Biogenic facies in the Antarctic glacimarine environment: Basis for a polar glacimarine summary

Palaeogeography, Palaeoclimatology, Palaeoecology, 63 (1988): 357-372

357

Elsevier Science Publishers B.V., Amsterdam - - Printed in The Netherlands

BIOGENIC FACIES IN THE ANTARCTIC GLACIMARINE ENVIRONMENT: BASIS FOR A POLAR GLACIMARINE SUMMARY E. W. D O M A C K

Geology Department, Hamilton College, Clinton, NY 13323 (U.S.A.) (Received October 31, 1986; revised and accepted June 5, 1987)

Abstract Domack, E. W., 1988. Biogenic facies in the Antarctic glacimarine environment: basis for a polar glacimarine summary. Palaeogeogr., Palaeoclimatol., Palaeoecol., 63: 357-372. The association of glacial deposits (diamicts or tillites) with biogenic and/or chemically derived sedimentary rocks in the ancient record is one which continues to challenge interpretation. An analogous association can be found on portions of the East Antarctic continental shelf, where relict glacial and glacimarine diamictons are overlain by a varied and complex suite of modern biogenic deposits. Carbonate bioclastic material can be found in two distinct zones and its deposition may have been contemporaneous with partial glacial expansion onto the shelf. A nearshore carbonate facies consists of in situ and partially reworked barnacles, bryozoans, ostracods, pelycepods, gastropods, planktonic and benthonic foraminifera. A siliceous biogenic component is intermixed within the shallow water ( < 250 m) carbonates and may locally predominate. Calcareous deposits are also found as graded sands and gravel lags on portions of shelf banks and adjacent continental slopes. A deep water (> 500 m) siliceous facies, of pelagic, hemipelagic and sediment gravity flow (?) origin, is found within inner shelf depressions and consists predominantly of diatom frustules and radiolaria. These accumulations of siliceous ooze and mud clearly post date glacial recession from the shelf and contain laminated, cross-laminated, and massive interbeds which are accumulating at rates of 0.3 cm/yr. The extreme polar climate inhibits meltwater production and terrigenous influx into marine basins. This allows for the accumulation of biogenic rich material which is also favored by the high seasonal productivity inherent in the south polar oceans. Ancient analogs to these modern deposits exist and their association with glacimarine strata could be used to infer polar climatic settings and glacial retreat that is related to eustatic sea level, rather than regional warming.

Introduction The earth's glacial mode during the Latest Cenozoic has expressed itself in two major ways; through a long term glaciation of Antarctica and through short term cycles of northern hemisphere and mid-latitude glaciation. The former was primarily due to tectonic isolation over the pole and resultant oceanog r a p h i c c o o l i n g ( K e n n e t t e t al., 1974). T h i s event produced major and long-lasting changes 0031-0182/88/$03.50

in global ocean circulation and sedimentation ( K e n n e t t , 1977; I n g l e , 1981). T h e l a t e r c y c l e s a r e b e l i e v e d t o be c o n t r o l l e d b y f l u c t u a t i o n s i n t h e earth's axis and orbit and hence are responses t o r e g i o n a l w a r m i n g a n d c o o l i n g ( H a y s e t al., 1976; I m b r i e e t al., 1985). B e c a u s e o f t h e p o l a r glacial setting of the antarctic, peripheral fluctuations of glacial ice along the continental shelf are tied to the relationship between eustatic sea level and ice-shelf and/or ice tongue d y n a m i c s ( H o l l i n , 1962), I n t h i s i n s t a n c e g l a c i a l

© 1988 Elsevier Science Publishers B.V.

358 recession, though partial, is not so much a climatic response to regional warming but a response to rising sea level. Therefore, a primary contrast between Antarctic ice sheet recession and recession of the northern hemisphere ice sheets, during the late Pleistocene, is one of limited ablation and meltwater sedimentation for the former. To further our understanding, we need to evaluate ancient (pre-Cenozoic) glaciations in the same light. The first step would be to make a climatic distinction between temperate, subpolar, and polar glacial sequences; a problem which can be approached from the perspective of facies analysis. Ice-recessional marine deposits best reflect the variable role of meltwater processes, and thus climate, within each of these settings (Domack, 1984). For example, rapidly deposited sands and muds are the dominant facies in both ice-proximal and ice-distal positions in temperate glacimarine environments, where mean summer temperatures are significantly above 0°C and sea ice is lacking (Powell, 1981, 1983; Molnia and Hein, 1982; Molnia, 1983). Subpolar environments lie within seasonal sea-ice limits and have mean summer temperatures _>0°C. Meltwater sands and muds, therefore, occur primarily in ice-proximal positions (Gilbert, 1982; Pfirman, 1985). Polar glacial marine environments, such as those surrounding Antarctica, are characterized by extensive winter sea ice and mean summer temperatures < 0°C. Modern sedimentation under such a regime is dominantly biogenic. The purpose of this paper is to document the nature of biogenic sediments on a portion of the Antarctic continental shelf and to discuss the role of eustacy, productivity, sediment gravity flow, sea ice and limited terrigenous supply on the development of facies patterns. These biogenic deposits provide a fundamental contrast to glacial retreat sequences developed under more moderate climates.

Background Samples described in this study were retrieved during Operation Deep Freeze 1979

aboard the USCGC Glacier (Anderson et al., 1979a). The area of study includes the part of the East Antarctic continental shelf between 140 ° and 150 ° E Longitude (Fig.l). H e r e , a portion o f the world's largest ice-sheet terminates along the coast in the form of grounded ice cliffs (Ad~lie coast) and floating ice tongues of the Mertz and Ninnis outlet glaciers. The climatic conditions of the area are polar but not extreme for the Antarctic. The average surface temperature is approximately - 15.0°C with mean summer temperatures of - 2.0°C. As a consequence little surface meltwater is generated and extensive areas of sea ice form, melt back and disperse seasonally. Open water forms persistently during the austral summer only in the area west of the Mertz Glacier, while further east ice bound conditions persist for most of the year. Oceanographic and bathymetric data are discussed further by Domack (1980, 1982), Domack and Anderson (1983), Dunbar et al. (1985) and Chase et al. (1987). Relict glacial and glacimarine diamictons are the predominant subsurface deposit on the shelf and their origins are discussed in detail in Anderson et al. (1980a), Domack (1982) and Hampton (1987). Recent chronologic determinations (14C dating) suggest that glacial ice may have retreated from the shelf sometime between the latest Pleistocene and middle Holocene (Domack et al., 1988). Holocene sediments of varied composition and thickness overlie the glacial diamicton facies and are the subject of this paper.

Carbonate bioclastic facies Nearshore sediments

Carbonate bioclastic material can be found in two distinct zones on the continental shelf (Fig.2). A nearshore region, boardering the ice sheet, contains in situ and reworked accumulations of calcareous shell material. These sediments are found above 250 m, to depths as shallow as 10 m along the Ad~lie Land coast (Chapman, 1922; Mawson, 1940). Shell debris so

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far recognized from samples 8 and 9 consists of barnacles, bryozoans, ostracods, pelycepods, gastropods, foraminifera, and calcareous algae. Such carbonate particles dominate (> 80%) the sand and gravel fraction of these poorly sorted sediments (Figs.3 and 4). The finer fraction consists of carbonate, terrigenous silt (quartz), clay and siliceous material composed of sponge spicules with lesser abundant diatoms and radiolarians. The total organic carbon (TOC) content for some of these samples is high, at 1.9%, considering their coarse nature and is evidence of high percents of organic matter in the finer fraction (> 6~/o or more, Fig.4). Locally such sediments may be as old as mid-Holocene since a 14C age of 5019_+ 179 has been obtained on foraminifera from sample 9 (A. J. T. Jull, Univ. Arizona, pers. comm., 1986).

Shelf bank and slope sediments Bioclastic rich sediments also occur on the upper continental slope and in areas of the shelf adjacent to shallow banks (Figs.2 and 5). These sand and gravel units are thin (0.2-2.0 m) and are characterized by unlithified sedimentary clasts, of mud or sand, and well developed normal grading (Domack, 1980, 1982; Wright et al., 1983). The biogenic (carbonate) content varies from approximately 10% to as

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T. Jull, pers. comm., 1986). Therefore, bioclastic sand deposition may be contemporaneous with at least partial glaciation of the shelf (Domack, 1982; Domack et al., 1988). Biogenic (carbonate) rich surface samples were also recovered adjacent to shallow knolls on the outer continental shelf. This bryozoan rich sand and gravel appears to represent in situ accumulations or lags according to grain size data (Domack, 1980; Dunbar et al., 1985). However, since piston cores (subsurface samples) were not recovered in these locations the role of downslope transport of these sediments cannot be fully assessed. A 14C date of 3230_ 200 yr B.P. from surface sample 25 suggests relatively recent carbonate deposition in this area (A. J. T. Jull, pers. comm., 1986).

Below approximately 500m in depth fine grained siliceous-rich sediments are accumulating (Figs.1 and 2) in a zone which extends up to the floating terminus of the Mertz Glacier. Bottom profiles, using a 12 kHz source, reveal a rather uniform (6 m thick) and draped layer of this sediment within most shelf basins and a distinct thickening (up to 30 m) in the western end of the Mertz-Ninnis Trough (Domack and Anderson, 1983). Thicknesses of a few cm to several meters have been recovered in piston and trigger (gravity) cores which also reveal sharp contacts of the siliceous facies with underlying diamictons (see Fig.7; Domack, 1980, 1982; Hampton, 1986). Thus, deposition of this unit postdates a period of more widespread glaciation on the shelf. Texturally these sediments range from sandy muds to muds while compositionally they range from diatom and radiolarian oozes to siliceous muds (Fig.6). The dominant biogenic constituents consist of planktonic diatoms, minor benthic diatoms,-radiolarians, silicoflagellates and sponge spicules. Calcium carbonate is almost completely lacking within these deposits. Fecal pellets comprise up to 25°//0 of the sediment, where they have yet to disintegrate in shallow portions of cores, and terrigenous content averages about 50%, as determined by microscopic examination of the sand and silt fraction (Fig.6). The biogenic content is noticeably greater within laminated intervals (Figs.6 and 7). TOC contents are generally greater than 1.0% (i.e., 1.17% and 1.20% for cores 12-266 cm and 13-470 cm). Sedimentary structures, as determined from X-ray radiographs, are varied and include horizontal laminations (approximately 0.25-O.5cm thick), cross-laminations, convoluted laminations, and massive units (Figs.7 and 8). Scour structures (erosional surfaces) are also present. Annual deposition of up to 0.5 cm (the average laminae pair thickness) may be occurring as 21°pb measurements for

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cores 12 and 13 indicate an accumulation rate of 0.3 cm/yr (Domack and Anderson, 1983). Poorly sorted sand and gravel (ice-rafted?, debris, IRD) is rare within the thick ooze deposits but is significantly greater in core 13 than core 12 (Figs.7 and 8). Discussion

Terrigenous supply and productivity A number of factors control the distribution and depositional processes f o r the biogenic sediments described above (Fig.9). Of these, shelf physiography, high productivity and a lack of meltwater derived terrigenous sediment probably are the most important, but gravity flow processes, water mass circulation (temperature) and sea ice also play an important role. It is known from several studies that high turbidity is an important environmental stress for carbonate secreting organisms. Such turbidity would be even more limiting for biogenic carbonate production in frigid (polar) waters

where growth rates are slow and carbonate solubility is high. Thus, the bioclastic carbonate facies found just offshore of the Ad~lie land coast would suggest limited turbidity in this ice proximal setting. Though data on suspended particulates along the East Antarctic ice sheet are lacking it would seem intuitive that an absence of a normal coastal zone and an interior which is covered by a subfreezing ice sheet would severely restrict terrigenous supply to the continental shelf. Exceptions would include ice rafting and ice marginal dumping, along the base of ice cliffs (ZnachkoYavorskiy, 1979; Anderson et al., 1983). In the area surveyed during Deep Freeze-79 surface turbidity was limited to diatom blooms and other plankton. Hence productivity must also play a role in shelf sedimentation, particularly in regards to the siliceous-rich basinal facies. Yet the limited dilution of the biogenic component, by terrigenous mud in these sediments, again emphasizes the apparent absence of significant meltwater derived sedimentation. Any productivity related model for the

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origin of the siliceous facies must also explain the absence of calcium carbonate, varied sedimentary structures, moderate T.O.C. and the moderate but persistent terrigenous content of the siliceous basinal facies. Though the basinal siliceous facies is predominantly pelagic in origin other processes may be important. The lack of pelagic carbonate within these deposits, particularly plank-

6 -Fig.8. Sedimentary structures as revealed by x-ray radiographs for core 13. Note abundance of ice rafted gravel (large dots) and sand (small dots).

tonic foraminifera, is most likely related to dissolution at depths greater than 500 m on the continental shelf (Kennett, 1966). The lack of calcium carbonate might also be explained as a result of high rates of carbonic acid production related to aerobic respiration (Aller, 1982). Yet this process should also be occurring within the shallow water sediments which clearly have significant amounts of preserved organic carbon and calcium carbonate, even within the mud fraction (Fig.5). Also, the presence of moderate TOC contents and well preserved laminations in the deep-water oozes would

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suggest an absence of infaunal organisms and anoxic bottom waters within these silled basins. However, there is no direct evidence for this as yet. In fact, Gordon and Tchernia (1972) have determined that the basinal waters are near oxygen saturation, based upon their measurements of slope water believed to be derived from the basin. Instead the origin of the horizontal laminations is likely related to rapid accumulation (high productivity) dominating over bioturbation. Preliminary results demonstrate that the laminations represent a rapid biogenic event since each laminae pair, in contrast to massive portions of the cores, are essentially lacking in terrigenous material (Figs.5-7). Laminae and organic carbon preservation within these sediments may therefore be due to rapid fluxes to the bottom, perhaps over a period of weeks or months, along with the very low bottom water temperatures ( < - 1 . 5 ° C ) which would limit bacterial metabolism. Sources for the terrigenous component could include sediment gravity flow, meltwater, aeolian transport, ice-rafting and bottom current erosion and transport.

Sediment gravity flows The presence of normally graded, bioclasticrich, sands and gravels attest to the importance of gravity flow mechanisms to the dispersal of coarse grained sediments on the

shelf. Such units are interpreted as turbidites. The generative processes for such currents are believed to be related to post-glacial readjustment (slumping and debris flow) of local relief features on the shelf (Domack, 1982; Wright et al., 1983). It is suggested here that storm related currents (surges) may also have been responsible for graded sand deposition on the shelf. In any event, the responsible currents were apparently of sufficient energy to prevent the deposition of upper divisions of the Bouma sequence (B-D) since nowhere do sands grade texturally into silts. Rather, overlying the graded units are siliceous muds which are hydrodynamically in disequilibrium with underlying sands and gravels (Domack, 1980). The fate of suspended fines produced by the turbidity currents may lie on the continental s].ope/rise and within deeper and lower relief portions of the shelf. Evidence of sediment gravity flow deposition on the continental slope and rise is abundant and includes graded sands and gravels and massive (debris flow) diamictons (Figs.1 and 5; Anderson et al., 1981). However, the distribution of biogenic detritus within such units is limited to regions of the slope directly adjacent to shelf banks (Fig.l), thus suggesting a source related to biogenic accumulation on the banks. In contrast, slope areas adjacent to major pathways for glacial drainage (i.e., George V Basin; Fig.l) contain clastic dominated lithologies including laminated muds (Anderson et al., 1981).

365 Massive units within the basinal siliceousrich facies (Figs.7 and 8) may represent sediment gravity flows similar to massive unifites recognized by other authors (Stanley, 1981; and others). However, if the massive siliceous units are the distal equivalents of bioclastic sands then they should contain some calcium carbonate mud. The deseminated presence of granules and coarse sand within massive intervals (Figs.7 and 8) would also argue against fully turbulent processes of transport, as coarse grains should be concentrated near the base of units. Coarse grains could be supported by the matrix strength of dense debris-type flows, but the resulting deposit should be distinct in composition, carbonate and water content from pelagic portions of the core. No such contrasts were recognized within freshly cut cores or samples. Thus given the present data, the role of fine grained sediment gravity flows within the basin remains problematic. The thickest part of the basin infill (Fig.9) has yet to be cored. Logically it should contain a greater proportion of gravity flow muds than the uniform and draped layers recovered in cores 12 and 13 (Fig.9). Perhaps the process of fine grained sediment gravity flow was more important during the early stages of basin infill, under lower sea level, when shelf turbidity currents reasonably would have been more frequent. A 14C age of 14,261 ___136 for core 4 along w i t h the fact that some coarse turbidites are overlain by siliceous sediment supports the view of more active gravity flow transport earlier in the Holocene and perhaps late Pleistocene. Water mass circulation

On the outer portion of the continental shelf, impinging Circum Polar Deep Water (CDW) has effectively sorted relict diamicton, into coarse to medium sands and gravels, and has redistributed modern biogenic detritus (Domack, 1980; Domack and Anderson, 1983). This process is aided significantly by bioturbation of the substrate (Singer and Anderson, 1984). Fine grained sediment~ removed in this manner

may end up within the basinal siliceous facies as it is transported shelfward along the isopycnal (0 ° isotherm in Fig.9). Sediment budget studies (Dunbar et al., 1985) have determined that the terrigenous content within the siliceous sediment could have been derived entirely from this source. Since ice-rafting is occurring (Fig.8; Anderson et al., 1980b) it seems the role of meltwater as a sediment source is further diminished. Based upon preliminary study, the distribution o f ice-rafted material also seems to be related to water mass distribution. Surface and subsurface water temperatures on the majority of the shelf are between - 1 . 0 and -1.9°C (S. Jacobs, pers. comm., 1979; data in Domack, 1980). Such temperatures severely limit the melt rate of icebergs (Morgan and Budd, 1978) and hence rates of debris release. Only where CDW intrudes onto the outer shelf and upper slope are temperatures above 0.0°C encountered (Fig.9). These intruding warm tongues are dynamic features which seem to fluctuate in position from year to year. However, the position of core 13 relative to a warm tongue observed in 1979 (Domack, 1980) would seem to explain the high percentage of IRD in this core (Fig.8). Sea ice

Sea ice plays a major role in controlling surface productivity (Hart, 1934); as recently demonstrated in other regions of the Antarctic including the Ross Sea and Weddell Sea (E1Sayed and Taguchi, 1981; Nelson et al., 1984; Smith and Nelson, 1985). Low salinity surface waters are derived by the melting back of sea ice during the austral summer. According to Smith and Nelson (1985) this sea ice meltwater creates environmental stability within the photic zone which is quickly manifested by dense diatom blooms. High productivity is also stimulated by seeding of epontic (in ice) diatoms released into the sea. Areas which undergo seasonal fluctuations in sea ice c o v e r may therefore also contribute significant quantities of biogenic detritus (silica and organic matter) to the bottom. This mechanism may be

366 important in controlling the siliceous facies distribution as well as influencing the carbonate facies by the advection of planktonic productivity to the benthos, as demonstrated in McMurdo Sound (Dayton and Oliver, 1977). However, it is not known whether the seasonal conditions of open water in the western area off the George V coast (i.e., where the siliceous ooze is thickest) are due to wind forced breakout of sea ice, and hence thorough mixing of the surface waters, or melting. The graded sands found in this area do however contain a lot of biogenic carbonate while graded sands found to the east, where sea ice is present year round, lack a biogenic fraction. Therefore, source areas for such sediments, vary in their content of organisms and thus biogenic accumulation. If the productivity model of Smith and Nelson (1985) is applicable to the open water area off the George V coast, then it may also explain the development of laminated intervals within the siliceous-rich facies. Initial analysis of diatom floras within individual lamina reveal strong contrasts in species diversity and size. Laminae differ in that small epontic diatoms (i.e., Nitzchia kerguelensis) and resting spores(?) dominate within one of each laminae pair. Such layers may represent the initial pelagic flux induced by melting sea ice, while overlying(?) laminae represent productivity during the rest of the austral summer. This concurs with the results of Wilson and Smith (1985) who found similar contrasts in diatom assemblages between high-productivity (sea ice marginal) zones and open antarctic waters. Bodungen et al. (1986) have also concluded that such dense (short lived) diatom blooms along the northern Antarctic Peninsula are quickly fluxed to the bottom. The massive portions of cores 12 and 13 may represent hemipelagic sedimentation during periods, several years (?), when seasonal contrasts in productivity are not reflected within the sediments. This could be the result of locally more extensive sea ice or, alternately, the persistence of open water or polynas. Rates of biogenic (pelagic) sedimentation would be lower under such conditions thus allowing for

the increased terrigenous content and effective bioturbation. The seasonal growth and scour effect of sea ice also controls the extent of the Shallow water carbonate facies. This process has been reported to restrict biogenic accumulation to depths of > 15 m (Gruzov, 1977; Picken, 1985). Ice-rafting by sea ice within the nearshore zone is also important (Mawson, 1940; Anderson et al., 1980b).

Comparison to the rest of the Antarctic The biogenic facies model presented above is generally applicable to the rest of the Antarctic, exclusive of the Antarctic Peninsula. In the Ross Sea, basinal areas are also accumulating siliceous rich facies though the terrigenous content is greater and bottom currents play a greater role in biogenic silica distribution (Myers, 1982; Anderson et al., 1983; Dunbar et al., 1985). Shallow banks have calcareous debris and siliceous (sponge) spicule mats (Elverhoi and Roaldset, 1983) as well as thin turbidites (Wright et al., 1983). Terrigenous supply is also greater along the mountainous coast of Victoria Land due to eolian transport by katabatic winds, which move across exposed bedrock, and local volcanic sources (Brake, 1983). Pre-Holocene biogenic sediments (siliceous oozes) have also been preserved along the East Antarctic coast despite subsequent overriding by glacial ice (Adamson and Pickard, 1983). The general shelf physiography, glacial and sea ice conditions of the George V-Ad~lie margin are quite similar to the rest of the East Antarctic shelf, which comprises about 60~/o of the antarctic margin. Therefore, despite the fact that the vast majority of the shelf remains unsampled the facies relationships discussed above are expected to be representative but must be considered as a regional summary at this time.

Other regions The continental shelves of Labrador, British Columbia and the Barents Sea contain exten-

367 sive accumulations of biogenic, calcareous, debris which overlie relict glacial diamictons (Muller and Milliman, 1973; Nelson and Bornhold, 1983; Elverhoi and Solheim, 1983). Coastal Labrador and British Columbia cannot today be called glacimarine as glaciers no longer reach the sea. Therefore, bioclastic carbonates in these settings represent true interglacial deposits (Elverhoi and Roaldset, 1983). The Barents Sea shelf has certain similarities to, yet important differences from, that of the Antarctic. The region is polar in latitude (70~0°N) yet subpolar in climate. Thus, icecap and valley glacier systems of Nordaustlandet and Spitsbergen contribute significant quantities of terrigenous sediment to the marine environment, though meltwater muds are deposited relatively close (< 15 km) to the ice front (Pfirman, 1985). Like the Antarctic, seasonal sea ice fluctuations within the Barents Sea control phytoplankton productivity (Rey and Leong, 1985) which undoubtedly contributes to fine grain deposition. However, productivity is noticeably restricted by the limited supply of nutrients within the water column (Rey and Leong, 1985); in contrast nutrient supply is not a limiting factor in Antarctic waters (Nelson et al., 1984; Smith and Nelson, 1985). The above conditions in the Barents Sea limit bioclastic carbonates to very distal settings with respect to the coast ( > 50 km) and result in basinal (> 200 m deep) muds and sandy muds which are dominated by terrigenous material (Elverhoi and Solheim, 1983). These regions therefore lack the facies association of ice-proximal and shallow water carbonates, bioclastic turbidites, and basinal siliceous sediments found in the Antarctic. Thus it seems justified in proposing a polar glacimarine model (Fig.9) which has as its main components a continent wide ice sheet, average summer temperatures of < 0°C, and extensive yet dynamic sea ice cover. Since continental wide ice sheets in polar settings have existed a number of times through earth history (Hambrey and Harland, 1981) it seems as though some sedimentologic record similar to the present day Antarctic should exist.

Early Permian of Tasmania Early Permian (Sakmarian) sedimentary rocks in Tasmania possess abundant evidence for glacimarine sedimentation in the form of striated pavements, tillite and dropstone bearing stratified sediments (Banks, 1962; Crowell and Frakes, 1971; Fig.10). The Wynyard Tillite, and subsequent units, were deposited over a highly irregular sub-Permian surface with a relief of several hundred meters (Banks, 1962; Clarke and Banks, 1975). This paleotopography was marked by a linear deep, or basin, and an outer shelf high toward the present northeast. Of particular interest in this sequence is the presence of biogenic sedimentary rocks above the Wynyard diamictite (tillite). In basinal positions the biogenic facies consist of Tasmanite, a unique unit of the Quamby Mudstone Group (Figs.10 and 11). The Tasmanite unit consists of at least two beds, several meters thick, which can be traced laterally over several hundred km 2. Internal structures include fine scale horizontal laminations, cross-laminations and scour surfaces (M. Banks, pers. comm., 1985). It is essentially an oil shale with TOC (kerogen) contents in excess of 8% (Reid, 1924). It is almost lacking in terrigenous material, except for scattered ice-rafted clasts and thin silt laminae, and instead is composed of the organic-walled microfossil Tasmanites. Tasmanites was not correctly assigned taxonomically until 1962 when Wall established their relationship to living marine forms of blue-green algal cysts; later described by Parke (1966). Therefore, Tasmanites likely occupied a position within the phytoplankton similar to modern diatoms. However, the origin of the Tasmanite itself is still unknown. Carey and Ahmad (1961) suggested that the Tasmanite "organisms" flourished in basinal, hypersaline, environments created by the seasonal freezing of sea ice; thus their analogy with the Antarctic. Williams (1978) and Combaz (1967) however, proposed that melting glacial ice provided low salinity marine settings where the Tasmanites could also have flourished. Yet modern glacial

368

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369 marine settings where glacial ice is melting, to the point of diluting sea water (i.e., Gulf of Alaska), are dominated by terrigenous sedimentation. It is suggested here that the sea ice melting and plankton blooms described from the Antarctic are more appropriate to the origin of the Tasmanite (Fig.ll). Bioclastic carbonates and interbedded, thin, conglomerates are also found within the sequence, though they are believed to be slightly younger stratigraphically (Fig.10). These units are described as containing ice-rafted clasts and an abundant but low diversity fauna dominated by Eurydesma shells (pelecypods), brachiopods, and bryozoans (Rao, 1981). The carbonates comprise irregular beds 1-2 m thick and are restricted in their distribution to paleogeographic highs in the northeast (Fig.10). Thin bryozoan rich turbidites are also present and were apparently deposited by density currents flowing off the highlands into the basin (Banks, 1962; Rao, 1981). This biogenic facies association is almost identical, except for specific species content, to that described previously for the Antarctic. Further, the paleomagnetic data interpreted by Embleton (1973) and Irving (1966) place Tasmania at approximately 80°S Latitude during the Early Permian. Thus there is every reason to suggest a true polar glacimarine setting, replete with sea ice, for the Early Permian of Tasmania.

Concluding remarks The application of the polar glacimarine model presented here can aid in addressing the following specific questions. First, detailed sedimentologic studies on the Tasmanite could resolve the problems facing the interpretation of the antarctic siliceous oozes, namely the relative role of pelagic vs sediment gravity flow deposition. Second, physical oceanographic conditions in a high latitude sea of the Early Permian could be addressed, in as much as modern antarctic sediments strongly reflect conditions of productivity, sea ice extent and water mass circulation.

Further, debate continues on the paleoclimatic significance of ancient glacial and glacial marine diamictites which are associated with organic rich (biogenic) and/or chemically derived sedimentary rocks (Walter and Bauld, 1983). Interpretations of various sequences include: rapid climatic warming following deglaciation (Williams, 1979; Fairchild and Hambrey, 1984; Tucker, 1986), complete deglaciation (via ablation) but continued cold during the biogenic interval (Deynoux et al., 1985), and transgression induced productivity events (Tucker, 1983). A comparison to the modern antarctic suggests that similar sequences could be interpreted as true polar glacial marine deposits. This would imply not only a frigid, polar, setting but also the ice recessional pattern indicative of an ice shelf/ice tongue response to a eustatic rise in sea level. Also, those oceanographic characteristics of the southern ocean, such as sea ice fluctuations and bottom water production, could be inferred during these ancient glacial episodes. Thus, an understanding of the true climatic significance of these rocks would substantially improve our knowledge of polar and glacial climates through earth history.

Acknowledgements This work was originally supported by National Science Foundation grant DPP-77-26407 and DPP-80-80242 to Dr. John Anderson (Rice University). Gratitude is extended to D. Cassidy at the Florida State University Antarctic Core Facility for his help in providing samples, S. Jacobs for providing oceanographic data and M. Banks for his review of the manuscript. Carbon-14 dates were obtained from the University of Arizona tandem accelerator-mass spectrometer; analyses were partially supported by NSF grant DPP-86-13565 and a grant from Hamilton College. This paper was presented in abstract form during the glacigenic deposystem symposium at the GSA northcentral meeting in DeKalb, Illinois, 1985. Portions of this paper were excerpted from earlier, unpublished, material.

370

References Adamson, D. and Pickard, J., 1983. Late Quaternary ice movement across the Vestfold Hills, East Antarctica. In: R.L. Oliver, P.R. James and J.B. Jago (Editors), Antarctic Earth Science. Aust. Acad. Sci., Canberra, pp. 465-469. Aller, R. C., 1982. Carbonate dissolution in nearshore terrigenous muds, the role of physical and biological reworking. J. Geol., 90: 79-95. Anderson, J. B., Balshaw, K., Domack, E., Kurtz, D., Milam, R. and Wright, R., 1979. Geological survey of east antarctic margin aboard USCGS Glacier. Antarct. J. U.S., 14: 142-144. Anderson, J. B., Kurtz, D. D., Domack, E .W. and Balshaw, K. M., 1980a. Glacial and glacial marine sediments of the Antarctic continental shelf. J. Geol., 88: 399-414. Anderson, J. B., Domack, E.W. and Kurtz, D.D., 1980b. Observations of sediment laden icebergs in Antarctic waters: implications to glacial erosion and transport. J. Glaciol., 25: 387-396. Anderson, J. B., Davis, S. B., Domack, E. W., Kurtz, D. D., Balshaw, K. M. and Wright, R., 1981. Marine sediment core descriptions IWSOE 68, 69, 70: Deep Freeze 79. Rice University, Houston, Tex., 60 pp. (unpublished). Anderson, J. B., Brake, C., Domack, E. W., Myers, N. and Wright, R., 1983. Development of a polar glacial-marine sedimentation model from Antarctic Quaternary deposits and glaciological information. In: B.F. Molnia (Editor), Glacial-Marine Sedimentation. Plenum Press, New York, N.Y., pp. 233-264. Banks, M.R., 1962. Permian (Tasmania). J. Geol. Soc. Aust., 9: 189-216. Bodungen, B.V., Smetacek, V.S., Tilzer, M.M. and Zeitzschel, B., 1986. Primary production and sedimentation during spring in the Antarctic Peninsula region. Deep-Sea Res., 33: 177-194. Brake, C.F., 1983. Sedimentology of the North Victoria Land continental margin. Thesis. Rice University, Houston, Tex., 175 pp. Carey, S. W. and Ahmad, N., 1961. Glacial marine sedimentation: In: G. Raasch (Editor), Proc. First Int. Symp. on Arctic Geology, 2. Univ. Toronto Press, pp. 865-894. Chapman, E., 1922. Sea-floor deposits. In: D. Mawson (Editor), Reports Australian Antarctic Expedition, 1911-1914, Ser. A:2(1). Chase, T. E., Seekins, B. A. and Young, J. D., 1987. Marine topography of offshore Antarctica. In: S. L. Eittreim and M.A. Hampton (Editors), The Antarctic Continental Margin: Geology and Geophysics of Offshore Wilkes Land. Circum-Pac. Counc. Energy Miner. Resour., Houston, Tex., pp. 147-150. Clarke, M.J., 1968. Reappraisal of the Lower Permian Type section, Golden Valley, Tasmania. Rec. Geol. Surv. Tasmania, 8: 7-32. Clarke, M. J. and Banks, M. R., 1975. The stratigraphy of the lower (Permo-Carboniferous) parts of the Parmeener super-group, Tasmania. In: K. S. W. Cambell (Editor), Gondwana Geology. Aust. Natl. Univ. Press, Canberra, pp. 453-467.

Combaz, A., 1967. Leiosphaeridiacea Eisenack, 1954 et Protoleiosphaeridiaceae Timofeev, 1959--leurs aflinit~s, leur rSle s~dimentologique et g~ologique. Rev. Palaeobot. Palynol., 1: 309-321. Crowell, J.C. and Frakes, L.A., 1971. Late Paleozoic glaciation of Australia. J. Geol. Soc. Aust., 17: 115-155. Dayton, P. K. and Oliver, J. S., 1977. Antarctic soft-bottom benthos in oligotrophic and eutrophic environments. Science, 197: 55-58. Deynoux, M., Sougy, J. and Trompette, R., 1985. Lower Paleozoic rocks of west Africa and the western part of central Africa. In: C. H. Holland (Editor), Lower Paleozoic of North-western and West Central Africa. Wiley, New York, N.Y., pp. 337-496. Domack, E. W., 1980. Glacial marine geology of the George V-Ad~lie continental shelf, East Antarctica. Thesis. Rice Univ., Houston, Tex., 142 pp. Domack, E. W., 1982. Sedimentology of glacial and glacial marine deposits on the George V-Ad~lie continental shelf, East Antarctica. Boreas, 11: 79-97. Domack, E.W., 1984. Rhythmically bedded glaciomarine sediments on Whidbey Island, Washington. J. Sediment. Petrol., 54: 589-602. Domack, E. W. and Anderson, J. B., 1983. Marine geology of the George V continental margin: combined results of Deep Freeze 1979 and the 1911-1914 Australasian expedition: In: R. L. Oliver, P. R. James and J. B. Jago (Editors), Antarctic Earth Science. Aust. Acad. Sci., Canberra, pp. 402-406. Domack, E. W., Jull, A. J. T., Anderson, J. B. and Linick, T.W., 1988. Mid-Holocene ice-sheet recession from the Wilkes Land continental shelf, East Antarctica. In: Fifth Antarct. Earth Sci. Syrup., Cambridge, in press. Dunbar, R. B., Anderson, J. B., Domack, E. W. and Jacobs, S.S., 1985. Oceanographic influences on sedimentation along the Antarctic continental shelf. In: S.S. Jacobs (Editor), Oceanology of the Antarctic continental shelf (Antarctic Research Series). 43. Am. Geophys. Union, Washington, D.C., pp. 291-312. E1-Sayed, S. S. and Taguchi, S., 1981. Primary production and standing crop of phytoplankton along the ice-edge in the Weddell Sea. Deep-Sea Res., 28A: 1017-1032. Elverhoi, A. and Roaldset, E., 1983. Glaciomarine sediments and suspended particulate matter, Weddell Sea shelf, Antarctica. Polar Res., 1: 1-21, Elverhoi, A. and Solheim, A., 1983. The physical environment Western Barents Sea 1:1,500,000: Sheet A, surface sediment distribution. Nor. Polarinst. Skr., 179, 23 pp. Embleton, B. J.J., 1973. The paleolatitude of Australia through Phanerozoic time. J. Geol. Soc. Aust., 19: 475-482. Fairchild, I.J. and Hambrey, M.J., 1984. The Vendian succession of northeastern Spitsbergen: Petrogenesis of a dolomite-tillite association. Precambrian Res., 26: 111-167. Gilbert, R., 1982. Contemporary sedimentary environments on Baffin Island, N.W.T., Canada: Glaciomarine processes in fiords of eastern Cumberland Peninsula. Arct. Alp. Res., 14: 1-12. Gordon, A.L. and Tchernia, P., 1972. Waters of the

371 continental margin off Ad~lie coast, Antarctica. In: D. E. Hayes (Editor), Antarctic Oceaonography II: The Australian New Zealand sector (Antarct. Res. Ser., 9). Am. Geophys. Union, Washington, D.C., pp. 59-69. Gruzov, E. N., 1977. Seasonal alterations in coastal communities in the Davis Sea. In: G. A. Llano (Editor), Proc. Third SCAR Syrup. on Antarctic Biology, pp. 262-278. Hambrey, M.J. and Harland, W.B. (Editors), 1981. The Earth's Pre-Pleistocene Glacial Record. Cambridge Univ. Press, Cambridge, 1024 pp. Hampton, M. A., 1987. Geology of sediment cores from the George V continental margin, Antarctica. In: S.L. Eittreim and M.A. Hampton (Editors), The Antarctic Continental Margin: Geology and Geophysics of Offshore Wilkes Land. Circus-Pac. Counc. Energy Miner. Resour., Houston, Tex., pp. 151-174. Hart, T. J., 1934. On the phytoplankton of the South-West Atlantic and the Betlingshausen Sea, 1929-1931. Discovery Rep., 8: 1-268. Hays, J. D., Imbrie, J. and Shackleton, N.J., 1976. Variations in the Earth's orbit: pacemaker of the ice ages. Science, 194: 1121-1132. Hollin, J. T., 1962. On the glacial history of Antarctica. J. Glaciol., 4: 173-195. Imbrie, J., Shackleton, N.J., Pisias, N.G., Morley, J.J., Prell, W. L.; Martinson, D. G., Hays, J. D., McIntyre, A. and Mix, A. C., 1985. The orbital theory of Pleistocene climate: support from a revised chronology of the marine 1sO record. In: A.L. Berger et al. (Editors), Milankovitch and Climate. I. Reidel, Dordrecht, pp. 269-305. Ingle, J. C., Jr., 1981. Origin of Neogene diatomites around the North Pacific rim. In: R.E. Garrison and R.G. Douglas et al. (Editors), The Monterey Formation and Related Siliceous Rocks of California, Pac. Section Soc. Econ. Paleontol, Mineral., Los Angeles, Cal., pp. 159-179. Irving, E., 1966. Paleomagnetism of some Carboniferous rocks from New South Wales and its relation to geological events. J. Geophys. Res., 71: 6025-6051. Kennett, J. P., 1966. Foraminiferal evidence of a shallow calcium carbonate solution boundary, Ross Sea, Antarctica. Science, 153: 191-193. Kennett, J. P., 1977. Cenozoic evolution of Antarctic glaciation, the circum-Antarctic Ocean, and their impact on global paleoceanography. J. Geophys. Res., 82: 3843-3860. Kennett, J.P., Houtz, R.E., Andrews, P.B., Edwards, A. R., Gostin, V. A., Hajos, M., Hampton, M. A., Jenkins, D.G., Margolis, S.V., Ovenshine, A.T. and PerchNielsen, K., 1974. Development of the circum-Antarctic current. Science, 186: 144-147. Mawson, D., 1940. Marine biological programme and botanical activities. In: D. Mawson (Editor), Report Australasian Antarctic Expedition 1911-1914, Ser. A, 2, (5): 127-165. Molnia, B.F., 1983. Subarctic glacial-marine sedimentation--a model. In: B. F. Molnia (Editor), Glacial-Marine Sedimentation. Plenum, New York, N.Y., pp. 95-144. Molnia, B.F. and Hein, J., 1982. Clay mineralogy of a

glacially dominated subarctic continental shelf: northwestern Gulf of Alaska. J. Sediment. Petrol., 52: 515-527. Morgan, V. I. and Budd, W. F., 1978. Distribution, movement and melt rates of Antarctic icebergs. In: A.A. Husseiny (Editor), Int. Conf. on Iceberg Utilization, Iowa State Univ. Pergamon Press, New York, N.Y., pp. 220-228. Muller, J. and Milliman, J. D., 1973. Relict carbonate-rich sediments on southwestern Grand Bank, Newfoundland. Can. J. Earth Sci., 10: 1744-1750. Myers, N. C., 1982. Marine geology of the Western Ross Sea: implications for Antarctic glacial history. Thesis. Rice Univ., Houston, Tex., 234 pp. Nelson, C. S. and Bornhold, B. D., 1983. Temperate skeletal carbonate sediments on Scott Shelf, Northwestern Vancouver Island, Canada. Mar. Geol., 52: 241-266. Nelson, D.M., Gordon, L.I. and Smith, W.O., 1984. Phytoplankton dynamics of marginal ice zone of the Weddell Sea, November and December 1983. Antarct. J. U.S., 19: 105-107. Parke, M., 1966. The genus Pachysphaera (Prasinophyceae). In: H. Barnes (Editor), Some Contemporary Studies in Marine Sciences. Allen and Unwin, London, pp. 555-563. Pfirman, S. L., 1985. Modern sedimentation in the northern Barents Sea: Input, dispersal, and deposition of suspended sediments from glacial meltwater. Thesis. Woods Hole Oceanographic Institute, Woods Hole, Mass., 378 pp. Picken, G. B., 1985. Benthic research in Antarctica: Past present, and future. In: J. S. Gray and M. E. Christianson (Editors), Marine Biology of Polar Regions and Effects of Stress on Marine Organisms. Wiley, Chichester, pp. 167-184. Powell, R. D., 1981. A model for sedimentation by tidewater glaciers. Ann. Glaciol., 2: 129-134. Powel], R. D., 1983. Glacial-marine sedimentation processes and lithofacies of temperate tidewater glaciers, Glacier Bay, Alaska. In: B. F. Molnia (Editor), GlacialMarine Sedimentation. Plenum, New York, N.Y., pp. 185-232. Rao, C. P., 1981. Criteria for recognition of cold-water carbonate sedimentation: Barriedale Limestone (Lower Permian), Tasmania, Australia. J. Sediment. Petrol., 51: 491 506. Reid, A. M., 1924. The oil shale resources of Tasmania, I. Mineral Resources. Geol. Surv. Tasmania, 8. Rey, F. and Leong, H., 1985. The influence of ice and hydrographic conditions on the development of phytoplankton in the Barents Sea. In: J.S. Gray and M. E. Christianson (Editors), Marine Biology of Polar Regions and Effects of Stress on Marine Organisms. Wiley, Chichester, pp. 49-63. Singer, J. K. and Anderson, J. B., 1984. Use of total grain size distributions to define bed erosion and transport for poorly sorted sediments undergoing simulated bioturbation. Mar. Geol., 57: 335-359. Smith, W.O. and Nelson, D.M., 1985. Phytoplankton bloom produced by a receding ice edge in the Ross Sea:

372 spatial coherence with the density field. Science, 227: 163-166. Stanley, D. J., 1981. Unifites--structureless muds of gravity-flow origin in Mediterranean basins. Geo-Mar. Lett., 1: 77-83. Tucker, M. E., 1983. Sedimentation of organic-rich limestones in the Late Precambrian of southern Norway. Precambrian Res., 22: 295-315. Tucker, M. E., 1986. Formerly aragonitic limestones associated with tillites in the Late Proterozoic of Death Valley, California. J. Sediment. Petrol., 56: 818-830. Wall, D., 1962. Evidence from recent plankton regarding the biological affinities of Tasrnanites Newton 1875 and Leiosphaeridia Eisenack 1958. Geol. Mag., 99: 353-362. Walter, M.R. and Bauld, J., 1983. The association of sulphate evaporites, stromatolitic carbonates and glacial sediments: examples from the Proterozoic of Australia and the Cainozoic of Antarctica. Precambrian Res., 21: 129-148.

Williams, G. E., 1979. Sedimentology, stable-isotope geochemistry and paleoenvironment of dolostones capping late Precambrian glacial sequences in Australia. J. Geol. Soc. Aust., 26: 377-386. Williams, G. L., 1978. Dinofiagellates, acritarchs, and tasmanitids. In: B.U. Haq and A. Boersma (Editors), Introduction to Marine Micropaleontology. Elsevier, New York, N.Y., pp. 293-326. Wilson, D. L. and Smith, W. O., 1985. Species-specific productivity in an ice-edge phytoplankton bloom in the Ross Sea. Antarct. J. U.S., 20: 132-133. Wright, R., Anderson, J. B. and Fisco, P. P., 1983. Distribution and association of sediment gravity flow deposits and glacial/glacial-marine sediments around the continental margin of Antarctica. In: B. F. Molnia (Editor), Glacial-Marine Sedimentation. Plenum Press, New York, N.Y., pp. 265-300. Znachko-Yavorskiy, G. A., 1979. Characteristics of the initial stages of recent sedimentation in the coastal regions of East Antarctica. Antarktika, 8: 45-53.