Hills. There are serious problems with radiocarbon dates from the sediments, expressed as reversed stratigraphical age successions (Burgess et al., 1994). While most dated samples give mid-late Holocene ages for the deglaciation, a single date gave 9,4 ka BP and another 25 ka BP. The Amery Ice Shelf (Fig. 1), the largest in East Antarctica, has been present throughout the Holocene since retreat from the mid-shelf LGM terminal moraines in Prydz Bay (O´Brian & Harris, 1996; Domack et al., 1998; Harris et al., 1998; O´Brien et al., 1999). The Holocene grounding line retreat is constrained by a number of radiocarbon dates to having occurred successively since c. 11 ka BP (O´Brien & Harris, 1996; Domack et al., 1999b).
Vestfold Hills - The last deglaciation of Vestfold Hills (Fig. 1) has been determined by radiocarbon dates on molluscs from raised marine deposits and moraines, on marine algal sediments from numerous lake basins, and on fossil mosses (Adamson & Pickard, 1983,1986; Pickard, 1985; Pickard & Seppelt; 1984, Pickard et al., 1986; Bird et al., 1991; Bronge, 1992). A prominent feature of the Vestfold Hills is its numerous lakes. There are about 300 lakes, with fresh to saline and hypersaline water. Many were formerly marine inlets and became isolated by isostatic uplift following the glacial retreat. The oldest radiocarbon dates, giving minimum ages for the initial deglaciation and the incursion of marine water onto coastal areas, as well as for the initiation of aquatic moss growth, are between 8,6-8,4 ka BP (Pickard & Seppelt, 1984; Fitzsimons & Domack, 1993; Bird et al., 1991; Fulford-Smith & Sikes, 1996). According to Adamson & Pickard (1986) and Pickard et al. (1986), ice retreat thereafter was slow or stepwise, averaging 1-2 m/yr, with, 20% of the land area exposed by 8 ka BP, 50% by 5 ka BP, and the ice margin reaching its present position in the last 1000 years. Relative sea-level stood at the marine limit in Vestfold Hills, at c. 9 m a.s.l., at c. 6 ka BP (Zwartz et al., 1998). Fitzsimons & Domack (1993) and Fitzsimons & Colhoun (1995) maintain that the southern part of the Vestfold Hills deglaciated prior to 8,6 ka BP.
Domack et al. (1991b) found evidence on the shelf for a middle Holocene readvance of floating ice tongues some time within the interval 7,3-3,8 ka BP. There is no dated evidence for this from the Vestfold Hills, but Adamson & Pickard (1986) suggested that a mid-Holocene glacial advance may be recognised in the form of moraine ridges in the area. A late Holocene ice advance, termed the Chelnok Glaciation, is poorly-dated but probably occurred after 2.5 ka BP and before 1 ka BP (Pickard et al., 1984; Adamson and Pickard, 1986; Zhang, 1992; McMinn et al., 2001). Fitzsimons and Colhoun (1995) found evidence for minor (Ł 500 m) late Holocene (post 0.7 ka BP) ice-marginal fluctuations in the form of a discontinuous series of ice-cored moraine ridges; they may correlate with the Little Ice Age. Pickard et al. (1984, 1986) concluded that the ‘post-glacial’ climate of the Vestfold Hills had been very stable and similar to that of today. Zhang (1992), however, on the basis of marine fossil assemblages and geomorphological criteria, proposed that a climatic optimum had occurred here between 6,2 and 3,7 ka BP. In the geomorphological record, Pickard (1982) found evidence that the predominant wind direction had been stable for the past 4 ka.
Björck et al. (1996a) re-interpreted the lake sediment evidence of Pickard et al. (1986) in terms of palaeoclimate development for the Vestfold Hills. They suggested that the deglaciation prior to 4.7 ka BP occurred under an arid and cold environment (low lake levels, high salinity), followed by a relatively warm and humid climate between 4,7-3,0 ka BP which caused intense melting of dead ice in the vicinity of the lakes. Massive input of fresh water into the basins caused salinity to fall and lake levels to rise. After 3 ka BP the climate again turned arid and cold, with decreased fresh water input and lowered lake levels. Björck et al. (1996a) thus conclude that a late Holocene (4.7 - 3.0 BP) climate optimum occurred in the Vestfold Hills. Roberts & McMinn (1996, 1999) used transfer functions for the reconstruction of past lakewater salinity from fossil diatom assemblages. They found that since 5,2 BP, one lake studied had undergone cycles of varying salinity. The data indicate a long period with relatively low salinity, possibly indicating warmer and wetter conditions in the time interval c. 4.2 - 2.2 ka BP. McMinn et al. (2001) studied marine sediment cores from the Vestfold Hills area. They recognised a mid-late Holocene warm period (>3.6 - 2.5 ka BP) and a significant cooling and increase in sea-ice cover after 2.5 ka BP. They correlate the cooling with the Chelnok Glaciation event, and suggest it can also be identified in lake cores (Roberts & McMinn, 1997, 1999) and marine cores (McMinn, 2000; McMinn et al., 2001), as well as in the geomorphological record (Pickard et al., 1984; Adamson & Pickard, 1986).
Bunger Hills - Numerous radiocarbon dates on total organic carbon from lacustrine and marine sediments indicate that the initial deglaciation of the southern part of the Bunger Hills oasis (Fig. 1) dates back to between 10 - 8 ka BP (Bolshiyanov et al., 1990, 1991; Melles et al., 1994; Melles et al., 1997; Kulbe et al., 2001), and radiocarbon dates of stomach oil deposits at nest sites of snow petrels show occupation as early as 9.5 ka BP (Bolshiyanov et al., 1991; Verkulich & Hiller, 1994). The glacial retreat was partly controlled by the rise of sea level, which caused a relatively rapid collapse of the ice-sheet margin (Colhoun & Adamson, 1992a; Verkulich & Melles, 1992). Melles et al.,(1997) conclude that the first phase of deglaciation was also associated with climatic warming, indicated by high diatom concentrations in the sediments and a large meltwater input to the basins. Kulbe et al. (2001) recognised the period after initial deglaciation, between c. 9 - 7 ka BP, to have been relatively warm. The southern part of the Bunger Hill Oasis deglaciated at c. 9 ka BP, and synchronous accumulation of subaquatic sediments in basins and stomach oil deposits on ice-free hills precludes the proximity of any large ice masses since c. 8.5 ka BP (Kulbe et al., 2001). Between 9 - 7.7 ka BP the sea flooded all the major inlets in Bunger Hills (Colhoun & Adamson, 1991, Melles et al., 1997; Kulbe et al., 2001). The breeding colonies of snow petrels expanded continuously, following ice retreat and the down-wasting of dead ice (Verkulich & Hiller, 1994), with the most intense phases of colonisation between 6.7-4.7 ka BP. Before 5.6 ka BP when the sea stood at the marine limit at between 9-7 m, glaciers were at or behind their present margins (Colhoun & Adamson, 1992a, Colhoun & Adamson, 1992b).
Evidence on the mid-late Holocene glacial and climatic evolution of Bunger Hills is, however, somewhat controversial. Bolshiyanov et al. (1991) suggested that the area was re-glaciated several times during the Holocene, causing damming of tributary valleys with periodical lake sediment deposition. They based their conclusions on lake sediment thickness, as well as on fluctuations in the growth rate of aquatic mosses and algae, and on shifting salinity conditions as reflected by diatom assemblages and the geochemistry of lake sediments. Verkulich & Hiller (1994) found no evidence of any major Holocene glacial advance in the Bunger Hills and Colhoun & Adamson (1992a), Fitzsimons & Colhoun (1995) and Fitzsimons (1997) concluded that ice margins on the southern boundary of Bunger Hills had been fairly stable since the last deglaciation. At the western margin, however, glacier expansions of a few hundred metres resulted in the formation of the so-called Older Edisto Moraines, post-dating 6.2 ka BP (Colhoun & Adamson, 1992a).
Verkulich & Melles (1992), Melles et al. (1997) and Kulbe et al. (2001) have studied sediment cores from freshwater lakes and from marine basins in the Bunger Hills Oasis. Melles et al. (1997) report high and stepwise increasing biogenic production between 4.7 to 2.0 ka BP, taken to indicate increased temperatures and correlated with the Antarctic Peninsula climate optimum at about the same time. Melles et al. (1997) also conclude that the glacial advance forming the Older Edisto Moraines pre-dated the 4.7 ka BP warming. This constrains the Edisto advance to between 6.2-4.7 ka BP, which is coincident with the cooling and glacial readvance in the Antarctic Peninsula area (see above). Roberts et al. (2000) recognised a period between 4-2 ka BP of decreased salinity in lake sediments cores from a Bunger Hill lake, which they thought signified more humid and warmer conditions. Rozycki (1961) and Colhoun & Adamson (1992a, 1992b) interpreted beach morphology to suggest that wave action may have been more important in the mid-Holocene, and that the extent and duration of sea ice has increased in the late Holocene. A recent study by Kulbe et al. (2001) suggest that there was a significant mid-late Holocene climatic variability in the Buger Oasis, with a climate optimum between c. 3,5-2,5 ka BP. It was followed by an abrupt and dramatic cooling, that took place within less than, 200 years.
Finally, Colhoun & Adamson (1992a) described a glacial advance during the last few centuries, leading to the formation of the so-called ‘Younger Edisto Moraines’. Melles et al.,(1997) concluded that the ‘Younger Edisto Moraines’ post-date 1,1 ka BP. Marine shell fragments collected from the moraines dated to 0,2 ka BP (Colhoun & Adamson, 1992a), suggesting a Little Ice Age glacial event.
Windmill Islands - Radiocarbon dating of bulk samples from basal lake sediments provide minimum estimates for the deglaciation of the Windmill Islands (Fig. 1) (Goodwin, 1993). These indicate that the southern part of the islands were deglaciated before 8 ka BP, while the northern islands were only deglaciated some time before 5,5 ka BP. Very little is known about the post-glacial Holocene climate development at Windmill Islands. Goodwin (1993) interpreted the onset of algal growth in the lakes to indicate warmer conditions than at present between 2-1 ka BP. He also found indications of higher lake levels during that period.
There is evidence for a readvance of the Law Dome ice sheet margin (Fig. 1) onto part of the Windmill Islands between 4-1 ka BP (Goodwin, 1996). The overriding advance of the ice margin incorporated frozen coastal sediments from raised beach, lacustrine and proglacial environments together with slabs of marine ice from a palaeo-ice shelf, during the marginal transition from fringing ice shelf to grounded ice sheet. Goodwin (1996) attributed the readvance to a positive mass balance on the Law Dome caused by high precipitation rates during the Holocene.
Fig. 15. Summary of glacial and environmental changes since LGM in Eastern Antarctica.
(Modified after Ingólfsson et al. 1998)
Summary of the East Antarctic data - The oldest deglaciation dates show open marine and lacustrine environments, as well as avian habitats, developing within inner shelf basins and coastal areas between 10 – 8 ka BP (Fig. 15). The deglaciation history of Vestfold Hills is controversial. One reconstruction specifies it as successive and slow and mainly post-dating 8 ka BP. In the other reconstruction, deglaciation is advanced by 8.6 ka BP. There are indications from the Vestfold Hills of a mid-Holocene glacial advance. The combined data from Bunger Hills show an early initial deglaciation phase, and retreat of the continental ice sheet margin to its present position by 10 ka BP. The retreat was coupled with collapse of ice over the marine inlets while the land was still depressed below present sea level around the Pleistocene-Holocene transition. Most of the down-wasting of the dead ice may have occurred successively after 8.8 ka BP. There are indications of a mid-Holocene glacial re-advance; expansion of the Edisto Glacier occurred after 6.2 ka BP and before 4.7 ka BP. The Windmill Islands were successively deglaciated between 8 - 5 ka BP. There are indications from both Vestfold Hills and Bunger Hills of a climate warmer and wetter than the present in the interval c. 4.7 - 2.5 ka BP, and the Windmill Islands record may tentatively be interpreted as indicating the same for the 4-1 ka BP interval. The records from both Vestfold Hills and Bunger Hills strongly suggest Neoglacial cooling after c. 2.5 ka BP. The East Antarctic data also show minor fluctuations of the ice margins over the past few hundred years (Goodwin, 1998), possibly correlative with the Little Ice Age advances in the Northern Hemisphere.
Fig. 16. The Ross Sea-Victoria Land region, with minimum ages for deglaciation. Underlined 14C ages relate to dated grounding line recession.
(Modified after Ingólfsson et al. (1998). Location of the LGM grounding line after Conway et al. (1999) and Anderson et al. (2002))
7.3 The deglaciation of Ross Sea and Victoria Land
Conway et al. (1999), Hall & Denton (2000) and Denton & Hughes (2000) have recently reviewed the retreat of ice in Ross Sea and along the Victoria Land coast. Several radiocarbon dates constrain the marine-based ice sheet retreat in the Ross Sea from the LGM-grounding line after 17 ka BP (Fig. 16). Minimum ages for the deglaciation of the area between 75°S and 76°S come from offshore sediments close to the present Drygalski ice tongue, and are c. 11,4 ka BP (Licht et al., 1996). At the LGM, the Terra Nova Bay region was occupied by coalescing outlet glaciers, draining the EAIS and joining the marine based Ross Ice Sheet (Baroni & Orombelli, 1991; Orombelli et al., 1991). Glacial drift deposits relating to the LGM indicate that the ice surface in the bay was about 400 m above the present sea level. Even though the grounding line of the ice sheet had retreated to south of Terra Nova Bay by 11.4 ka BP, the break-up of the shelf ice was probably considerably delayed. The minimum age for the deglaciation of Terra Nova Bay is some time before 6.2 ka BP, when the coasts there became free of ice and marine habitats could begin to develop (Baroni & Orombelli, 1991; Orombelli et al., 1991). The oldest penguin rookeries in the bay date to 5.8 ka BP.
Denton et al. (1991), Hall (1997) and Hall & Denton (2000) date the presence of Ross Sea Ice Sheet in Explorers Cove, southern Victoria Land, damming lakes along the valley threshold, at 8.3 ka BP. Daes from penguin rookeries on Ross Island and in the McMurdo Sound area provide minimum ages for the deglaciation in the southern Ross Sea. Rookeries at Cape Bird were occupied by 6.8-5.7 BP (Speir & Cowling, 1984; Heine & Speir, 1989). Marine fossils collected from debris bands on the McMurdo Ice Shelf date the grounding line recession of the Ross Sea ice sheet to a position south of Black Island by 6.3 ka BP (Kellogg et al., 1990), which is in line with the penguin data. Radiocarbon dates on fossil marine molluscs from raised beaches along the southern Victoria Land coast, McMurdo Sound and the islands in the south-western Ross Sea, as well as additional dates on penguin occupation in McMurdo Sound, all give minimum ages for the deglaciation as c. 6.3-6.1 ka BP (Stuiver et al., 1981; Speir & Cowling, 1984; Denton et al., 1989; Colhoun et al., 1992; Baroni & Orombelli, 1994a; Kellogg et al., 1996). A relative sea-level curve for the coast north of Explorers Cove indicates unloading of grounded ice by 6.3 ka BP and provides one minimum estimate for the deglaciation of coastal southern Victoria Land (Conway et al., 1999; Hall & Denton, 1999). It coincides well with the oldest radiocarbon date on a bivalve from a core in McMurdo Sound, which gives 6.5 ka BP as the minimum age of grounding line retreat from the Sound (Licht et al., 1996). The marine limit becomes gradually younger southwards, which might indicate lagging of shelf ice recession behind that of the grounding line. The age of the marine limit is 6.2 ka BP in Terra Nova Bay, 5.4 ka BP at Marble Point/South Stream and 5 ka BP in the Explorer’s Cove area (Stuiver et al., 1981; Denton et al., 1989; Orombelli et al., 1991; Berkman, 1997; Hall, 1997).
Bockheim et al. (1989) concluded that Holocene ice-surface lowering of the Hatherton Glacier in the Transantarctic Mountains, corresponding in time with grounding line recession in the south-western Ross embayment, occurred before c. 5.3 ka BP. Numerous radiocarbon dates on marine macrofossils recovered from dirt bands in shelf ice in southern McMurdo Sound indicate that the grounding line had retreated to an unknown position south of Minna Bluff by 2.7 ka BP (Kellogg et al., 1990). Conway et al. (1999) estimated the rate of grounding line retreat to have been c. 120 m y-1 for the past 7.5 ka. At present the Ross Shelf grounding line is located c. 900 km south of McMurdo Sound, still retreating. Conway et al. (1999) suggested that continued recession might lead to the complete disintegration of the WAIS within the present interglacial.
Mid-late Holocene glacier oscillations - There is some information available as to mid- to late-Holocene glacier variations in Victoria Land. After the deglaciation in Terra Nova Bay, the ice shelves entering the bay were less extensive than today (Orombelli et al., 1991, Baroni, 1994, Baroni & Orombelli, 1994b). The ice margins stood 2 - 5 km inside their present margins between 6.2 and 5.3 ka BP. A readvance across raised beaches took place some time after 5.3 ka BP (Baroni & Orombelli, 1994b). There was then a renewed withdrawal phase between c. 1 - 0.5 ka BP, correlated by Baroni (1994) and Baroni & Orombelli (1994b) with the Northern Hemisphere Medieval Warm Period (ca. 1000-1300 A.D.). Moraine ridges containing fossil marine shells then show a glacial readvance in Terra Nova Bay after 0.5 ka BP, which Baroni & Orombelli (1994b) suggested might correspond to the Little Ice Age glacial expansion in the Northern Hemisphere. Möller (1995) described a system of minor, fresh-looking thrust moraines in Granite Harbour, Victoria Land, which he suggested were formed by repeated oscillations during general retreat of the ice front. The youngest thrust moraine post-dates, 1910 A.D., when the English Terra Nova Expedition surveyed the area, and Möller (1995) concluded that the moraine ridge system was formed in connection with a Little Ice Age glacial expansion.
Alpine glaciers in the Dry Valleys probably did not contribute to the Ross Sea Ice Sheet, and at present some of these are at their maximum frontal positions since the LGM (Stuiver et al., 1981). The Wilson Piedmont Glacier in southern Victoria Land was contiguous with the Ross Ice Sheet at LGM (Hall & Denton, 2000) and merged with it north of Explorers Cove. The Wilson Piedmont glacier retreated inside the coast between 5.7 - 5 ka BP, but re-advanced in late Holocene times, and at a number of places it advanced across raised beaches which date back to 5,4-5 ka BP (Nichols, 1968, Stuiver et al., 1981, Denton et al., 1989, Hall, 1997).
Climatic optimum in Victoria Land - The best information on palaeoclimatic development in Victoria Land during the latter part of the Holocene comes from the spatial and temporal distribution of penguin rookeries (Baroni & Orombelli, 1994a). They document the continuous presence of Adélie penguins after c. 7 ka BP, but the greatest diffusion of rookeries occurred between 3.6 - 2.6 ka BP. They termed this interval the ‘penguin optimum’, and concluded that it was a period of particularly favourable environmental conditions. It was followed by a sudden decrease in the number of penguin rookeries shortly after 2.6 ka BP, particularly in southern Victoria Land, attributed to an increase in sea-ice. A mid-late Holocene decrease in sea ice and longer seasons with open water in southern Victoria Land is supported by the observations of Nichols (1968). He described raised beaches, clearly high-energy type, at many sites south of Granite Harbour. At many of these sites today the ice-foot rarely breaks up and recent beaches are of low-energy type. All beaches described by Nichols (1968) are younger than 5.4 ka BP (Stuiver et al., 1981, Berkman, 1997). Abandoned Adélie penguin rookeries, occupied during the penguin optimum, occur at Cape Ross and Marble Point (Baroni & Orombelli, 1994a), where Nichols (1968) described high-energy beach ridges. The raised beach deposits at Marble Point have been dated to 5.4 ka BP (Stuiver et al., 1981, Berkman, 1997). Cunningham et al. (1999) studied the palaeoceanographic changes in the western and west-central Ross Sea from >12 ka BP to the present. They recognised a distinct warm period during the mid-Holocene, c. 6-3 ka BP, and increasingly cooler late Holocene after 3 ka BP.
Studies on lake histories from the McMurdo Dry Valleys of Southern Victoria Land (Doran et al., 1994; Lyons et al., 1998) give different reconstructions of the late Holocene climate in southern Victoria Land. They recognise a relatively mild period in the interval c. 3 - 2 ka BP, followed by a cool, dry period between 2 - 1.2 ka BP and succeeded by a mild period from 1.2 - 0 ka BP. This discrepancy between the records is difficult to understand, but it might be explained by either a combination of microclimatic and topographical variables controlling hydrological, ecological and chemical balances in the lakes, or increased distality of the area to the retreating Ross Ice Shelf. Records of thinning of perennial ice thickness on lakes in southern Victoria Land have been used to substantiate warming and increased summer melting during the past few decades (Webster et al., 1996; Wharton et al., 1992). Meteorological data for the past ca. 50 years show that Antarctic surface air temperature trends towards cooling in the interior and warming in the coastal regions, and Doran et al. (2002) have described ecosystem response in areas affected by this cooling.
Fig. 17. A broad synthesis for glacial and environmental changes in Antarctica since the LGM.
8. Summary and conclusions
Detailed reconstruction of the glacial and climatic history of Antarctica since the LGM is hampered by the scarcity of available archives, low resolution in many datasets and chronological problems. Therefore, any synthesis (Fig. 17) must be regarded as a tentative description of the environmental development. However, the following broad pattern can be recognised:
a) There is evidence of much more extensive ice cover in Antarctica prior to c. 20-14 ka BP. LGM glaciers from the WAIS and the Antarctic Peninsula generally grounded on the middle-outer shelf, at water depths of ł 400 m, and were probably fringed by ice shelves extending over the outer shelf and the shelf break. East Antarctic glaciers expanded to mid-shelf positions, and in some areas there was little or no LGM glacial expansion.
b) Ice retreat from the LGM positions was under way by 17-14 ka BP. The initial ice retreat was probably eustatically-controlled.
c) Deglaciation occurred mainly during the time period >14-6 ka BP. Outer- and middle shelf areas deglaciated between 14-8 ka BP, while most inner shelf areas, fjords, bays and most currently ice-free coastal land areas deglaciated prior to 8-6 ka BP. This suggests that Antarctica could have contributed to global sea-level rise at least until mid-Holocene times (cf. Flemming et al., 1998; Ingólfsson & Hjort, 1999). There are signs in the East Antarctic record (Bunger Hills) that early Holocene glaciation may have been at least partly driven by rising temperatures.
d) There are no signs in Antarctica of a Younger Dryas glacial or cold climatic event.
e) The transition from glacial to interglacial conditions in Antarctica, indicated by ice configuration becoming similar to or less than at present, by the onset of lake sediment accumulation in ice-free basins, by moss-bank growth on the islands off the peninsula, and by penguin occupation of coastal rookeries, was broadly completed by 6 ka BP.
f) Mid-Holocene glacial- and ice-shelf re-advances have been described from a number of sites.
g) Terrestrial palaeoclimatic records, based on stratigraphical variables in lake- and moss-bank archives, as well as proliferation of penguin rookeries, suggest a circum-Antarctic climate optimum occurring broadly in the period 4,5-2.5 ka BP. It is best constrained in the stratigraphical records from the Antarctic Peninsula (4-3 ka BP), Bunger Hills in East Antactica (3,5-2,5 ka BP) and Victoria Land (3,6-2,6 ka BP).
h) Available data suggest that after c. 2.5 ka BP a distinct Neoglacial cooling occurred, and that many glaciers and ice-shelves expanded during the late Holocene. The Ross Ice Shelf probably continued retreating even after 2.5 ka BP and is still in a retreat mode.
i) A rapid and significant warming has occurred in the Antarctic Peninsula region for the past 100 years, which might be causing ice-shelf instability, and coastal areas in the Ross Sea also show signs of warming. The interior East Antarctica is slightly cooling.
The available data suggest that Antarctic deglaciation and Holocene climate development lagged that of the Northern Hemisphere (Hjort et al., 1998; Ingólfsson et al., 1998; Bentley, 1999; Hall & Denton, 1999, 2000). The late Holocene climate optimum, c. 4.5 – 2.5 ka BP is recognised as broadly synchronous environmental changes in the Antarctic Peninsula, East Antarctica and Victoria Land, which supports the notion of a circum-Antarctic climate optimum. This optimum is not recognised in the Antarctic ice cores. The ice-core records on Holocene climate variability from Byrd, Vostok and Taylor Dome differ considerably in terms of temperature trends for the past, 20 ka and through the Holocene (Blunier et al., 1998, Thompson et al., 1998, Petit et al., 1999, Steig et al., 1998, 2000), and cannot be used as proxies for environmental changes in the coastal areas of the continent. A study by Ciais et al. (1994) on Holocene temperature variations in Antarctica, as expressed by data from 6 ice-cores, indicated only subtle temperature variations (Ł±1oC) for the past 10 ka BP. Although that study recognised somewhat warmer-than-present conditions between c. 4 - 2 ka BP, it places the Antarctic Holocene climate optimum between c. 10 - 8 ka BP. The Palmer Deep record (Domack et al., 2001a), on the contrary, marks this as a period of cool climate. A study on 11 ice-core isotope records for the reconstruction of Holocene climate variability in Antarctica (Masson et al., 2000) likewise demonstrated a widespread early Holocene climate optimum (11.5 - 9 ka calibrated ice-core years: c. 10 - 8.5 ka radiocarbon years), in addition to a second period of climate amelioration in the Ross Sea area between 7 - 5 ka cal. yr BP (c. 6 - 4.5 ka radiocarbon years BP). These periods of relatively warm climate are not apparent in the geological stratigraphical record. The ice-core data further suggests an East Antarctic climate optimum between 6 - 3 ka cal. yr BP (Masson et al., 2000), whereas the geological record from Bunger Hill Oasis indicates the period 7.6 - 4.5 ka cal. yr BP as being relatively cold, and the climate optimum occurring between 3.5 - 2.5 ka cal. yr BP (Kulbe et al., 2001).
The large Antarctic system allows for significant regional variability in the Holocene development of climate and glaciation. Ice-core records from the high-altitude inland continental plateau may not respond to environmental changes in the peripheral maritime areas of the system. White & Steig (1998) suggested that the inland-plateau ice cores might not be entirely representative of Antarctic climate development, and that records from far more sites are needed to complete the story. Perhaps the poor fit between the ice-core and geological records is an illustration of just that problem.
Abreu, V.S. & Anderson, B.J. (1998). Glacial eustasy during the Cenozoic: Sequence stratigraphic implications. American Association of Petroleum Geologists Bulletin, 82, 1385-1400.
Adamson, D. & Pickard, J. (1983). Late Quaternary ice movement across the Vestfold Hills, Antarctica. In: Oliver, R.L., James, P.R. & Jago, J.B. (eds), Antarctic Earth Science. Canberra: Australian Academy of Sciences, 465-469.
Adamson, D. & Pickard, J. (1986). Cainozoic history of the Vestfold Hills. In: Pickard, J. (ed), Antarctic Oasis. Sidney, Academic Press, 63-98.
Adamson, D. A., Mabin, M.C.G. & Luly, J.G. (1997). Holocene isostasy and late Cenozoic development of landforms including Beaver and Radok Lake basins in the Amery Oasis, Prince Charles Mountains, Antarctica. Antarctic Science, 9, 299-306.
Alonso, B. Anderson, J.B., Díaz, J.I. & Bartek, L.R. (1992). Pliocene-Pleistocene seismic stratigraphy of the Ross Sea: evidence for multiple ice sheet grounding episodes. In: Elliot, D.H. (ed.), Contributions to Antarctic Research III: Antarctic Research Series, 57. Washington, Geophysical Union, 93-103.
Anderson, J.B. (1999). Antarctic Marine Geology. Cambridge, Cambridge University Press, 289 pp.
Anderson, J.B. & Andrews, J.T. (1999). Radiocarbon constraints on ice sheet advance and retreat in the Weddell Sea, Antarctica. Geology, 27, 179-182.
Anderson, J.B. & Bartek, L.R. (1992). Cenozoic glacial history of the Ross Sea revealed by intermediate resolution seismic data combined with drill information. Antarctic Research Series, 56, 231-263.
Anderson, J.B. & Shipp, S.S. (2001). Evolution of the West Antarctic Ice Sheet. In: Alley, R. & Bindschadler, R. (eds.), The West Antarctic Ice Sheet, Behaviour and Environments, Antarctic Research Series, 77. Washington DC, American Geophysical Union, 45-57.
Anderson, J.B., Shipp, S.S., Bartek, L.R. & Reid, D.E. (1992). Evidence for a grounded ice sheet on the Ross Sea continental shelf during the late Pleistocene and preliminary paleodrainage reconstruction. In: Elliot, D.H. (ed), Contributions to Antarctic Research Series III, 57, Washington DC, American Geophysical Union, 39-42.
Anderson, J.B., Brake, C.F. & Myers, N.C. (1984). Sedimentation on the Ross Sea continental shelf, Antarctica. Marine Geology, 57, 295-333.
Anderson, J.B., Kurtz, D.D., Domack, E.W., & Balshaw, K.M. (1980). Glacial and glacial-marine sediments of the Antarctic continental shelves. Journal of Geology, 88, 399-414.
Anderson, J. B., Kennedy, D. S., Smith, M. J., and Domack, E. W. (1991a). Sedimentary facies associated with Antarctica’s floating ice masses. In Anderson, J. B. and Ashley, G. M. (eds.), Palaeoclimatic interpretation of glacial marine deposits. Geological Society of America Special Paper, 256, 1-25.
Anderson, J.B., Andrews, B.A., Bartek, L.R. & Truswell, E.M. (1991b). Petrology and palynology of glacial sediments: implications for subglacial geology of the eastern Weddell Sea, Antarctic. In: Thomson, M.R.A., Crame, J.A. & Thomson, J.W. (eds), Geological Evolution of Antarctic. Cambridge, Cambridge University Press, 231-235.
Anderson, J.B., Shipp, S.S. & Siringan, F.P. (1992). Preliminary seismic stratigraphy of the northwestern Weddell Sea continental shelf. In: Yoshida, Y., Kaminuma, K. & Shiraishi, K. (eds), Recent progress in Antarctic earth sciences. Tokyo, Terra Scientific Publishing Company, 603-612.
Anderson, J.B., Shipp, S.S., Lowe, A.L., Wellner, J.S. & Mosola, A.B. (2002). The Antarctic Ice Sheet during the Last Glacial Maximum and its subsequent retreat history: a review. Quaternary Science Reviews, 21, 49-70.
Andrews, J.T., Domack, E.W., Cunningham, W.L., Levanter, A., Licht, K., Jull, A.J.T., DeMaster, D.J. & Jennings, A.E. (1999). Problems and possible solutions concerning radiocarbon dating of surface marine sediments, Ross Sea, Antarctica. Quaternary Research, 52, 206-216.
Askin, R.A. (1992). Late Cretaceous-early Tertiary Antarctic outcrop evidence for past vegetation and climate. Antarctic Research Series, 56, 61-73.
Augustinus, P.C., Gore, D.B., Leisman, M.R., Zwartz, D. & Colhoun, E.A. (1997). Reconstruction of ice flow across the Bunger Hills, East Antarctica. Antarctic Science, 9, 347-354.
Banfield, L.A. & Anderson, J.B. (1995). Seismic facies investigation of the Late Quaternary glacial history of Bransfield Basin, AntarctiC. Antarctic Research Series, 68, 123-140.
Barcena, M. A., Gersonde, R., Ledesma, S., Fabrés, J., Calafat, A.M., Canals, M., Javier, F. & Flores, J.A. (1998). Record of Holocene glacial oscillations in Bransfield Basin as revealed by siliceous microfossil assemblages. Antarctic Science, 10, 269-285.
Baroni, C. (1994). Notes on late-glacial retreat of the Antarctic Ice Sheet and Holocene environmental changes along the Victoria Land coast. Memoirs of National Institute of Polar Research Special Issue, 50, 85-107.
Baroni, C. & Orombelli, G. (1991). Holocene raised beaches at Terra Nova Bay, Victoria Land, Antarctica. Quaternary Research, 36, 157-177.
Baroni, C. & Orombelli, G. (1994a). Abandoned penguin rookeries as Holocene palaeoclimatic indicators in Antarctica. Geology, 22, 23-26.
Baroni, C. & Orombelli (1994b). Holocene glacier variations in the Terra Nova Bay area (Victoria Land, Antarctica). Antarctic Science, 6, 497-505.
Barrera, E. & Huber, B.T. (1993). Eocene to Oligocene oceanography and temperatures in the Antarctic Indian Ocean. Antarctic Research Series, 56, 49-65.
Barrett, P.J. (1997). Antarctic palaeoenvironment through Cenozoic times: a review. Terra Antarctica, 3, 103.119.
Barret, P.J, Hambrey, M.J. & Robinson, P.R. (1991). Cenozoic glacial and tectonic history from CIROS-1, McMurdo Sound. In: Thomson, M.R.A., Crame, J.A. & Thomson, J.W. (eds), Geological Evolution of Antarctica. Cambridge: Cambridge University Press, 651-656.
Barrett, P.J., Adams, C.J., McIntosh, W.C., Swisher, C.C. & Wilson, G.S. (1992). Geochronological evidence supporting Antarctic deglaciation three million years ago. Nature, 359, 816-818.
Barron, J., Larsen, B. & Baldauf, J.G. (1991). Evidence for late Eocene to early Oligocene Antarctic glaciation and observations on late Neogene glacial history of Antarctica: results from Leg 119. In: Barron, J. & Larsen, B. (eds), Proceedings of the Ocean Drilling Program, Scientific Results, 119. Texas, College Station, 869-891.
Barsch, D. & Mäusbacher, R. (1986a). Beiträge zur Vergletscherungsgeschichte und zur Reliefentwicklung der Südshetland Insel. Zeitschrift für Geomorphologie 61, 25-37.
Barsch, D. & Mäusbacher, R. (1986b). New data on the relief development of the South Shetland Islands, Antarctic. Interdisciplinary Science Reviews, 11, 211-218.
Bart, P.J. & Anderson, J.B. (1995). Seismic record of glacial events affecting the Pacific margin of the nortwestern Antarctic Peninsula. Antarctic Research Series, 68, 75-79.
Bentley, M.J. (1999). Volume of Antarctic ice at the Last Glacial Maximum, and its impact on global sea level change. Quaternary Science Reviews, 18, 1569-1595.
Bentley, M.J. & Anderson, J.B. (1998). Glacial and marine geological evidence for the ice sheet configuration in the Weddell Sea-Antarctic Peninsula region during the Last Glacial Maximum. Antarctic Science, 10, 309-325.
Bentley, M.J., Kubik, P.W. & Sugden, D.E. (2000). Geomorphological evidence and cosmogenic 10Be/26 Al exposure ages for the Last Glacial Maximum configuration of the Antarctic Peninsula Ice Sheet. In: Hjort, C. (ed.), Glaciation of the Weddell Sea basin. ANTIME-Workshop, Abisko, Sweden, 16-19 September 2000. Lund, Department of Quaternary Geology, 5.
Berger, A. & Loutre, M.F. (1991). Insolation values for the climate of the last 10 million years. Quaternary Science Reviews, 10, 297-317.
Berkman, P. (1994). Geochemical signatures of meltwater in mollusc shells from Antarctic coastal areas during the Holocene. Memoirs of National Institute of Polar Research, Special Issue, 50, 11-27.
Berkman, P. (1997). Ecological variability in Antarctic coastal environments: past and present. In: Battaglia, B., Valencia, J. & Walton, D.W.H. (eds), Antarctic Communities, Structure and Survival. Cambridge: Cambridge University Press, 349-357.
Berkman, P.A. & Forman, S.L. (1996). Pre-bomb radiocarbon and the reservoir correction for calcareous marine species in the Southern Ocean. Geophysical Research Letters, 23, 363-366.
Berkman, P.A., Andrews, J.T., Björck, S., Colhoun, E.A., Emslie, S.D., Goodwin, I.D., Hall, B.L., Hart, C.P., Hirakawa, K., Igarashi, A., Ingólfsson, Ó., López-Martínez, J., Lyons, W.B., Mabin, M.C.G., Quilty, P.G., Taviani, M. & Yoshida, Y. (1998). Circum-Antarctic coastal environmental shifts during the Late Quaternary reflected by emerged marine deposits. Antarctic Science, 10, 345-362.
Bird, M.L., Chivas, A.R., Radnell, C.J. & Burton, H.R. (1991). Sedimentological and stable-isotope evolution of lakes in the Vestforld Hills, Antarctica. Palaeogeography, Palaeoclimatology, Palaeoecology, 84, 109-130.
Birkenmajer, K. (1981). Lichenometric dating of raised marine beaches at Admirality Bay, King George Island (South Shetland Islands, West Antarctica). Bulletin de l´Academie Polonaise des Sciences 29, 119-127.
Birkenmajer, K. (1987). Oligocene-Miocene glaciomarine sequences of King George Island (South Shetland Islands), Antarctica. Palaeontologia Polonica 49, 9-36.
Birkenmajer, K. (1988). Tertiary glacial and interglacial deposits, South Shetland Islands, Antarctica: geochronology versus biostratigraphy (a process report). Bulletin of the Polish Academy of Sciences, 36, 133-145.
Birkenmajer, K. (1991). Tertiary glaciation in the South Shetland Islands, West Antarctica: evaluation of data.In: Thomson, M.R.A., Crame, J.A. & Thomson, J.W. (eds), Geological Evolution of Antarctica. Cambridge: Cambridge University Press, 627-632.
Birkenmajer, K., Ochyra, R., Olsson, I.U. & Stuchlik, L. (1985). Mid-Holocene radiocarbon dated peat at Admirality Bay, King George Island (South Shetland Islands, West Antarctica. Bulletin of the Polish Academy of Sciences, 33, 7-13.
Björck, S., Hjort, C., Ingólfsson, Ó. & Skog, G. (1991a). Radiocarbon dates from the Antarctic Peninsula region - problems and potential. Quaternary Proceedings, 1, 55-65.
Björck, S., Malmer, N., Hjort, C., Sandgren, P., Ingólfsson, Ó., Wallen, B., Smith R.I.L. & Liedberg-Jönsson, B. (1991b). Stratigraphic and palaeoclimatic studies of a 5,500 years old moss bank on Elephant Island, Antarctica. Arctic and Alpine Research 23, 361-374.
Björck, S., Sandgren, P. & Zale, R. (1991c). Late Holocene Tephrochronology of the Northern Antarctic Peninsula. Quaternary Research, 36, 322-328.
Björck, S., Hĺkansson, H., Zale, R., Karlén, W.and Liedberg-Jönsson, B. (1991d). A late Holocene lake sediment sequence from Livingston Island, South Shetland Islands, with palaeoclimatic implications. Antarctic Science, 3, 61-72
Björck, S., Hĺkansson, H., Olsson, S., Barnekow, L. & Jansens, J. (1993). Paleoclimatic studies in South Shetland Islands, Antarctica, based on numerous stratigraphic variables inlake sediments. Journal of Paleolimnology, 8, 233-272.
Björck, S., Hĺkansson, H., Olsson, S., Ellis-Evans, C., Humlum, O. & Lirio, J.M. (1996a). Late Holocene palaeoclimatic records from lake sediments on James Ross Island, Antarctica. Palaeogeography, Palaeoclimatology, Palaeoecology, 113, 195-220.
Björck, S., Hjort, C. Ingólfsson, Ó., Zale, R. & Ising, J. (1996b). Holocene deglaciation chronology from lake sediments. In: López-Martínez, J., Thomson, M.R.A. & Thomson, J.W. (eds.), Geomorphological map of Byers Peninsula, Livingston Island. BAS GEOMAP Series, Sheet 5-A, 1:25 000. Cambridge: British Antarctic Survey. 65 pp.
Blunier, T., Chappellaz, J., Schwander, J., Dallenbach, A., Stauffer, B., Stocker, T.F. & Raynaud, D. (1998). Asynchrony of Antarctic and Arctic climate change during the last glacial period. Nature, 394, 793-743.
Bolshiyanov, D.Y., Verkulich, S.R., Klokov, V., Makeev, V. & Arslanov, H. (1990). Radiocarbon dating of sediments from the Bunger oasis (East Antarctica). In: Wand, U. & Strauch, G. (eds.), Proceedings of 5th Working Meeting Isotopes in Nature. Leipzig: Zentralinstitut für Isotopenforschung, 47-52.
Bolshiyanov, D., Verkulich, S., Pushina, Z. & Kirienko, E. (1991). Some features of the Late Pleistocene and Holocene history of the Bunger Hills (East Antarctica). Abstracts Sixt International Symposium on Antarctic Earth Science. Tokyo: National Institute of Polar Research, 66-71.
Brachfeld, S. A., Banerjee, S. K, Guydo, Y. & Acton, G. D. (2002). A 13 200 year history of century to millennial-scale paleoenvironmental change magnetically recorded in the Palmer Deep, western Antarctic Peninsula. Earth and Planetary Science Letters, 194: 311-326.
Broecker, W.S. (1963). Radiocarbon ages of Antarctic materials. Polar Record 11, 472-473.
Bronge, C. (1992). Holocene climatic record from lacustrine sediments in a freshwater lake in the Vestfold Hills, Antarctica. Geografiska Annaler, 74A, 47-58.
Burgess, J.S., Spate, A.P. & Shevlin, J. (1994). The onset of deglaciation in the Larsemann Hills, Eastern Antarctica. Antarctic Science, 6, 491-495.
Burgess, J., Carson, C., Head, J. & Spate, A. (1995). Larsemann Hills: Not heavily glaciated during the Last Glacial Maximum. In: Ricci, C.A. (ed.), The Antarctic Region: Geological Evolution and Processes. Siena: Universtŕ degli Studi di Siena, 841-843.
Cameron, R.L., Loken, O. & Molholm, J. (1959). Wilkes Station Glaciological Data 1957-58. Ohio State University Research Foundation Report, 825-1, Part 3.
Canals, M., Urgeles, R. & Calafat, A.M. (2000). Deep sea-floor evidence of past ice streams off the Antarctic Peninsula. Geology, 28, 31-34.
Carrara, P. (1979). Former extent of glacial ice in Orville Coast region, Antarctic Peninsula. Antarctic Journal of the United States, 14, 45-46.
Carrara, P. (1981). Evidence for a former large ice sheet in the Orville Coast-Ronne Ice Shelf area, AntarctiC. Journal of Glaciology, 27, 487-491.
Ciais, P., Jouzel, J., Petit, J.R., Lipenkov, V. & White, J.W.C. (1994). Holocene temperature variations inferred from six Antarctic ice cores. Annals of Glaciology, 20, 427-436.
Clapperton, C.M. (1990). Quaternary glaciations in the Southern Ocean and Antarctic Peninsula area. Quaternary Science Reviews, 9, 229-252.
Clapperton, C.M. & Sugden, D.E. (1982). Late Quaternary glacial history of George VI Sound area, West Antarctica. Quaternary Research, 18, 243-267.
Clapperton, C.M. & Sugden, D.E. (1988). Holocene glacier fluctuations in South America and Antarctica. Quaternary Science Reviews, 7, 185-198.
Clapperton, C.M. & Sugden, D.E. (1990). Late Cenozoic glacial history of the Ross embayment, Antarctica. Quaternary Science Reviews, 9, 253-272.
Clark, J.A. & Lingle, C.S. (1979). Predicted sea-level changes (18,000 years BP to present). caused by late-glacial retreat of the Antarctic Ice Sheet. Quaternary Research, 9, 265-287.
Colhoun, E.A. (1997). A review of geomorphological research in Bunger Hills and expansion of the East Antarctic ice sheet during the Last Glacial Maximum. In: Ricci, C.A. (ed.), The Antarctic Region: Geological Evolution and Processes. Siena: Universtŕ degli Studi di Siena, 801-807.
Colhoun, E.A. & Adamson, D.A. (1991). Raised beaches of the Bunger Hills. In: Gilleson, D. & Fitzsimons, S. (eds.), Quaternary Research in Australian Antarctica: Future directions. Special Publication No.3. Canberra: Department of Geography and Oceanography, University College, Australian Defence Force Academy, 79-84.
Colhoun, E.A. & Adamson, D.A. (1992a). The Quaternary history of the Bunger Hills, East Antarctica. In: Yoshida, Y., Kaminuma, K. & Shiraishi, K. (eds.), Recent Progress in Antarctic Earth Science. Tokyo: Terra Scientific Publishing Company, 745-750.
Colhoun, E.A. & Adamson, D.A. (1992b). Raised beaches of the Bunger Hills. Australian National Antarctic Research Expedition Report, 136, 47 pp.
Colhoun, E.A., Mabin, M.C.G., Adamsson, D.A. & Kirk, R.M. (1992). Antarctic ice volume and contribution to sea-level fall at 20,000 yr BP from raised beaches. Nature 358, 316-319.
Conway, H., Hall, B.L., Denton, G.H., Gades, A.M. & Waddington, E.D. (1999). Past and future grounding-line retreat of the West Antarctic Ice Sheet. Science, 286, 280-283.
Cooper, A.P.R. (1997). Historical observations of Prins Gustav Ice Shelf. Polar Record, 33, 285-294.
Cunningham, W.L., Levanter, A., Andrews, J.T., Jennings, A.E. & Licht, K.J. (1999). Late Pleistocene-Holocene marine conditions in the Ross Sea, Antarctica: evidence from the diatom record. The Holocene, 9, 129-139.
Curl, J.E. (1980). A glacial history of the South Shetland Islands, AntarctiC. Institute of Polar Studies Report 63. Columbus: Institute of Polar Studies, Ohio State University, 129 pp.
Denton, G.H. (1979). Glacial history of the Byrd-Darwin Glacier area, Transantarctic Mountains. Antarctic Journal of the United States, 14, 57-58.
Denton, G.H. & Hughes, T.J (2000). Reconstruction of the Ross ice drainage system, Antarctica, at the last glacial maximum. Geografiska Annaler, 82A, 143-166.
Denton, G.H. & Marchant, D.R. (2000). The geologic basis for a reconstruction of a grounded ice sheet in McMurdo Sound, Antarctica, at the last glacial maximum. Geografiska Annaler, 82A, 167-212.
Denton, G.H., Armstrong, R.L. & Stuiver, M. (1971). The Late Cenozoic Glacial History of Antarctica. The Late Cenozoic Glacial Ages. New Haven, Yale University Press, 267-306.
Denton, G.H., Andersen, B.G. & Conway, H.W. (1986). Late Quaternary surface fluctuations of Beardmore Glacier, Antarctica. Antarctic Journal of the United States, 21, 90-92.
Denton, G.H., Bockheim, J.G., Wilson, S.C. & Stuiver, M. (1989). Late Wisconsin and Early Holocene glacial history, inner Ross Embayment, Antarctica. Quaternary Research, 31, 151-182.
Denton, G.H., Prentice, M.L. & Burckle (1991). Cainozoic history of the Antarctic ice sheet. In: Tingey, R.J. (ed.), Geology of Antarctica. Oxford: Oxford University Press, 365-433.
Denton, G.H., Sugden, D.E., Marchant, D.R., Hall, B.L. & Wilch, T.I. (1993). East Antarctica Ice Sheet sensitivity to Pliocene climate change from a Dry Valleys perspective. Geografika Annaler, 75A, 155-204.
Domack, E.W. (1992). Modern carbon-14 ages and reservoir corrections for the Antarctic Peninsula and Gerlache Strait area. Antarctic Journal of the United States, 27, 63-64.
Domack, E.W., Jull, A.J.T., Anderson, J.B., Linick, T.W. & Williams, C.R. (1989). Application of Tandem Accelerator Mass-spectrometre dating to Late Pleistocene-Holocene sediments of the East Antarctic Continental Shelf. Quaternary Research, 31, 277-287.
Domack, E.W., Jull, A.J.T. & Donahue, D.J. (1991). Holocene chronology for the unconsolidated sediments at Hole 740A: Prydz Bay, East Antarctica. Proceedings of the Ocean Drilling Program Scientific Results, 119B (207), 1-7.
Domack, E.W., Ishman, S.E., Stein, A.B., McClennen, C.E. & Jull, A.J.T. (1995). Late Holocene advance of the Müller Ice Shelf, Antarctic Peninsula: sedimentological, geochemical and palaeontological evidence. Antarctic Science, 7, 159-170.
Domack, E.W., O´Brien, P. Harris, P., Taylor, F., Quilty, P.G., DeSantis, L. & Raker, B. (1998). Late Quaternary sediment facies in Prydz Bay, East Antarctica, and their relationship to glacial advances onto the continental shelf. Antarctic Science, 10, 236-246.
Domack, E,W., Jacobson, E.A., Shipp, S. & Anderson, J.B. (1999a). Sedimentologic and stratigraphic signature of the late Pleistocene/Holocene fluctuations of the West Antarctic Ice Sheet in the Ross Sea: a new perspective Part 2. Geological Society of America Bulletin, 111, 1486-1516.
Domack, E.W., O’Brien, P., Harris, P., Taylor, F., Quilty, P.G., DeSantis, L. & Raker, B. (1999b). Late Quaternary sediment facies in Prydz Bay, East Antarctica, and their relationship to glacial advance onto the continental shelf. Antarctic Science, 10, 236-246.
Domack, E., Hall, B.L. & Hayes, J.M. (2000). Accurate Antarctic dating technique sought by Quaternary community. EOS, 80(49), 591 and 596.
Domack, E., Levanter, A., Dunbar, R., Taylor, F., Brachfield, S., Sjunneskog, C. & ODP Leg 178 Scientific Party (2001a). Chronology of the Palmer Deep site, Antarctic Peninsula: a Holocene palaeoenvironmental reference for the circum-Antarctic. The Holocene, 11, 1-9.
Domack, E., Levanter, A., Gilbert, R., Brachfeld, S., Ishman, S., Camerlenghi, A., Gavahan, K., Carlson, D. & Barkoukis, A. (2001b). Cruise reveals history of Holocene Larsen Ice Shelf. EOS, 82(2), 13 and 16-17.
Doran, P.T., Wharton Jr., R.A. & Lyons, W.B. (1994). Paleolimnology of the McMurdo Dry Valleys, Antarctica. Journal of Paleolimnology, 10, 85-114.
Doran, P.T., Berger, G.W., Lyons, W.B., Wharton Jr., R.A., Davisson, M.L., Southon, J. & Dibb, J.E. (1999). Dating Quaternary lacustrine sediments in the McMurdo Dry Valleys, Antarctica. Palaeogeography, Palaeoclimatology, Palaeoecology, 147, 223-239.
Doran, P. T., Priscu, J. C., Lyons, W. B., Walsh, J. E., Fountain, A. G., McKnight, D. M., Moorhead, D. L., Virginia, R. A., Wall, D. H., Clow, G. D., Fritsen, C. H., McKay, C. P. & Parsons, A. N. (2002). Antarctic climate cooling and terrestrial ecosystem response. Nature 517.
Dowsett, H.J. Barron, J.A. Poore, R.Z. (1996). Middle Pliocene sea-surface temperatures: a global reconstruction. Marine Micropalaeontology, 27, 13-25.
Drewry, D.J. (1983). Antarctica: Glaciological and Geophysical Folio. Cambridge, Scott Polar Research Institute, 9 sheets.
Drewry, D.J. (1979). Late Wisconsin reconstruction for the Ross Sea region, Antarctica. Journal of Glaciology, 24, 231-244.
Drewry, D.J., Jordan, S.R. & Jankowski, E. (1982). Measured properties of the Antarctic ice sheet: Surface configurations, ice thickness, volume and bedrock characteristics. Annals of Glaciology, 3, 83-91.
Elverhři, A. (1981). Evidence for a late Wisconsin glaciation of the Weddell Sea. Nature, 293, 641-642.
Elverhři, A. & Maisey, G. 1983. Glacial erosion and morphology of the eastern and southeastern Weddell Sea shelf. In: Oliver, R.L., James, P.R. & Jago, , J.B. (eds.), Antarctic Earth Science. Canberra, Australian Academy of Sciences, 483-487.
Emiliani, C. (1969). Interglacial high sea levels and the control of Greenland ice by the precession of the equinoxes. Science, 166, 1503-1504.
Emslie, S.D. (1995). Age and taphonomy of abandoned penguin rookeries in the Antarctic Peninsula. Polar Record, 31, 409-418.
Emslie, S.D. (2001). Radiocarbon dates from abandoned penguin colonies in the Antarctic Peninsula region. Antarctic Science, 13, 289-295.
Emslie, S.D., Fraser, W., Smith, R.C. & Walker, W. (1998). Abandoned penguin colonies and environmental change in the Palmer Station area, Anvers Island, Antarctic Peninsula. Antarctic Science, 10, 257-268.
Fabres, J., Calafat, A., Canals, M., Barcena, M.A. & Flores, J.A. (2000). Bransfield Basin fine-grained sediments: late-Holocene sedimentary processes and Antarctic oceanographic conditions. The Holocene, 10, 703-718.
Fairbanks, R.G., (1989). A 17,000-year glacio-eustatic sea level record. Nature, 342, 637-642.
Fenton, J.H.C. (1980). The rate of peat accumulation in Antarctic moss banks. Journal of Ecology, 68, 211-228.
Fenton, J.H.C. (1982). The formation of vertical edges on Antarctic moss peat banks. Arctic and Alpine Research, 14, 21-26.
Fitzsimons, S.J. (1997). Depositional models for moraine formation in East Antarctic coastal oases. Journal of Glaciology, 43, 256-264.
Fitzsimons, S.J. & Colhoun, E.A. (1995). Form, structure and stability of the margin of the Antarctic ice sheet, Vestfold Hills and Bunger Hills, East Antarctica. Antarctic Science, 2, 171-179.
Fitzsimons, S.J. & Domack, E.W. (1993). Evidence for early Holocene deglaciation of the Vestfold Hills, Antarctica. Polar Record, 29, 237-240.
Flemming, R.F. & Barron, J.A. (1996). Evidence of Pliocene Nothofagus in Antarctica from Pliocene marine sedimentary deposits (DSDP Site 274). Marine Micropalaeontology, 27, 227-236.
Flemming, K., Johnston, P., Zwartz, D., Yokoyama, Y., Lambeck, K. & Chappell, J. (1998). Refining the eustatic sea-level curve since the Last Glacial Maximum using far- and intermediate-field sites. Earth and Planetary Science Letters, 163, 327-342.
Francis, J.E. (1995). 381 growth forms of fossil trees and climatic implications for the Pliocene Sirius Group, Dominion Range, Transantarctic Mountains. VII International Symposium on Antarctic Earth Sciences Abstracts. Siena, Typografia Senese, 142
Fraser, W.R., Trivelpiece, W.Z., Ainley, D.G. & Trivelpiece S.G. (1992). Increases in Antarctic penguin populations: Reduced competition with whales or a loss of sea ice due to environmental warming. Polar Biology, 11, 525-531.
Fulford-Smith, S.P. & Sikes, E.L. (1996). The evolution of Ace Lake, Antarctica, determined from sedimentary diatom assemblages. Palaeogeography, Palaeoclimatology, Palaeoecology, 124, 73-86.
Fütterer, D.K. & Melles, M. (1990). Sediment patterns in southern Weddell Sea: Filchner shelf and Filchner depression. In: Bleil, U. & Thiede, J. (Eds), Geological History of the Polar Oceand: Arctic versus Antarctic. Boston, Kluwer Academic Publishers, 381-401
Gersonde, R. & Zielinski, U. (2000). The reconstruction of late Quaternary Antarctic sea-ice distribution – the use of diatoms as a proxy for sea-ice. Palaeogeography, Palaeoclimatology, Palaeoecology, 162, 263-286.
Gillieson, D.S. (1991). An environmental history of two freshwater lakes in the Larsemann Hills, Antarctica. Hydrobiologia, 214, 327-331.
Gingele, F., Kuhn, G., Maus, B., Melles, M. & Schöne, T. (1997). Holocene ice retreat from the Lazarev Sea shelf, East Antarctica. Continental Shelf Research, 17(2), 137-163.
Goodwin, I.D. (1993). Holocene deglaciation, sea-level change, and the emergence of the Windmill Islands, Budd Coast, Antarctica. Quaternary Research, 40 70-80.
Goodwin, I. D. (1996). A mid to late Holocene readvance of the Law Dome ice margin, Budd Coast, East Antarctica. Antarctic Science, 8, 395-406.
Goodwin, I.D. (1998). Did changes in Antarctic ice volumes influence late Holocene sea-level lowering. Quaternary Science Reviews, 17, 319-332.
Goodwin, I.D. & Zweck, C. (2000). Glacio-isostasy and glacial ice load at Law Dome, Wilkes Land, East Antarctica. Quaternary Research, 53, 285-293.
Gordon, J.E. & Harkness, D.D. (1992). Magnitude and geographic variation of the radiocarbon content in Antarctic marine life: implications for reservoir corrections in radiocarbon dating. Quaternary Science Reviews, 11, 697-708.
Gore, D.B. (1997). Blanketing snow and ice; constraints on radiocarbon dating deglaciation in East Antarctic oases. Antarctic Science, 9, 336-346.
Grobe, H. & Mackensen, A (1992). Late Quaternary climate cycles as recorded in sediments from the Antarctic continental margin. Antarctic Research Series, 56, 349-376.
Grobe, H., Huybrechts, H. & Futterer, D.K. (1993). Late Quaternary record of sea-level changes in the Antarctic. Geologische Rundshau, 82, 263-275.
Hĺkansson, H., Olsson, S. & Björck, S. (1995). Diatom records and sediment chemistry of a shallow glaciolacustrine basin on James Ross Island, AntarctiC. PACT, 50, 417-430.
Hall, B.L. (1997). Geological Assessment of Antarctic Ice Sheet Stability. Ph.D. thesis, University of Maine, Orono, 324 pp. (unpublished).
Hall, B.L. & Denton, G.H. (1999). New relative sea-level curves for the southern Scott Coast, Antarctica: evidence for Holocene deglaciation of the Ross Sea. Journal of Quaternary Science, 14, 641-650.
Hall, B.L. & Denton G.H. (2000). Radiocarbon chronology of Ross Sea Drift, Eastern Taylor Valley, Antarctica: Evidence for a grounded ice sheet in the Ross Sea at the Last Glacial Maximum. Geografiska Annaler, 82, 305-336.
Hambrey, M.J. (1993). Cenozoic sedimentary and climatic record, Ross Sea Region, Antarctica. Antarctic Research Series, 60, 91-124.
Hambrey, M.J., Larsen, B. & Ehrmann, W.U. (1989). Forty million years of Antarctic glacial history yielded by Leg 119 of the Ocean Drilling Program. Polar Record 25, 99-106.
Hambrey, M.J., Ehrmann, W.U. & Larsen, B. (1991). Cenozoic glacial record of the Prydz Bay continental shelf, East Antarctica. Proceedings of the Ocean Drilling Programme, Scientific Results, 119, 77-132.
Hansom, J.D. & Flint, C.P. (1989). Short notes. Holocene ice fluctuations on Brabant Island, Antarctic Peninsula. Antarctic Science, 1, 165-166.
Harden, S.L., DeMaster, D.J. & Nittrouer, C.A. (1992). Developing sediment geochronologies for high-latitude continental shelf deposits: a radiochemical approach. Marine Geology, 103, 69-97.
Harris, P.T. & O’Brien, P.E. (1996). Geomorphology and sedimentology of the continental shelf adjacent to to Mac Robertson Land, East Antarctica. Geo-Marine Letters, 16, 287-296.
Harris, P.T. & O´Brien, P.E. (1998). Bottom currents, sedimentation and ice-sheet retreat facies successions on the Mac Robertson shelf, East Antarctica. Marine Geology, 151, 47-72.
Harris, P.T., Taylor, F., Domack, E., DeSantis, L., Goodwin, I., Quilty, P.G. & O’Brien, P.E. (1997). Glacimarine siliclastic muds from Vincennes Bay, East Antarctica; preliminary results of an Exploratory Cruise in 1997. Terra Antarctica, 4, 11-20.
Hayashi, M. & Yoshida, Y. (1994). Holocene raised beaches in the Lützow-Holm Bay region, East Antarctica. Memoirs of National Institute of Polar Research Special Issue, 50, 49-84.
Heine, J.C. & Speir, T.W. (1989). Ornithogenic soils of the Cape Bird Adélie penguin rookeries, Antarctica. Polar Biology, 10, 89-99.
Hiemstra, J. F., (2001). Microscopic analyses of Quaternary glacigenic sediments of Marguerite Bay, Antarctic Peninsula. Arctic, Antarctic and Alpine Research, 33, 258-265.
Hill, R.S & Trustwell, E.M. (1993). Nothofagus fossils in the Sirius Group, Transantarctic Mountains; leaves and pollen and their climatic implications. Antarctic Research Series, 60, 67-73.
Hiller, A., Wand, U., Kämpf, H. & Stackebrandt, W. (1988). Occupation of the Antarctic continent by petrels during the past 35000 years: inferences from a 14C study of stomach oil deposits. Polar Biology, 9, 69-77.
Hirakawa, K., Ono, Y., Hayashi, M., Aniya, M., Iwata, S., Fujiwara, K., Moriwaki, K. & Yoshida, Y. (1984). Antarctic Geomorphological Map of Langhovde. Tokyo, National Institute of Polar Research, 63 pp.
Hjort, C. & Ingólfsson, Ó. (1990). Studies of the Glacial History in the Antarctic Peninsula Area. In: Karlqvist, A. (ed), Swedish Antarctic Research Programme 1988/89 – A Cruise Report. Stockholm, Swedish Polar Research Secretary, 76-80.
Hjort, C., Ingólfsson, Ó., Möller, P. & Lirio, J.M. (1997). Holocene glacial history and sea-level changes on James Ross Island, Antarctic Peninsula. Journal of Quaternary Science, 12, 259-273.
Hjort, C., Björck, S., Ingólfsson, Ó. & Möller, P. (1998). Holocene deglaciation and climate history of the northern Antarctic Peninsula region - a discussion of correlations between the Southern and Northern Hemispheres. Annals of Glaciology, 27, 110-112.
Hjort, C., Ingólfsson, Ó. Bentley, M. J., and Björck, S. in press: The late Pleistocene and Holocene glacial and climate history of the Antarctic Peninsula region as documented by land and lake sediment records – a review. In Domack E. W., Leventer, A., Burnett, A., Kirby, and Bindschadler, R. (eds.), Antarctic Peninsula Climate Variability: a Historical and Paleoenvironmental Perspective. American Geophysical Union, Antarctic Research Series XX: 000-000.
Hjort, C., Bentley, M.J. & Ingólfsson, Ó. (2001). Holocene and pre-Holocene temporary disappearance of the George VI Ice Shelf, Antarctic Peninsula. Antarctic Science, 13, 296-301.
Hollin, J.T. (1962). On the glacial history of Antarctica. Journal of Glaciology, 4, 173-195.
Hughes, T.J. (1975). The West Antarctic Ice Sheet: Instability, disintegration and the initiation of ice ages. Reviews of Geophysics and Space Physics, 15, 1-46.
Hughes, T.J. (1987). Deluge II and the continent of doom: rising sea level and collapsing Antarctic ice. Boreas, 16, 89-100.
Hughes, T.J., Denton, G.H., Andersen, B.G., Schilling, D.H., Fastook, J.L. & Lingle, C.S. (1981). The last great ice sheets: A global view. In: Denton, G.H. & Hughes, T.J. (eds.), The Last Great Ice Sheets. New York: John Wiley & Sons, 263-317.
Igarashi, A., Harada, N. & Moriwaki, K. (1995). Marine fossils of 30-40 m raised beach deposits and late Pleistocene glacial history around Lutzow-Holm Bay, East Antarctica. Proceedings NIPR Symposium on Antarctic Geosciences, 8, 219-229.
Ingólfsson, Ó., Hjort, C., Björck, S. & Smith, R.I.L. (1992). Late Pleistocene and Holocene glacial history of James Ross Island, Antarctic Peninsula. Boreas, 21, 209-222.
Ingólfsson, Ó., Hjort, C., Berkman, P., Björck, S., Colhoun, E., Goodwin, I.D., Hall, B., Hirakawa, K., Melles, M., Möller, P. & Prentice, M. (1998). Antarctic glacial history since the Last Glacial Maximum: an overview of the record on land. Antarctic Science, 10, 326-344.
Ingólfsson, Ó. & Hjort, C. (1999). The Antarctic contribution to Holocene global sea-level rise. Polar Research, 18, 323-330.
Ingólfsson, Ó. & Hjort, C. (in press). Glacial and climate history of the Antarctic Peninsula since the Last Glacial Maximum. Arctic, Antarctic and Alpine Research, 34, 000-000.
John, B.S. (1972). Evidence from the South Shetland Islands toward a glacial history of West Antarctica. Institute of British Geographers, Special Publication, 4, 75-89.
John, B.S. & Sugden, D.E. (1971). Raised marine features and phases of glaciation in the South Shetland Islands. British Antarctic Survey Bulletin, 24, 45-111.
Johnson, C.G. & Vestal, J.R. (1991). Photosythetic carbon incorporation and turnover in Antarctic cryptoendolithic microbial communities: are they the slowest-growing communities on Earth? Applied and Environmental Microbiology, 57, 2308-2311.
Jones, V.J., Hodgson, D.A. & Chepstow-Lusty, A. (2000). Palaeolimnological evidence for marked Holocene environmental changes on Signey Island, Antarctica. The Holocene, 10, 43-60.
Jonsson, S. (1988). Observations on physical geography and glacial history of the Vestfjella nunataks in western Dronning Maud Land, Antarctica. Naturgeografiska Institutionen, Stockholms Universitet, Rapport, 68. 57 pp.
Jouzel, J., Lorius, C., Petir, J.R., Genthon, C., Barkov, N.I., Kotlyakov, V.M. & Petrov, V.M. (1987). Vostok ice core: a continuous isotope temperature record over the last climatic cycle (160,000 years). Nature, 329, 403-408.
Kellogg, T.B., Kellogg D.E. & Stuiver, M. (1990). Late Quaternary history of the southwestern Ross Sea: evidence from debris bands on the McMurdo Ice Shelf, Antarctica. In: Bentley, C.R. (ed.), Contributions to Antarctic Research I. Washington D.C.: American Geophysical Union, 25-56.
Kellogg, T.B., Hughes, T. & Kellogg, D.E. (1996). Late Pleistocene interactions of East and West Antarctic ice-flow regimes: evidence from the McMurdo Ice Shelf. Journal of Glaciology, 42, 486-500.
Kennedy, D.S. & Anderson, J.B. (1989). Glacial-marine sedimentation and Quaternary glacial history of Marguerite Bay, Antarctic Peninsula. Quaternary Research, 31, 255-276.
Kennett, J.P. (1977). Cenozoic evolution of Antarctic glaciation, the circum-Antarctic ocean, and their impact on global palaeoceanography. Journal of Geophysical Research, 82, 3843-3860.
Kennett, J.P. & Barker, P.F. (1990). Latest Cretaceous to Cenozoic climate and oceanographic developments in the Weddell Sea, Antarctica: an ocean drilling perspective. Proceedings of the Ocean Drilling Programme, Scientific Results, 113, 937-960.
Kennett, J.P. & Hodell, D.A. (1993). Evidence for relative climatic stability of Antarctica during the early Pliocene: a marine perspective. Geografiaka Annaler, 75A, 205-220.
Kennett, J.P. & Stott, L.D. (1990). Proteus and Proto-oceanus: ancestral Paleogene oceans as revealed from Antarctic stable isotopic results. Proceedings of the Ocean Drilling Program, Scientific Results, 113, 865-880.
King, J.C. (1994). Recent climate variability in the vicinity of the Antarctic Peninsula. International Journal of Climatology, 14, 357-369.
Kirby, M. E., Domack, E. W. & McClennen, C. E. (1998): Magnetic stratigraphy and sedimentology of Holocene glacial marine deposits in the Palmer Deep, Bellingshausen Sea, Antarctica: implications for climatic change. Marine Geology, 152, 247-259.
Kulbe, T., Melles, M., Verkulich, S.R. & Pushina, Z.V. (2001). East Antarctic climate and environmental variability over the last 9400 years inferred from marine sediments in the Bunger Oasis. Arctic, Antarctic, and Alpine Research, 33, 223-230.
Larter, R.D. & Barker, P.F. (1989). Seismic stratigraphy of the Antarctic Peninsula margin: a record of Pliocene-Pleistocene ice volume and palaeoclimate. Geology, 17, 731-734.
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. & Vanneste, L.E. (1995). Relict subglacial deltas on the Antarctic Peninsula outer shelf. Geology, 23, 33-36.
Levanter, A., Domack, E. W., Ishman, S. E., Brachfeld, S., McClennen, C. E. & Manley, P. (1996). Productivity cycles of 200-300 years in the Antarctic Peninsula region: understanding linkages among the sun, atmosphere, oceans, sea ice, and biota. Geological Society of America Bulletin, 108, 1642-1644.
Licht, K.J., Jennings, A.E., Andrews, J.T. & Williams, K.M. (1996). Chronology of the late Wisconsin ice retreat from the western Ross Sea, Antarctica. Geology, 24, 223-226.
Licht, K.J., Dunbar, N.W., Andrews, J.T. & Jennings, A.E. (1999). Distinguishing subglacial till and glacial marine diamictons in the western Ross Sea, Antarctica: implications for last glacial maximum grounding line. Geological Society of America Bulletin, 111, 91-103.
Lintinen, P. & Nenonen, J. (1997). Glacial history of the Vestfjella and Heimefrontfjella nunatak ranges in western Dronning Maud Land, Antarctica. In: Ricci, C.A. (ed.), The Antarctic Region: Geological Evolution and Processes. Siena: Universtŕ degli Studi di Siena, 845-852.
López-Martínez, J., Thomson, M.R.A., Arche, A., Björck, S., Ellis-Evans, J.C., Hathway, B., Hernández-Cifuentes, F., Hjort, C., Ingólfsson, Ó., Ising, J., Lomas, S., Martínez de Pisón, E., Serrano, E., Zale, R. & King, S. (1996). Geomorphological map of Byers Peninsula, Livingston Island. BAS GEOMAP Series, Sheet 5-A, 1:25 000, with supplementary text. Cambridge, British Antarctic Survey, 65 pp.
Lorius, C., Jouzel, J., Ritz, C., Merlivat, L., Narkov, N.I., Korotkevich, Y.S. & Kotlyakov, V.M. (1985). A 150,000-year climatic record from Antarctic ice. Nature 316, 591-596.
Lorius, C., Jouzel, J. & Raynaud, D. (1993). Glacials-interglacials in Vostok: climate and greenhouse gases. Global Planetary Change, 7, 131-143.
Lyons, W.B., Tyler, S.W., Wharton, R.A., McKnight, D.M. & Vaughn, B.H. (1998). Late Holocene desiccation of Lake Hoare and Lake Fryxell, McMurdo Dry Valleys, Antarctica. Antarctic Science, 10, 247-256.
Mabin, M.C.G. (1985). 14C ages for ”heroic era” penguin and seal bones from Inexpressible Island, Terra Nova Bay, North Victoria Land. New Zealand Antarctic Record, 6, 24-25.
Mabin, M.C.G. (1986). The Ross Sea section of the Antarctic ice sheet at 18,000 BP: evidence from Holocene sea-level changes along the Victoria Land coast. South African Journal of Science, 82, 506-508.
Mackensen, A., Grobe, H., Kuhn, G. & Fütterer, D.K. (1989). Stable isotope stratigraphy from the Antarctic continental margin during the last one million years. Marine Geology, 87, 315-321.
Maemoku, H., Miura, H., Saigusa, S. & Moriwaki, K. (1997). Stratigraphy of Late Quaternary raised beach deposits in the northern part of Langhovde, Lutzow-Holm Bay, East Antarctica. Proceedings NIPR Symposium on Antarctic geosciences, 10, 178-186.
Martinez-Macchiavello, J.C., Tatur, A., Servant-Vildary, S. & Del Valle, R. (1996). Holocene environmental change in a marine-estuarine-lacustrine sediment sequence, King George Island, South Shetland Islands. Antarctic Science, 8, 313-322.
Masson, V., Vimeux, F., Jouzel, J., Morgan, V., Delmotte, M., Ciais, P., Hammer, C., Johnsen, S., Lipenkov, V.Y., Mosley-Thompson, E., Petit, J.R., Steig, E.J., Stievenard, M. & Vaikmae, R. (2000). Holocene climate variability in Antarctica based on 11 ice-core isotopic records. Quaternary Research, 54, 348-358.
Mäusbacher, R. (1991). Die Jungkvartäre Relief- und Klimageschichte im Bereich der Fildeshalbinsel, Süd-Shetland-Inseln, Antarktis. Heidelberger Geographische Arbeiten, 89, 207 pp.
Mäusbacher, R., Müller, J. & Schmidt, R. (1989). Evolution of postglacial sedimentation in Antarctic lakes. Zeitschrift für Geomorphologie, 33, 219-234.
Mayewski, P.A. (1975). Glacial geology and late Cenozoic history of the Transantarctic Mountains, Antarctica. Institute of Polar Studies Report, 56, 1-168.
McMinn, A. (2000). Late Holocene increase in sea ice extent in fjords of the Vestfold Hills, eastern Antarctica. Antarctic Science, 12, 80-88.
McMinn, A., Hejnis, H., Harle, K. & McOrist, G. (2001). Late-Holocene climate change recorded in sediment cores from Ellis Fjord, eastern Antarctica. The Holocene, 11, 291-300.
McKelvey, B.C. (1991). The Cainozoic glacial record in south Victoria Land: a geological evaluation of the McMurdo Sound drilling projects. In: Tingey, R.J. (ed.), The Geology of Antarctica. Oxford, Clarendon Press, 434-454.
Melles, M. (1991). Paläoglaziologie und Paläozeanographie im Spätquartär am Kontinentalrand des südlichen Weddellmeres, Antarktis. Berichte zur Polarforshung, 81, 1-190.
Melles, M., Verkulich, S.R. & Hermichen, W-D. (1994). Radiocarbon dating of lacustrine and marine sediments from the Bunger Hills, East Antarctica. Antarctic Science, 6, 375-378.
Melles, M., Kulbe, T., Verkulich, S.R., Pushina, Z.V. & Hubberten, H.-W. (1997). Late Pleistocene and Holocene environmental history of Bunger Hills, East Antarctica, as revealed by fresh-water and epishelf lake sediments. In: Ricci, C.A. (ed.), The Antarctic Region: Geological Evolution and Processes. Siena: Universtŕ degli Studi di Siena, 809-820.
Mercer, J.H. (1968). Antarctic ice and Sangamon sea level. Commission of Snow and Ice Reports and Discussions, Publication no. 79 De L´Association Internationale D´Hydrologie Scientifique, 217-225.
Mercer; J.H. (1972). Some observations on the glacial geology of the Beardmore Glacier area. In: Adie, R.D. (ed.), Antarctic Geology and Geophysics. Oslo, Universitetsforlaget, 427-433.
Mercer, J.H. (1978). Glacial development and temperature trends in the Antarctic and in South America. In: van Zinderen, E.M. (ed.), Antarctic Glacial History and World Paleoenvironments. Rotterdam, A.A. Balkema, 73-79.
Mercer, J.H. (1981). Terrestrial deposit of the Ross Ice Shelf area, Antarctica. In: Hambrey, M.J. & Harland, W.B. (eds.), Earths Pre-Pleistocene Glacial Record. Cambridge, Cambridge University Press, 204-207.
Miller, K.G., Wright, J.D. & Fairbanks, R.G. (1991). Unlocking the ice house: Oligocene-Miocene oxygen isotopes, eustasy, and margin erosion. Journal of Geophysical Research, 96, 6829-6848.
Moriwaki, K., Yoshida, Y. & Harwood, D.M. (1992). Cenozoic glacial history of Antarctica - a correlative synthesis. In: Yoshida, Y., Kaminuma, K. & Shiraishi, K. (eds.), Recent Progress in Antarctic Earth Science. Tokyo, Terra Scientific Publishing Company, 773-780.
Möller, P. (1995). Subrecent moraine ridge formation on Cuff Cape, Victoria Land, Antarctica. Geografiska Annaler, 77A, 83-94.
Nichols, R.L. (1968). Coastal geomorphology, McMurdo Sound, Antarctica. Journal of Glaciology, 7, 449-478.
Nienow, J.A. & Friedmann, E.I. (1993). Terrestrial lithophytic (rock). communities. In: Friedmann, E.I. (ed.), Antarctic Microbiology. New York, John Wiley & Sons, 343-412.
O´Brien, P.E. & Harris P.T. (1996). Patterns of glacial erosion and deposition in Prydz Bay and the past behaviour of the Lambert Glacier. Papers and Proceedings of Royal Society of Tasmania, 130, 79-85.
O´Brien, P.E., DeSantis, L., Harris, P.T., Domack, E. & Quilty, P. (1999). Ice shelf grounding zone features of western Prydz Bay, Antarctica: sedimentary processes from seismic and sidescan images. Antarctic Science, 11, 78-91.
Omoto, K. (1977). Geomorphic development of the Sôya Coast, East Antarctica. Science Reports of Tohoku University, 7th Series, 27, 95-148.
Omoto, K. (1983). The problem and significance of radiocarbon geochronology in Antarctica. In: Oliver, R.L, James, P.R. & Jago, J.B. (eds.), Antarctic Earth Science. Cambridge, Cambridge University Press, 205-209.
Orombelli, G., Baroni, C. & Denton, G.H. (1991). Late Cenozoic glacial history of the Terra Nova Bay region, Northern Victoria Land, Antarctica. Geografia Fisica Dinamica Quaternaria, 13, 139-163.
Payne, A.J.D., Sugden, D.E. & Clapperton, C.M. (1989). Modelling the growth and decay of the Antarctic Peninsula ice sheet. Quaternary Research, 31, 119-134.
Petit, J.R, Basile, I., Leruyuet, A., Raynaud, D., Lorius, C., Jouzel, J., Stievenard, M., Lipenkov, V.Y., Barkov, N.I., Kudryashov, B.B., Davis, M., Saltzman, E. & Kotlyakov, V. (1997). Four climate cycles in Vostok ice core. Nature, 387, 359-360.
Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J-M., Basile, I., Bender, M. & Chappallaz, J. (1999). Climate and atmospheric history of the past 420,000 years from the Vostok ice core, AntarctiC. Nature, 399, 429-436.
Pickard, J. (1982). Holocene winds of the Vestfold Hills, Antarctica. New Zealand Journal of Geology and Geophysics, 25, 353-358.
Pickard, J. (1985). The Holocene fossil marine macrofauna of the Vestfold Hills, East Antarctica. Boreas, 14, 189-202.
Pickard, J. & Seppelt, R.D. (1984). Holocene occurrence of the moss bryum algens Card. in the Vestfold Hills, Antarctica. Journal of Bryology, 13, 209-217.
Pickard, J., Adamson, D.A. & Heath, C.W. (1986). The evolution of Watts Lake, Vestfold Hills, East Antarctica, from marine inlet to freshwater lake. Palaeogeography, Palaeoclimatology, Palaeoecology, 53, 271-288.
Pickard, J., Selkirk, P.M. & Selkirk, D.R. (1984). Holocene climates of the Vestfold Hills, Antarctica and Macquarir Island. In: Vogel, J.C. (ed.), Late Cainozoic Paleoclimates of the Southern Hemisphere. Rotterdam, A.A.Balkema, 173-182.
Pope, P.G. & Anderson, J.B. (1992). Late Quaternary glacial history of the northern Antarctic Peninsula´s western continental shelf: evidence from the marine record. American Geophysical Union, Antarctic Research Series, 57, 63-91.
Powell, R.D., Hambrey, M.J. & Krissek, L.A. (1998). Quaternary and Miocene glacial and climatic history of the Cape Roberts drillsite region, Antarctica. Terra Antarctica, 5, 341-351.
Prentice, M.J. & Fastook, J.L. (1990). Late Neogene Antarctic ice sheet dynamics; ice model-data convergence. EOS, 71, 1378.
Prentice, M.J. & Mathews, R.K. (1988). Cenozoic ice-volume history: development of a composite oxygen isotope record. Geology, 17, 963-966.
Pudsey, C.J. & Evans, J. (2001). First survey of Antarctic sub-ice shelf sediments reveals mid-Holocene ice shelf retreat. Geology, 29, 787-790.
Pudsey, C.J., Barker, P.F. & Larter, R.D. (1994). Ice sheet retreat from the Antarctic Peninsula Shelf. Continental Shelf Research, 14, 1647-1675.
Pudsey, C.J., Domack, E.W., Evans, J., Morris, P. & Del Valle, R.A. (2001). Bathymetry and acoustic facies beneath the former Larsen-a and Prince Gustav Ice Shelves, NW Weddell Sea. Antarctic Science, 13, 312-322.
Quilty, P.G. (1992). Late Neogene sediments of coastal East Antarctica. In: Yoshida, Y., Kaminuma, K. & Shiraishi, K. (eds.), Recent Progress in Antarctic Earth Science. Tokyo, Terra Scientific Publishing Company, 699-705.
Rabassa, J. (1983). Stratigraphy of the glacigenic deposits in northern James Ross Island, Antarctic Peninsula. In: Evenson, E., Sclüchter, C. & Rabassa, J. (eds.), Tills and Related Deposits. Rotterdam, A.A. Balkema Publishers, 329-340.
Richter, W. & Bormann, P. (1995). Geomorphology. In: Bormann, P. & Fritzsche, D. (eds.), The Schirmacher Oasis, Queen Maud Land, East Antarctica, and its surroundings. Gotha, Justus Perthes Verlag, 171-190.
Robert, C. & Kennett, J.P. (1994). Antarctic subtropical humid episode at the Paleocene-Eocene boundary; clay mineral evidence. Geology, 22, 211-214.
Roberts, D. & McMinn, A. (1996). Relationships between surface sediment diatom assemblages and water chemistry gradients in saline lakes of the Vestfold Hills, Antarctica. Antarctic Science, 8, 331-341.
Roberts, D. & McMinn, A. (1999). A diatom-based palaeosalinity history of Ace Lake, Vestfold Hills, Antarctica. The Holocene, 9, 401-408.
Roberts, D., McMinn, A. & Zwartz, D. (2000). An initial palaeosalinity history of Jaw Lake, Bunger Hills, based on a diatom-salinity transfer function applied to sedimentary cores. Antarctic Science, 12, 172-176.
Rott, H., Rack, W., Nagler, T. & Skvarca, P. (1998). Climatically induced retreat and collapse of northern Larsen Ice Shelf, Antarctic Peninsula. Annals of Glaciology, 27, 86-92.
Rozycki, S. Z. (1961). Changements Pleistocenes de l´extesion de l´Inlandis en Antactide Orientale d´aprčs l´étude des anciennes plages élevées de l´Oasis Bunger, Queen´s Mary Land. Biuletyn Peryglacjalny, 10, 257-283.
Sawagaki, T. & Hirakawa, K. (1997). Erosion of bedforms by subglacial meltwater, Soya Coast, East Antarctica. Geografiska Annaler, 79(A), 223-238.
Scambos, T. A. (2000). The link between climate warming and break-up of ice shelves in the Antarctic Peninsula. Journal of Glaciology, 46, 516-530.
Scherer, R.P., Aldahan, A., Tulaczyk, S., Possnert, G., Engelhardt, H. & Kamb, B. (1998). Pleistocene collapse of the West Antarctic Ice Sheet. Science, 281, 82-85.
Schmidt, R., Mäusbacher, R. & Müller, J. (1990). Holocene diatom flora and stratigraphy from sediment cores of two Antarctic lakes (King George Island). Journal of Paleolimnology, 3, 55-74.
Sedwick, P.N., Harris, P.T., Robertson, L.G., McMurtry, G.M., Cremer, M.D. & Robinson, P. (2001). Holocene sediments from the continental shelf of MacRobertson Land, East Antarctica. Paleoceanography, 16, 212-225.
Schackleton, N.J. & Kennett, J.P. (1975). Paleotemperature history of the Cainozoic and the initiation of Antarctic glaciation: oxygen and carbon analyses in DSDP sites 277, 279 and 281. Initial Reports of the Deep Sea Drilling Project, 29, 743-755.
Shevenell, A.E., Domack, E.W. & Kernan, G.M. (1996). Record of Holocene palaeoclimate change along the Antarctic Peninsula: evidence from glacial marine sediments, Lallemand Fjord. Papers and Proceedings of the Royal Society of Tasmania, 130, 55-64.
Shipp, S.J. & Anderson, J.B. (1994). High-resolution seismic survey of the Ross Sea continental shelf: implications for ice-sheet retreat behaviour. Antarctic Journal of the United States, 29(5), 137-138.
Shipp, S.J., Anderson, J.B. & Domack, E.W. (1999). Seismic signature of the Late Pleistocene fluctuations of the West Antarctic Ice Sheet system in Ross Sea: a new perspective, Part I. Geological Society of America Bulletin, 111, 1486-1516.
Shipp, S.J., Anderson, J.B., DeSantis, L, Bartek, L.R. Alonso, B. & Zayatz, I. (1994). High- to intermediate resolution seismic stratigraphic analysis of mid-late Miocene to Pleistocene strata in eastern Ross Sea: implications for changing glacial/climate regime. Terra Antarctica, 1, 381-384.
Skvarca, P. Rack, W., Rott, H. & Donangelo, T.L. (1999). Climatic trends and the retreat and disintegration of ice shelves on the Antarctic Peninsula: An overview. Polar Research, 18, 151-157.
Sloan, B.J., Lawer, L.A. & Anderson, J.B. (1995). Seismic stratigraphy of the Palmer Basin. Antarctic Research Series, 68, 235-260.
Smith, R.C., Stammerjohn, S.E. & Baker, K.S. (1996). Surface air temperature variations in the western Antarctic Peninsula region. In: Ross, R.M., Hoffman, E.E. & Quetin, L.B. (eds.), Foundations for Ecological Research West of the Antarctic Peninsula. Washington, American Geophysical Union, 105-121.
Smith, R.C., Ainley, D., Baker, K., Domack, E., Emslie, S., Fraser, S., Kennett, Levanter, A., Mosley-Thompson, E., Stammerjohn, S. & Vernet, M. (1999). Marine ecosystem sensitivity to climate change. BioScience, 49, 393-404.
Speir, T.W. & Cowling, J.C. (1984). Ornithogenic soils of the Cape Bird Adélie penguin rookeries, Antarctica. 1. Chemical properties. Polar Biology, 2, 199-205.
Squyres, S.W., Andersen, D.W., Nedell, S.S. & Wharton Jr., R.A. (1991). Lake Hoare, Antarctica: sedimentation through thick perennial ice cover. Sedimentology, 38, 363-379.
Stackebrandt, W. (1995). Moraines around lake Untersee - indicators of the Late Quaternary regional glacial history. In: Bormann, P. & Fritzsche, D. (eds.), The Schirmacher Oasis, Queen Maud Land, East Antarctica, and its surroundings. Gotha, Justus Perthes Verlag, 237-242.
Stark, P. (1994). Climatic warming in the central Antarctic Peninsula area. Weather 49, 215-220.
Steig, E.J., Brook, E.J., White, J.M.C., Sucher, C.M., Bender, M.L., Lehman, S.J., Morse, D.L., Waddington, E.D. & Clow, G.D. (1998). Synchronous climate changes in Antarctica and the North Atlantic. Science, 282, 92-95.
Steig, E.J., Morse, D.L., Waddington, E.D., Stuiver, M, Grootes, P.M., Mayewski, P.A., Twickler, M.S. & Whitlow, S.I. (2000). Wisconsinan and Holocene climate history from an ice core at Taylor Dome, western Ross Embayment, Antarctic. Geografiska Annaler, 82A, 213-232.
Stuiver, M., Denton, G-H., Hughes, T.J. & Fastook, J.L. (1981). History of the marine ice sheet in West Antarctica during the last deglaciation: A working hypothesis. In: Denton, G.H. & Hughes, T.J. (eds.), The Last Great Ice Sheets. New York, John Wiley & Sons, 319-436.
Sugden, D.E. & Clapperton, C.M. (1977). The maximum ice extent on island groups in the Scotia Sea, Antarctica. Quaternary Research, 7, 268-282.
Sugden, D.E. & Clapperton, C.M. (1981). An ice-shelf moraine, George IV Sound, Antarctica. Annals of Glaciology, 2, 135-141.
Sugden, D.E. & John, B.S. (1973). The ages of glacier fluctuations in the South Shetland Islands, Antarctica. In: van Zinderen Bakker, E.M. (ed.), Palaeoecology of Africa, the surrounding Islands and Antarctica, 8. Cape Town: Balkema, 141-159.
Sugden, D.E., Marchant, D.R. & Denton, G.H. (1993). The case for a stable East Antarctic Ice Sheet: The background. Geografiska Annaler, 75A, 151-153.
Taylor, F. & McMinn A. (2001). Evidence from diatoms for Holocene climate fluctuation along the East Antarctic margin. The Holocene, 11, 455-466.
Taylor, F., Whitehead, J. & Domack, E. (2000). Holocene paleoclimate change in the Antarctic Peninsula: evidence from diatom, sedimentary and geochemical record. Marine Micropaleontology, 41: 25-43
Thomas, R.H. & Bentley, C.R. (1978). A model for Holocene retreat of the West Antarctic Ice Sheet. Quaternary Research, 10, 150-170.
Thompson, L.G., Davis, M.E., Mosley-Thompson, E., Sowers, T.A., Henderson, K.A., Zagorodnov, V.S., Lin, P.N., Mikhalenko, V.N., Campen, R.K., Bolzan, J.F., Cole-Dai, J. & Francou, B. (1998). A 25,000-year tropical climate history from Bolivian ice cores. Science, 282, 1858-1864.
Trivelpiece, W.Z. & Fraser, W.R. (1996). The breeding biology and distribution of Adélie penguins: Adaptations to environmental variability. In: Ross, R.M., Hoffman, E.E. & Quetin, L.B. (eds.), Foundations for Ecological Research West of the Antarctic Peninsula. Washington, American Geophysical Union, 273-285.
Vaughan, D.G. & Doake, C.S.M. (1996). Recent atmospheric warming and retreat of ice shelves on the Antarctic Peninsula. Nature, 379, 328-331.
Vaughan, D.G. (2000). Bedmap subglacial and seabed topography for Antarctica. Cambridge, British Antarctic Survey (www.antarctica.ac.uk/BAS_Science/Highlights/2000/bedmap.html).
Verkulich, S.R. & Hiller, A. (1994). Holocene deglaciation of Bunger Hills revealed by 14C measurements on stomach oil deposits in snow petrel colonies. Antarctic Science, 6, 395-399.
Verkulich, S.R. & Melles, M. (1992). Composition and palaeoenvironmental implications of sediment in a fresh water lake and in marine basins of Bunger Hills, East Antarctica. Polarforschung, 60, 169-180.
Vincent, W.F., Howard-Williams, C. & Broady, P.A. (1993). Microbial communities and processes in Antarctic flowing waters. In: Friedman, E.I. (ed.), Antarctic Microbiology. New York, John Wiley & Sons, 543-569.
Waitt, R.B. (1983). Thicker West Antarctic ice sheet and peninsula ice cap in late-Wisconsin time – sparse evidence from northern Lassiter Coast. Antarctic Journal of the United States, 18, 91-93.
Wasell, A. (1993). Diatom stratigraphy and evidence of Holocene environmental changes in selected lake basins in the Antarctic and South Georgia. Stockholm University: Department of Quaternary Geology Research Reports, 23, 1-15.
Wasell, A. & Hĺkansson, H. (1992). Diatom stratigraphy in a lake on Horseshoe Island, Antarctica: a marine-brackish-fresh water transition, with comments on the systematics and ecology of the most common diatoms. Diatom Research, 7, 157-194.
Webb, P.N. & Harwood, D.M. (1991). Late Cenozoic glacial history of the Ross Embayment, Antarctica. Quaternary Science Reviews, 10, 215-223.
Weber, M.E., Bonani, G. & Fütterer, K.D. (1994). Sedimentation processes within channel-ridge systems, southeastern Weddell Sea, Antarctica. Paleoceanography, 9, 1027-1048.
Webster, J., Hawes, I., Downes, M., Timperley, M. & Howard-Williams, C. (1996). Evidence for regional climate change in the recent evolution of a high latitude pro-glacial lake. Antarctic Science, 8, 49-59.
Weiss, R.F., Oesterlund, H.G. & Craig, H. (1979). Geochemical studies of the Weddell Sea. Deep-Sea Research, 26(10A), 1093-1120.
Wharton, R.A., McKay, C.P., Clow, G.D., Andersen, D.T., Simmons, G.M. & Love, F.G. (1992). Changes in ice cover thickness and lake level of Lake Hoare, Antarctica – implications for local climate change. Journal of Geophysical Research, 97, 3503-3513.
White, J.W.C. & Steig, E.J. (1998). Timing is everything in a game of two hemispheres. Nature 394, 717-718.
Whitehouse, I.E., Chinn, T.J.H., Höfle, H.-C. & McSaveney, M.J. (1987). Radiocarbon contaminated penguin bones from Terra Nova Bay, Antarctica. New Zealand Antarctic Record, 8(3), 11-23.
Whitehouse, I.E., Chinn, T.J.H. & Höfle, H.-C. (1989). Radiocarbon dates from raised beaches, Terra Nova Bay, Antarctica. Geologisches Jarbuch, E38, 321-334.
Wilson, G.S. (1995). The Neogene East Antarctic ice sheet: a dynamic or stable feature. Quaternary Science Reviews, 14, 101-123.
Wilson, G.S., Roberts, A.P., Verosub, K.L., Florindo, F. & Sagnotti, L. (1998). Magnetostratigraphic chronology of the Eocene-Oligocene transition in the CIROS-1 core, Victoria Land margin, Antarctica: implications for Antarctic glacial history. Geological Society of America Bulletin, 110, 35-47.
Wisniewski, E. (1983). Bunger Oasis: the largest ice-free area in the Antarctic. Terra, 95, 178-187.
Yang, S.Y. & Harwood, D.M. (1977). Late Quaternary environmental fluctuations based on diatoms from Yanou Lake, King George Island, Fildes Peninsula, Antarctic. In: Ricci, C.A. (ed.), The Antarctic Region: Geological Evolution and Processes. Siena, Universtŕ degli Studi di Siena, 853-859.
Yoon, H. I., Park, B. K., Kim, Y. & Kim, D. (2000). Glaciomarine sedimentation and its palaeoceanographic implications along the fjord margins in the South Shetland Islands, Antarctica, during the last 6000 years. Palaeogeography, Palaeclimatology, Palaeoecology, 157: 189-211.
Yoshida, Y. (1983). Physiography of the Prince Olav and the Prince Harald Coasts, East Antarctica. Memoirs of National Institute of Polar Research, Series C, Earth Science, 13, 83 pp.
Yoshida, Y. & Moriwaki, K. (1979). Some consideration on elevated coastal features and their dates around Syowa Station, Antarctica. Memoirs of National Institute of Polar Research, Special Issue, 13, 220-226.
Yoshikawa, T. & Toya, H. (1957). Report on geomorphological results of the Japanese Antarctic Research Expedition, 1956-57. Nankyoku Shiryou, 1, 1-13.
Zale, R. (1993). Lake sediments around the Antarctic Peninsula, archives of climatic and environmental change. GERUM Naturgeogr. , 17, 1-21.
Zale, R. (1994a). Changes in size of the Hope Bay Adélie penguin rookery as inferred from Lake Boeckella sediment. Ecography, 17, 297-304.
Zale, R. (1994b). 14C age correction in Antarctic lake sediments inferred from geochemistry. Radiocarbon, 36, 173-185.
Zale, R. & Karlén, W. (1989). Lake sediment cores from the Antarctic Peninsula and surrounding islands. Geografiska Annaler, 71(A), 211-220.
Zhang, Q.S. (1992). Late Quaternary environmental changes in the Antarctic and their correlation with global change. In: Yoshida, Y., Kaminuma, K. & Shiraishi, K. (eds.), Recent Progress in Antarctic Earth Science. Tokyo, Terra Scientific Publishing Company, 781-785.
Zwartz, D., Bird, M., Stone, J. & Lambeck, K. (1998). Holocene sea-level change and ice-sheet history in the Vestfold Hills, East Antarctica. Earth and Planetary Science Letters, 155, 131-145.