Introduction

One key objective of the West Antarctic ice sheet (WAIS) initiative is documentation of eustatic sea-level change resulting from the possible disintegration of the West Antarctic ice sheet (Reference Bindschadler,Bindschadler, 1991). This effort requires detailed knowledge of ice-sheet history in West Antarctica, including the history of grounding-line and calving-margin retreat to the positions now occupied by these features and an accurate dynamic overview of the West Antarctic ice sheet, including present and former precipitation sources, flow regimes and changing interactions with ice derived from East Antarctica. The development of reliable predictive models for future ice-sheet retreat depends on these data. Yet, so far most WAIS studies have concentrated either on modern ice dynamics in West Antarctica or on elucidating former ice maxima.

This paper presents a series of preliminary observationally based reconstructions of interactions between the East and West Antarctic ice sheets during and since the last glacial maximum (LGM). These two ice sheets are separated by a dividing wall (the Transantarctic Mountains) and interact today primarily where through-flowing outlet glaciers from East Antarctica impinge on the Ross Ice Shelf. Our reconstructions are based on glacial geologic observations and the observed present-day ice dynamics of the McMurdo Ice Shelf (MIS), a part of the Ross Ice Shelf. The reconstructions aid in elucidating modern glaciological processes operating on this ice shelf and support suggestions by Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others (1990) that some existing MIS features are remnants from a former ice-flow regime. MIS data permit evaluation of two different published reconstructions of the flow regime during the LGM, and development of four new reconstructions for the Ross Sea embayment during the deglacial transition from the grounded late Wisconsin Ross Sea ice sheet to the modern Ross Ice Shelf, all of which include ice derived from both East Antarctica and West Antarctica.

West Antarctica is a key area for studying glacial history and dynamics, because the world’s largest extant marine-based ice sheet is located there. Of particular importance are the Dry Valleys of southern Victoria Land, situated between the seasonally open waters of the Ross Sea and McMurdo Sound on the east and the East Antarctic ice sheet to the west. These ice-free valleys are shielded from direct East Antarctic ice-sheet influence by the Transantarctk Mountains. It can be argued that a third major factor, in addition to East and West Antarctic ice fluctuations, operates in this area: local glaciers respond to the proximity of an open-water moisture source in the Ross Sea. This influence is actually controlled by West Antarctic ice extent, being insignificant during glacial maxima and reaching its greatest importance during interglacials when shelf- and sea-ice limits are least extensive. The availability of a local moisture source is important for this paper primarily with respect to fluctuations of Koettlitz and Skelton Glaciers. A detailed late Quaternary paleoclimatic record for the Dry Valleys region (Reference Denton,, Armstrong,, Stuvier and Turekian,Denton and others, 1971, Reference Denton,, Bockheim,, Wilson, and Stuvier.1989; Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others, 1981a) shows that the dominant features were periodic advances of the West Antarctic ice sheet, which thickened and grounded in the Ross Sea and McMurdo Sound, pushing lobes of grounded ice into the Dry Valleys and leaving a record of ice-dammed lakes, marginal and recessional moraines and raised beaches. Numerous conventional ¹⁴C dates on marine shells and terrestrial fresh-water algae suggest a late Quaternary age of ∼18-20 BP for the most recent advance of West Antarctic ice (Ross Sea ice sheet). Reference Drewry,Drewry (1979) suggested that grounded ice extent was considerably less extensive than portrayed by Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others (1981a) and hence that their ice-flow trajectories were unlikely. In particular, Reference Drewry,Drewry (1979) called for northward flow of grounded ice past Ross Island in McMurdo Sound, with no westward flow into the Dry Valleys.

The mechanism and history of deglaciation following the late Wisconsin ice maximum in West Antarctica is a significant glacial geologic problem, involving ice thinning and grounding-line retreat of the Ross Sea ice sheet in the central and outer Ross Sea. and flotation to form the Ross Ice Shelf in the central and southern Ross Sea. McMurdo Sound is of particular importance for understanding this transition, because apparent remnants of former grounded ice are preserved as part of the MIS (Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others, 1990).

Glacial Geology of the Mcmurdo Ice Shelf

The MIS is an extension of the Ross Ice Shelf that occupies southern McMurdo Sound (Reference Stuart, and BullStuart and Bull, 1963). It is bounded on the south by Minna Bluff, on the north by Ross Island and the seasonally open waters of MuMurdo Sound, on the west by Mount Discovery and the Dry Valleys of southern Victoria Land, and on the east by the Ross Ice Shelf and White Island (Figs 1 and 2). Koettlitz Glacier flows into the MIS west of Brown Peninsula. The eastern part of the MIS is dominated by net annual surface accumulation, like the remainder of the Ross Ice Shelf, but the western part is characterized by surface ablation (Reference Swithinbank,Swithinbank, 1970). Prominent bands and patches of sediment debris mantle the MIS surface in this area. The following extended discussion concentrates on the western ablation area and is based largely on Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others (1990).

Fig. 1. Index mар showing locations mentioned in text. See Figure. 2 for details of the MIS. Symbols: B = Blue Glacier; CB = Cape Bird; CR = Cape Royds; F = Ferrar Glacier; K = Koettlitz Glacier; MP = Marble Point.

Fig. 2. McMurdo Ice Shelf (MIS) area, showing locations mentioned in text and informal names for study areas (Black Island debris bands (BIDB), “Swirls”, central area, etc.). Note that because the MIS is so complex, only major debris bands and features are shown. Base map used here is modified from Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others (1981a), is based on U.S. Geological Survey topographic maps, aerial photographs and material from other sources.

The MIS debris bands comprise sediment (ranging in size from clay to boulders but dominated by sand and gravel), marine biotic remains (including: siliceous sponges, molluscs, barnacles, corals, bryozoans, Foram-inifera. diatoms and fish), and non-marine algae and diatoms that live in seasonal melt ponds of the MIS surface. Reference Debenham,Debenham (1919) proposed that marine sediment and biota were incorporated into the MIS by basal adfreezing, to emerge eventually on the MIS surface because of net annual surface ablation. Subsequent workers, who have reported biotic remains on the MIS, include: Reference Debenham,Debenham, 1949, Reference Denbenham,1961; Reference Swithinbank,Swithinbank and others, 1961; Reference Speden,Speden, 1962; Reference Gow,Gow and others, 1965; Reference Kellogg,, Stuvier,, Kellogg, and Denton,Kellogg and others, 1977; Reference Brady,Brady, 1978; Reference Hayward, and Taylor,Hayward and Taylor, 1984; Reference Kellogg, and Kellogg,Kellogg, 1987; Reference Howard-Williams,, Pridmore,, Downes, and Vincent,Howard-Williams and others, 1989, Reference Howard-Williams,, Pridmore,, Broady,, Vincent, and Kerry,1990; Reference Hart,Hart, 1990; Reference Suren,Suren, 1990. Reference Gow, and Epstein.Gow and Epstein (1972), Reference Suren,Stuiver and others (1981b) and Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others (1990) used oxygen-isotopie analyses of MIS ice to show that lowermost Koettlitz Glacier and most ice in the surface-ablation zone formed from frozen sea water, supporting Debenham’s hypothesis.

Ice-thickness data obtained by radio echo-sounding (Reference Swithinbank,Swithinbank, 1970) and drilling Reference Gow, and Epstein.Gow and Epstein, 1972) are shown in Figure. 3. Thicknesses between Mount Discovery and Black Island and north of Minna Bluff exceed 100 m. Further north, most thicknesses range from 20 to 50 m. This information, together with observed surface-ablation rates (0.5-1.0 m year⁻¹; Reference Gow,Gow, 1967; personal Communication from A. L. DeVries, 1987), suggests that ice in the northern part of die MIS must be less than 100 years in age but the presence of fossil shells (¹⁴C ages ≤7750BP; all dates in this paper are uncorrected) suggests that basal freezing and surface ablation must be nearly in equilibrium. The MIS probably did not disintegrate completely during the Holocene because ¹⁴C dated shells have progressively older ages towards the north along debris bands.

Fig. 3. MuMurdo Ice Shelf thickness and ice-flow data (after Reference Kellogg,, Kellogg,, Stuvier and Bentley,kellogg and others, fig. 17). Solid dots and arrows are thickness (in m) and velocity (m year⁻¹) measurements (Reference Swithinbank,Swithinbank, 1970), respectively; solid squares are drillholes (Reference Gow, and Epstein.Gow and Epstein, 1972), except 45 m ice depth near the easternmost Dailey Islands which was extrapolated by Reference Zotikov,, Gow, and Oura,Zotiko and Gow (1967); solid triangles are thicknesses rounded to the nearest 5 m from seismic soundings (Reference Cook,Сook, 1963); 35 m depth is from a drillhole through the ice shelf (Reference Paige,Paige, 1968). Dashed contour lines north of Minna Bluff are ice thicknesses (U.S. Geological Survey, 1972). Dashed lines labeled “A”, “B” and “C” are from Reference Swithinbank,Swithinbank (1970). “A” marks the boundary between rough ice with melt streams and considerable relief (in the west) and much smoother bare ice. “B” approximates the surface equilibrium line and is extended into the Minna Bluff area based on our observations. See Reference Swithinbank,Swithinbank (1970) for significance of “C”. Additional thickness and velocity data for the eastern part of the MIS (accumulation area) have been given by Reference McCrae,McCrae (1984). Ice-frontal position in 1947 is probably incorrect as explained by Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others (1990). Dashed lines on Brown Peninsula. Mount Discovery and Black Island are maximum extent of Ross Sea Drift (Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others, 1981a).

Very little ice is adverted today into the western MIS from the Ross Ice Shelf (Reference Swithinbank,Swithinbank, 1970; Reference Bentley,, Clough,, Jezek, and ShabtaieBentley and others, 1979) or from Koettlitz Glacier (Reference Gow, and Epstein.Gow and Epstein, 1972). Hence, the mass balance of this area is dominated by basal freezing and surface ablation, the latter caused primarily by adiabatically warmed katabatic winds flowing down on to the MIS from Minna Bluff and Mount Discovery (Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others, 1981a). Ice-velocity measurements in the western MIS are less than 5 m year⁻¹ towards the north between Mount Discovery and Black Island and north of the strait between White and Black Islands (Reference Swithinbank,Swithinbank, 1970). Higher velocities of 18-22 m year⁻¹ directed west-northwest to northwest were recorded near the calving margin between Ross Island and the Dry Valleys (Reference Swithinbank,Swithinbank, 1970). Most ice movement north of Black Island is towards the north or northwest, parallel to debris bands, especially along the striking debris band extending ∼30 km from Black Island to the calving margin Black Island debris bands: BIDB), because of the progression of ¹⁴C ages mentioned above (Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others, 1990). However, an anonymous reviewer suggested that flow might have been perpendicular to debris bands, in which case debris comprising the BIDB might have originated along the coasts of Black and White Islands. This “normal flow model” permits postulation of a flow regime entirely different from that discussed above, with different locations for debris and fossil Origin, and hence can have a significant bearing on interpretations of present and former ice flow lines.

Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others 1990 recognized two distinctly different types of debris bands on the MIS. The first group, discussed above, is characterized by thin (usually < 10cm thick) debris overlying clean shelf ice. in which marine shells yield ¹⁴C ages ranging up to 7700 BP. Ice underlying debris in this young group of debris bands usually has δ¹⁸O values ranging from −5.0 to +5.0%, indicating a marine origin. The presumed northward pattern of modern ice movement has been discussed above.

The second type of debris band is located primarily along the east coast of Brown Peninsula in the “Swirls” area (Fig. 2). Debris thickness here appears to exceed 10 cm but its full extent has not been determined, because it is ice-cemented below 5-10 cm. A significant part of the boulders here have lithologies indicating transport from the Transantarctic Mountains. Some of the bands cross the tide crack along the MIS grounding line and extend on to the flanks of Brown Peninsula and even cross the Brown Peninsula saddle (Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others, 1981a). Some shell dates from the northern part of the “Swirls” and adjacent bands and patches that are thought to have been former parts of the “Swirls” area yield ¹⁴C ages > 20000 BP. Some ice samples from this area yield δ¹⁸O values ranging from −25.35 to 48.96%, suggesting the presence of patches of glacial ice that formed at elevations above 3000 (Reference Morgan,Morgan, 1982) and was adverted to southern McMurdo Sound. The “central” area, located between Brown Peninsula and Black Island, where thick debris is also present, may also represent relict debris bands. The sparse ¹⁴C and δ¹⁸O data available from this area are not consistent with remnant ice but additional field work is required to test this hypothesis.

Sedimentologic, ¹⁴C and δ¹⁸O data described above, combined with the observed pattern of debris bands, suggest that the “Swirls” area includes deformed remnants from the grounded late Wisconsin Ross Sea ice sheet. Glacial flow was apparently directed from east to west across southern MeMurdo Sound, with an ablating margin pushing into the mouths of the Dry Valleys to the west (Fig. 4). The distribution of late Wisconsin-aged Ross Sea drift on Black Island and Brown Peninsula suggests that high peaks protruded through the ice sheet as nunataks (Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others, 1981a); Black and White Islands, and Brown Peninsula, probably served as effective barriers, retarding flow rates. Debris entrained in the grounded ice probably accumulated on the surface in the McMurdo Sound ablation area, producing some of the debris bands now preserved in the “Swirls”, because the prevailing surface-ablation regime here (Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others. 1981a) resulted in upward-directed ice flowlines. This concludes our summary of previous work on the MIS (Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others, 1990).

Fig. 4. Surface contours (solid lines; elevations in m a.s.l.) and flowlines (dotted) of the late Wisconsin Ross Sea ice sheet in McMurdo Sound (after Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others, 1981a. (Fig. 7-16).

We turn now to a discussion of possible ramifications of MIS and other geologic data for recognizing West Antarctic ice retreat. Grounded ice covered most of the Ross Sea continental shelf during the late Wisconsin maximum, as discussed later in this paper (Fig. 7). Deglaciation to the present configuration, with the floating Ross Ice Shelf covering the southern half of the Ross Sea, probably involved glacial thinning and ungrounding. We should be able to recognize two events at any particular location north of the present ice-shelf front: the times when (1) the grounding line and (2) the calving ice-shelf margin retreated past the location. In principle, ¹⁴C dating of basal postglacial sediments in Ross Sea cores should provide a detailed picture of grounding-line recession. This has not proved possible with conventional ¹⁴C dating because Ross Sea sediments have a low organic carbon content and calcareous shells are rare, and few AMS dates have been obtained. However, ¹⁴C dates from the MIS also provide limits on the timing of grounding-line recession. All these dates are either >20 000 BP or <7750 BP. If the absence of dates between these ages represents the time when grounded ice filled southern McMurdo Sound, the younger population of dates records events following thinning, grounding-line recession and the incursion of marine waters beneath the MIS. The older population provides minimum dates for preglacial marine conditions. Thus, the grounding line had receded to the vicinity of Brown Peninsula saddle by 7750 BP.

Calving-margin recession may be tracked using ¹⁴C dated raised beaches, which require seasonal open water to form. Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others (1981a) dated numerous raised beaches along the Ross Sea coast and offlying islands, and concluded that the calving margin extended from Marble Point to a position north of Beaufort and Franklin Islands shortly before 5650 CBP and had retreated to a position from Gape Bernacchi to near Mount Bird on Ross Island shortly after 5650 ¹⁴C BP. The first observations of the MIS during Scott’s 1910–13 expedition showed the western part of the calving margin extending from Cape Chocolate to near the present terminus of the Black Island debris band; the eastern part was much like today. Calving-margin retreat since 1910-13 has occurred primarily in this western section.

During the period from the LGM to about 7750 BP, thinning probably continued to concentrate englacial debris at the surface in southern McMurdo Sound. In addition, as the ice wasted, exposing more and more of Black and White Islands and Brown Peninsula, the ice-flow pattern changed. The former east-to-west pattern was probably replaced by nearly stagnant conditions or sluggish northward movement (on the order of a few m year⁻¹) between the islands and Brown Peninsula, and existing debris islands were probably deformed or bent into the shapes that give the “Swirls” area its name. Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others (1990) postulated that while most of the remainder of the MIS became progressively afloat, the “Swirls” remained grounded, because it rested on a wide shallow shelf east of Brown Peninsula and that ungrounding here is a very recent event.

Glaciological Reconstructions

Ice elevations and flowlines for our glaciological reconstructions are based on data for the area between 150° E and 180° and from 70° to 83°S (Fig. 1). that is, the Transantarctic Mountains from northern Victoria Land south to near Nimrod Glacier and adjacent areas of East and West Antarctica. We reconstructed ice flowlines and ice elevations in this sector to conform with these field data, not with a theoretically based computer ice-sheet model. Figure. 5 shows the bedrock topography for the modern ice sheet (Reference Drewry,Drewry, 1983) that constrains our reconstructions.

Fig. 5. Modern subglacial topography (after Reference Drewry,Drewry, 1983), contoured at 500 m intervals. Present sea level is shown by a heavy solid line where not covered by the ice sheet, and elsewhere by a heavy dashed line. Solid pattern indicates elevations above 1000 m; stippled pattern indicates depths greater than 1000 m. Note that these great depths are restricted to outlet-glacier troughs now occupied by Byrd, Mulock, David and Rennick Glaciers. Present northern margin of the Ross Ice shelf is indicated by a light dashed line.

The modern ice-sheet profile (Fig. 6) and the LGM reconstruction (Fig. 7) represent two end members in our series of reconstructions. They are constrained by observational data and/or a wealth of glacial geologic information collected along the Victoria Land coast and on offlying islands (e.g. Reference Denton,, Armstrong,, Stuvier and Turekian,Denton and others, 1971, Reference Denton,, Bockheim,, Wilson, and Stuvier.1989; Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others, 1981a, Reference Suren,b; Reference Kellogg,, Kellogg,, Stuvier and Bentley,Kellogg and others, 1990; Reference Baroni, and Orombelli.Baroni and Orombelli, 1991, Reference Baroni, and Orembelli1994; Reference Dochat,Dochat, 1994) including ages and elevations of Ross Sea drift and raised beaches, and locations of kenyite erratics, These data help to constrain LGM ice elevations along the coast of Victoria Land and in some cases suggest former flow directions. More geologic data from the Ross Sea would be highly desirable to refine our work but are currently unavailable. Our reconstructions for 16 000, 12 000, 8000 and 4000 BP (Figs 9–12) are less rigorous. They represent sequential modifications of the LGM reconstruction constrained by the limited available glacial geologic data.

Fig. 6. Present-day ice elevations (after Reference Drewry,Drewry 1983) contoured at 500, 1000 and at 100 m intervals above 2000 m. flowlines indicated by dashed lines. “D” = ice dome; “S” = ice saddle. Stippled pattern indicates bedrock elevations above 1000 m. Note the prominent dome west of northern Victoria Land. A smaller dome (Taylor Dome—not labeled) is situated west of MuMurdo Sound. Flow from Taylor Dome is down Skelton Glacier but terminates south of Minna Bluff. Flow from Mulock Glacier rounds Minna Bluff with most ice passing east of Ross Island.

Fig. 7. LGM ice-sheet reconstruction contoured at 100 m intervals (solid lines). “GL” = grounding line; other symbols and patterns as in Figure. 6. Note the similarity of ice elevations west of the Transantarctic Mountains to modern conditions. Major differences are in West Antarctica where a low-profile ice sheet occupied the entire Ross Sea. 7he prominent saddle north of Ross Island permitted southward ice flow into McMurdo Sound, consistent with the Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others (1981a) reconstruction. Flow from Taylor Dome (west of McMurdo Sound) was down Skelton Glacier but terminated south of Minna Bluff. Flow from Mulock Glacier rounded Minna Bluff and Ross Island before flowing south into McMurdo Sound.

Fig. 9. 16000 BP ice-sheet reconstruction. Symbols, patterns and contours as in Figure. 7. The grounding line has retreated from the continental shelf break, with deep embayments in troughs on the Ross Sea continental shelf, especially north of Coulman Island. An ice shelf still covers much of the northern part of the continental shelf. As discussed in the text, this reconstruction might represent the LGM, if the true LGM grounding line was just north of Coulman Island.

Fig. 10. 12 000 BP ice-sheet reconstruction. Symbols, patterns and contours as in Figure. 7. The grounding line adjacent to northern Victoria Land has retreated to the Vicinity of Cape Hickey and a small ice shelf occupies this embayment north to Drygalski Ice Tongue. Note that the LGM saddle north of Ross Island has been replaced by a dome, hence ice flow through McMurdo Sound (mostly from Mulock and Koettlitz Glaciers) is now northward.

Fig. 11. 8000 BP ice-sheet reconstruction. Symbols, patterns and contours as in Figure. 7. “R” = ice rises, where shelf ice grounds on shallows. The grounding line has retreated to a position in southern McMurdo Sound and all shelf ice in the sound is of local origin or from Koettlitz and Ferrar Glaciers. Position of the ice-shelf margin north of Franklin and Beaufort Islands is approximate.

Fig. 12. 4000 BP ice-sheet reconstruction. Symbols, patterns and contours as in Figure. 7. The grounding line has retreated to a position south of Minna Bluff and the Ross Ice Shelf margin is approaching its present position. Shelf ice in McMurdo Sound is from Koettlitz Glacier and local precipitation, and some Mulock Glacier ice. Ice-shelf marginal position in McMurdo Sound is poorly constrained but south of Cape Bird.

Modern ice-flow regime

One objective of our work is to develop a model of the dynamics of the modern MIS ice-flow regime. Of particular importance is ice advected into the MIS area, because ice derived from precipitation outside the immediate area may be expected to display different stable-isotope compositions and carry exotic erratic debris. MIS features which are not accounted for by the modern flow regime can be assumed to be relict, especially if the δ¹⁸O of ice in these areas has the signature of ice formed by precipitation at high elevations on the East Antarctic ice sheet (e.g. the “Swirls’” and possibly the central area). Knowledge of the modern flow regime permits us to test and evaluate conflicting hypotheses about grounded ice extent in the Ross Sea during the last glacial maximum (e.g. Reference Drewry,Drewry, 1979; Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others, 1981a).

Modern ice elevations (Reference Drewry,Drewry, 1983) were used to construct ice flowlines for the East Antarctic ice sheet ((Fig. 6)). Three prominent domes arc located in the study area; Talos Dome in northern Victoria Land, Dome Circe west of northern Victoria Land and Taylor Dome west of the Dry Valleys area. East Antarctic ice is discharged through the Transantarctic Mountains at a number of locations (notably Byrd, Mulock and David Glaciers), all of which coincide with deep bedrock troughs (Fig. 5). North of McMurdo Sound and Ross Island, this discharge is more or less directly into the sea, and the ice is removed as floating icebergs. However, south of Ross Island, the East Antarctic discharge feeds the floating Ross Ice Shelf where it is supplemented by local precipitation, These two sources of ice are balanced by bottom melting and calving from the ice-shelf front, at least in historic time. An exception is the western part of the MIS where a peculiar combination of basal adfreezing, surface ablation and calving-margin retreat have been observed as discussed above. Flowlines shown for the Ross Ice Shelf and consistent with observed flow directions (Reference Swithinbank,Swithinbank, 1970. personal communication; U.S. Geological Survey, 1972).

Possible contributions of ice advected to McMurdo Sound can be related to different possible precipitation sources and or ice-elevation fluctuations: (1) Koettlitz Glacier, representing local precipitation in southern Victoria Land, (2) Skelton Glacier, representing mostly local precipitation and precipitation further west on the flank of the East Antarctic ice sheet, and (3) Mulock Glacier, representing a nearly pure high-elevation East Antarctic precipitation source. We combine model results with δ¹⁸O analyses of MIS ice, ¹⁴C dates on MIS shells and MIS boulder-provenance data to evaluate the ways in which this ice-supply pattern changed in the past, during (Fig. 7) and since the LGM (Figs. 9–12). This analysis helps to detail the deglacial transition from the grounded Ross Sea ice sheet to shelf ice and may permit a closer correlation of late Quaternary Dry Valleys and Ross Sea sediment records.

We made some preliminary measurements and calculations for the modern Skelton and Mulock Glaciers drainage basins and the Ross Ice Shelf adjacent to the MIS to supplement the glaciologic model results in the preparation of our LGM reconstruction. Using Reference Drewry,Drewry’s (1983) surface-elevation data, we plotted surface flowlines to determine the Mulock Glacier catchment area. This catchment was subdivided into 15 50 km wide sectors, for each of which we measured area. The total catchment area is ∼95 000 km. We were unable to do a similar analysis for Skelton Glacier, because the drainage area is so small on Reference Drewry,Drewry’s (1983) map and so cluttered with Transantarctic Mountains topography and flight lines that we could not discern surface-elevation contours. Hence, we estimated a total catchment area for Skelton Glacier of ∼10 000 km², a value that we believe over-estimates the true area. Using surface mass-balance data Reference Giovinetto,, Bently, and Bull,Giovinetto and others, 1989, (Fig. 2), we assigned mass-balance rates (kg m⁻² year⁻¹) to each sector to calculate net accumulation. The total calculated for Mulock Glacier is ∼8.5 × 10⁸ kg year⁻¹ and that for Skelton Glacier is ∼1.5 × 10⁸ kg year⁻¹. The total flux through the Transantarctic Mountains was calculated for each glacier using a density of 0.9 Mg m⁻³. Total fluxes are ∼9.5 ×10⁵ m³ year⁻¹ for Mulock Glacier and ∼1.6 × 10⁵ m³ year⁻¹ for Skelton Glacier. In both cases, this volume of ice passes through a narrow channel approximately 10 km wide on to the Ross Ice Shelf. Grounding-line ice thicknesses are approximately 1000 and 850 m, respectively (Fig. 5).

Ice discharged from Skelton and Mulock Glaciers today becomes afloat as it passes through the Transantarctic Mountains. Thickness data for the Ross Ice Shelf (U.S. Geological Survey, 1972) show a complicated pattern related to the numerous ice streams, like Skelton, Mulock and Byrd Glaciers, but there is a general thinning from values close to 500 m where ice streams enter the ice shelf to less than 300 m near the calving margin. This pattern is complicated by surface accumulation and by basal melting, which is apparently very important within ∼200 km of the calving margin (except beneath the western MIS and which causes much of the noted thinning.

Tracing individual ice-stream components to the ice-shelf margin is not trivial because of these complex interactions, What is clear from the U.S. Geological Survey (1972) thickness data is that ice rounding Minna Bluff and flowing west between Minna Bulff and White Island today is not derived from Skelton Glacier but from Mulock Glacier and from local precipitation off Minna Bluff. (Reference Swithinbank,Swithinbank 1970, personal communication) believed that this flow is negligible today, with most Ross Ice Shelf ice passing east of White Island and Ross Island. Our analysis supports this view ((Fig. 8a). Assuming that the Skelton Glacier ice-flow band is 20 km wide, its area is approximately one-third that of the catchment area. Modest basal melting rates, on the order of 10 cm year⁻¹, should remove all the glacial ice before it rounds Minna Bluff. In any case, Skelton Glacier ice could not explain the scattered large negative δ¹⁸O values from the MIS. because the catchment area for this glacier is below 2500 m. In contrast, the Mulock Glacier catchment an a extends well up on to the East Antarctic Ice sheet (EAIS) to elevations over 3000 m. Most Mulock Glacier ice apparently either melts off the base of the Ross Ice Shelf or flows east of Minna Bluff and Ross Island in calve into the Ross Sea. This analysis demonstrates that the MIS ice samples with high negative δ¹⁸O values cannot be related to the modern flow regime.

Fig. 8. Сarloons illustrating different ice-flow patterns in McMurdo Sound: a. Modern ice-flow regime: b. Mid-Holocene (4000-8000 BP), c. LGM modified from Reference Drewry,Drewry (1979). d. LGM modified from Reference Stuvier,, Denton,, Hughes,, Fastook,, Denton, and Hughes,Stuvier and others (1981a).

Koettlitz Glacier today contributes ice only to the area west of Brown Peninsula, and even here, ice formed by precipitation in the local catchment area of southern Victoria Land mostly melts off the base of the glacier soon after it becomes afloat, so that the lower reaches are formed from frozen sea water (Reference Gow, and Epstein.Gow and Epstein, 1972). Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others 1981a suggested that Koettlitz Glacier is probably near its maximum extent today, and was even less advanced during the LGM. However, Koettlitz Glacier may have contributed more ice to the MIS during the Holocene (Fig. 8b), especially if the Ross Sea then was more open and hence a more important precipitation source than today (e.g. Reference Kellogg,, Truesdale, and Osterman,Kellogg and others, 1979).

Discussion

The observation-based reconstructions at the LGM in Figure. 7 and during deglaciation in Figures 9–12 have very low ice-sheet flowline profiles that are as low as or lower than any found in West Antarctica today. This can be explained in two ways. First, when sea level lowered 120 m or more during the last glaciation and caused the Ross Ice Shelf to become grounded in the Ross Sea embayment, the ice surface would have been virtually horizontal and many millennia would be required at present-day rates of accumulation to produce flowline profiles for nearly steady-state ice-sheet flow that are observed today for the East Antarctic ice sheet, especially if the bed became frozen upon grounding and remained frozen. In this case, the low flowline profiles existed because grounding was so late in the glaciation cycle that time was merely insufficient to produce the steeper steady-state flowline profiles typical of East Antarctica today. However, if the bed was thawed and remained thawed when the Ross Ice Shelf grounded, then East Antarctic outlet glaciers through the Transantarctic Mountains (specifically Byrd, Mullock and David Glaciers in our Study) may have continued as ice streams occupying the linear troughs that extend seaward across the continental shelf from these outlet glaciers. In this case, the low flowline profiles existed because water-saturated sediments in the troughs provided low traction that allowed the low flowline profiles typical of West Antarctic ice streams today. As in West Antarctica, these sediments must be either very thick or continuously resupplied in order to prevent fast ice-stream flow from becoming slow ice-sheet flow as glacial erosion transported the sediments seaward, thereby exposing high-traction bedrock in the troughs.

A field test of the two hypotheses would involve finding thick sediment wedges on the continental slope beyond the Ross Sea embayment. If none exist, ice-shelf grounding late in the glaciation cycle may not have allowed sufficient time for ice Streams to deposit their wedges before rising sea level forced grounding lines to retreat into the Ross Sea embayment. Field studies would then not provide a definitive test of the competing hypotheses.

Glaciological theory can provide a test of each hypothesis by using ice-sheet modeling to reconstruct flowline profiles that match the profiles in Figures 7 and 9–12 for the deglaciation hemi-cycle. Following Reference Hughes,Hughes (1992), the surface profile along a flowline can be calculated by combining the force-balance and mass-balance equations. A simultaneous solution of these two equations results in the following expression for the surface slope:

(1)

where Δh is the change of ice height h in horizontal length increment Δx along length x of a flowline, with x = 0 at the grounding line and x = L. at the ice divide, h ₁ is ice thickness along the flowline, w is width of the flow band that encloses the flow line, a is average accumulation rate along the flow band, and h ₀ and u ₀ are the respective grounding-line ice thickness and velocity at x = 0, τ_s is side shear stress alongside the flow band, τ_o is basal shear stress beneath the flow band, g_z is gravity acceleration, ρ_I is ice density, ρ_W is water density, P _I = ρ_Ig _Z h _I is basal ice pressure, P _W = ρ _W g _Z h _W is basal water pressure equivalent to that beneath a water column of height h _W and P _W/P _I = ρ _W h _W/ρ _I h _I is a basal buoyancy factor such that P _W/P _I = 0 for ice-sheet flow over a frozen bed, P _W/P _I = 1 for ice-shelf flow over sea water, and 0 < P _W/P _I < 1 for ice-stream flow over wet sediments. A and n give the respective hardness and viscoplastic properties of ice, the effect of longitudinal deviator stress is contained in the first term, the effect of longitudinal deviator stress gradient is contained in the second term, the effect of transverse shear stress σ _xy is contained in the third term, the effect of vertical shear stress σ _xz is contained in the fourth term, and R when transverse deviator stress is ignored.

Equation (1) can be used to reconstruct surface-ice profiles along the entire length of a flowline, including ice-sheet flow, ice-stream flow and ice-shelf flow, merely by allowing P _W/P _I to vary along the flowline. For example, consider the two extremes of ice-sheet flow over a frozen bed and ice-shelf flow over sea water in a flow band of constant width. Equation (1) gives flowline slopes for linear ice-sheet flow over a frozen bed, for which P _W/P _I = τ_s = 0:

(2)

Equation (1) also gives flowline slopes for linear ice-shelf flow, for which τ_o = 0 and P _W/P _I = 1:

(3)

Equation (2) can be integrated to give the parabolic flowline profile for ice-sheet flow when τ_o is constant and the bed is horizontal. Equation (3) gives the ice-thickness gradient Δh _I/Δx = (Δh/Δx)/(1 – ρ _I/ρ _W) for floating ice that Reference Oerlemans, and van der VeenOerlemans and Van der Veen (1984) derived for R = 1 and τ_s = 0.

The flow law of ice can be integrated to give the average creep velocity at position x along the flowline:

(4)

the sliding law of ice for ice-sheet flow can be modified to include ice-stream flow by including P_W /P_I . The sliding velocity is then (Reference Hughes,Hughes, 1992):

(5)

where B includes bed-roughness factor Λ/Λ′ for ice sliding over bedrock bumps of critical height Λ and spacing Λ′ and m = (n + 1)/2 in the Reference Weertman,Weertman (1957) theory of ice sliding over wet bedrock, P_W /P_I = 0 when no wet sediments cover bedrock, P_W /P_I = 1 when wet sediments bury bedrock bumps and 0 < P_W /P_I < 1 when some bumps are not buried. Note that τ_o → 0 as P _W/P _I → 1, so u_s is indeterminate in Equation (5) when ice-stream flow becomes ice-shelf flow beyond the grounding line.

For ice-sheet flow over bedrock, τ _o = (τ _o)_c for creep over a frozen bed according to Equation (4), and τ _o = (τ _o)_s for sliding over a thawed bed according to Equation (5), where (τ _o)_c > (τ _o)_s, because a frozen bed provides more traction to resist sliding. When a and w are constant and P_W /P_I → 0 along a flowline, Reference Hughes,, Denton, and Hughes,Hughes (1981) showed that (τ _o)_c and (τ _o)_s are:

(6)

(7)

If f is the thawed fraction of the bed, ϕ_s and ϕ_c are factors that allow line variable accumulation rates and flow-band widths along the flowline and P_W /P_I varies along the flow band, because wet sediments cover part or all of bedrock, then (τ_o)_s, in Equation (7) is multiplied by (1 – P _W/P _I)^4m/(2m+1) and τ_o for both ice-sheet flow and ice-stream flow is given by:

(8)

Reference Hughes,, Denton, and Hughes,Hughes (1981) included glacio-isostatic depression of bedrock beneath ice-sheet flowlines by introducing the ratio r of lowered bed elevation to lowered surface elevation due to isostatic depression. For a bed lowering beneath thickening ice:

(9)

where for isostatic equilibrium, t is time since the ice load was applied and t _o is the relaxation time. For a bed rebounding beneath thinning ice:

(10)

where ice began to thin at t = 0 and r _a is given at time t in Equation (9) when thickening ended and thinning began.

Equations (6)–(10) can be incorporated into a numerical solution of Equation (1) by setting Δh = h _i+1 – h_i and h _I = h – h _R, where i is the number of Δx steps from the grounding line to distance x along the flow line and h _R is bedrock height above (positive h_R ) or depth below (negative h_R present-day sea level. If h* is the isostatically lowered surface, is the isostatically lowered bed and h _M is h _R at the LGM ice margin where isostatic adjustments are small enough to ignore (Reference Hughes,, Denton, and Hughes,Hughes, 1981):

(11)

(12)

Since r varies much more slowly than Δh over distance Δx, Equation (1) can be written as the following initial-value finite-difference recursive formula for the glacio-isostatically lowered ice surface along a flowline:

(13)

All terms having a subscript i must be evaluated at each step i = x/Δx along the flowline. As with Equation (1), Equation (13) can give surface profiles for any combination of ice-sheet flow, ice-stream flow and ice-shelf flow along a flowline. In addition, it can give these profiles for any history of glacio-isostatic adjustments beneath the flowline. Solutions of Equation (13) begin at the grounding line, where i = 0 and h ₀ = (ρ _W/ρ _I)h _W in order to satisfy the buoyancy requirement.

Using Equation (13) to produce flowline profiles that match the observation-based profiles in Figure. 7 and Figures 9–12 requires knowing P_W /P_I , r, a, A, B, n, m, T_s , w. h_R , f, ϕ_s and ϕ_c along flowlines, where ϕ_s and ϕ_c depend on how a and w vary along flowlines and А, В, n and m are obtained from the flow and sliding laws of glacial ice. Some simplifications are possible. For example, τ_s = O for ice-sheet flow and can be ignored when w > > h ₁ for ice-stream flow. which is usually the case; however, see Reference Whillans, and van der Veen.Whillans and Van der Veen (1993). In addition, ice elevations are only weakly dependent on a and w, so τ_s = τ_c = 1 is a reasonable approximation, especially since a is not known accurately during deglaciation. This leaves P_W /P_I , r, f and h _R as the major variables to be specified along flowlines. Simultaneous solutions of the equations of heat transport and mass transport allow P_W /P_I and f to be calculated along flowlines, together with accurate determinations of A, B, n and m. This is possible, in principle, from time-dependent three-dimensional ice-sheet computer models but the modeling work has scarcely begun and the models must include the physics of wet deforming sediments beneath the ice (e.g. Reference Lingle,, Brown,, Van der Veen,, Oerlemans, and ReidelLingle and Brown, 1987; Reference MacAyeal,MacAyeal, 1989). Ideally, r should be obtained from a physically based global glacio-isostatic model (e.g. Reference Clark, and Lingle,Clark and Lingle, 1979; Reference Lingle, and Clark,Lingle and Clark, 1985) instead of empirically based Equations (9) and (10). Only h_R can be readily obtained from geophysical profiling but even this is lacking along the entire lengths of flowlines in Figure. 7 and Figures 9–12.

Can matching observation-based low flowline profiles in Figure. 7 and Figures 9–12 to theoretical profiles using Equation (13) decide between the two possibilities that we have postulated? The low profiles allow either ice that thickens slowly, because it is in nearly Stagnant ice-sheet flow over a frozen bed, or ice that cannot thicken at all, because it is in rapid ice-stream flow over water-soaked sediments. Low flowline profiles for nearly stagnant ice-sheet flow over a frozen bed are obtained from Equation (13) when P _W/P _I = 0, f = 0 and (τ_o)_c is small. Equation (6) gives (τ_o) and a in Equation (6) is an effective ice-accumulation rate given by the actual ice-accumulation rate ȧ minus the ice-thickening rate Ice is stagnant when the two rates are equal, so that Therefore, the more stagnant the ice, the lower (τ_o)_c becomes in Equations (6) and (13). Low flowline profiles for fast ice-stream flow over water-soaked sediments are obtained from Equation (13) when P _W/P _I → 1, f = 1, τ_s is small and w is large. A fast wide ice stream will have a small τ_s and a large w because thermal and strain softening become concentrated in lateral shear zones under these conditions. This is the case for the floating part of Byrd Glacier today (Reference Hughes,Hughes, 1977). Since the two possibilities are end members in the spectrum of ice-sheet dynamics, it is not unreasonable to expect ice-sheet models to distinguish between these possibilities, despite the formidable problems that the models must address.

Conclusions

The reconstructions presented here represent a first attempt to understand interactions between East Antarctic and West Antarctic ice during the complex grounding-line and calving-margin retreat that characterized deglaciation from the LGM in the Ross Sea. Sea-level rise certainly played a major role in this retreat. Although many of our marginal and grounding-line positions are poorly constrained, we believe that our approach, combining glaciological considerations with glacial geologic investigations of coastal and ice-shelf deposits, has the potential to yield additional insights into the interactions between these ice sheets. Our reconstructions are subject to refinement as additional constraining geologic data become available. An example is the uncertainty regarding the LGM grounding-line position in the northwestern Ross Sea. We suspect that future work will show the LGM position was near the shelf break rather than near Coulman Island. If we are correct, rapid grounding-line retreat probably occurred in this sector when sea level began to rise after 22 000 BP. This is compatible with the inherent instability of marine-based ice-sheet grounding lines much like modern fluctuations along the Siple Coast. Data from coastal deposits between McMurdo Sound and Terra Nova Bay should provide important information to better constrain reconstructions for the period from the LGM tei 8000 BP. The complete lack of data south of Minna Bluff currently hinders improvements in the grounding-line positions after about 8000 BP but no studies of this area have been made. In particular, much more sea-floor bathymetric data in MeMurdo Sound and beneath the MIS and Ross Ice Shelf are needed before a deglaciation model can be properly constrained. Also, more velocity measurements are needed on the MIS. These data could be obtained from Landsat images or aerial photographs, using the method of Reference Lucchitta,, Ferguson,, Schafer,, Ferrigno, and Williams,Lucchitta and others (1989) and Reference MacDonald,, Ferrigno,, Williams, and Lucchitta,MacDonald and others (1989). The WAIS initiative should be expanded to include these sorts of studies, if we are to ever realize the WAIS goals of fully understanding the Ross Sea deglaciation history.

Our analysis shows that glacial erratics derived from the Transantarctic Mountains and deposited on the MIS, and ice with high negative δ¹⁸Ό values imbedded in the MIS, cannot be explained by the modern ice-flow regime. They almost certainly record a former ice-flow regime when ice derived from Mulock Glacier was advected into southern McMurdo Sound. The locations where these samples were found, combined with our reconstructions, suggest that emplacement occurred during the LGM or perhaps during deglaciation but prior to about 4000 BP, by which time the modern ice-flow regime was established.

Our data allow us to evaluate two competing models for ice flow and extent during the LGM. The Reference Drewry,Drewry (1979) model fails to explain the boulder provenance and stable-isotope data, because ice from Skelton Glacier, rather than Mulock Glacier, would have dominated in southern MuMurdo Sound. We prefer the Reference Stuvier,, Yang,, Denton,, Kellogg, and McGinnis,Stuiver and others’ (198la) model, because it is easily explained by our glaciologic reconstruction, and a reasonable balance between preliminary ablation and precipitation data for the Skelton and Mulock Glaciers catchment areas, flow bands and lobes and, because the flow pattern appears compatible with stable-isotope and boulder-provenance data from the MIS. Additional supporting evidence is provided by a recent study of glacial and Holocene sediments and land forms al Cape Bird. Reference Dochat,Dochat (1994) concluded that ice flowed west and south around Cape Bird into McMurdo Sound during the LGM because no kenyite erratics were found at Cape Bird. The only known Antarctic source for kenyite is at Cape Royds on the west coast of Ross Island. On the mainland, kenyite was not deposited north of Marble Point, so flow into McMurdo Sound from a saddle on the ice divide north of Ross Island and flow around Ross Island met at Marble Point. If LGM ice flow followed the Drewry model, kenyite should be present in Ross Sea Drift at Cape Bird and no kenyite should exist south of Marble Point.

While we reject the Drewry model for the LGM. our data show that the northward through-flow of ice predicted by his model did prevail in McMurdo Sound during deglaciation. from about 12000 to 8000 BP. In fact, our glacial reconstruction for 12 000 BP for the area north and east of Ross Island resembles Reference Drewry,Drewry’s (1979) reconstruction. Hence, we suggest that his model is excellent for the middle stages of deglaciation, if the contribution of ice from Skelton Glacier to the area north of Minna Bluff is omitted. The low flowline profiles in Figure. 7 and Figures 9–12 can be explained as resulting from ice that was either largely stagnant or became highly dynamic after grounding, depending on whether the bed became frozen or remained thawed. Deciding between these two possibilities is a difficult but not impossible problem in ice-sheet modeling.