The heavy mineral record in the Pliocene to Quaternary sediments of the CIROS-2 drill core, McMurdo Sound, Antarctica

ELSEVIER

Sedimentary Geology 128 (1999) 223–244

The heavy mineral record in the Pliocene to Quaternary sediments of the CIROS-2 drill core, McMurdo Sound, Antarctica Werner Ehrmann Ł , Kerstin Polozek Institut fu¨r Geologische Wissenschaften und Geiseltalmuseum, Domstraße 5, D-06108 Halle (Saale), Germany Received 23 July 1998; accepted 11 June 1999

Abstract The heavy mineral assemblages of the Lower Pliocene to Quaternary glacial sediments of the CIROS-2 drill site, situated near the mouth of Ferrar Glacier on the McMurdo Sound shelf, Antarctica, were analysed by optical means in order to reconstruct the source areas and dynamics of the late Cenozoic Antarctic ice masses. The assemblages are dominated by pyroxenes, amphiboles, altered minerals and opaque minerals. Within the pyroxene group, clinopyroxenes are most abundant; within the amphibole group, green and brown hornblendes are dominant. Other important heavy minerals present in minor amounts are zircon, titanite, epidote, garnet and apatite. The downcore distribution of the heavy minerals allows a subdivision of the sedimentary sequence into two major units, which can both be further subdivided into two subunits. The Pliocene interval between 166 and 110 mbsf is characterized by relatively high concentrations of apatite, zircon, titanite, garnet, epidote and green hornblende. This assemblage accounts for some 40–60% of the heavy minerals and points to a source to the west of the drill site, in the Transantarctic Mountains, where intrusive, metamorphic and sedimentary rocks cover large areas. The interval 166–137 mbsf (Lower Pliocene) contains additional brown hornblende and palagonite, indicative of a volcanic, probably hyaloclastitic source, which possibly has to be sought beneath the ice of the Ferrar Glacier. In contrast, the uppermost Pliocene and Quaternary interval between 110 mbsf and the top of the core is characterized by high concentrations of clinopyroxenes, altered minerals and opaque minerals. This record points to a main source area in the south or east, where large areas are occupied by basaltic rocks of the McMurdo Volcanic Group. A slight change in the composition of the pyroxene group occurs within this interval, at 48 mbsf. Thus, the heavy minerals in the sediments of core CIROS-2 document different source areas and therewith major changes in the ice dynamics. During the Pliocene the ice discharged through the Transantarctic Mountains into the McMurdo Sound. The source area shifted in the Quaternary to the south, to the region of the present-day Ross Ice Shelf.  1999 Elsevier Science B.V. All rights reserved. Keywords: Antarctica; McMurdo Sound; heavy minerals; source area; ice dynamics

1. Introduction The Antarctic ice sheet is one of the most important features controlling the climate of our planet. It Ł Corresponding

author. Present address: Institut fu¨r Geophysik und Geologie, Talstraße 35, D-04103 Leipzig, Germany. Tel.: C49-341-9732886; Fax: C49 341 9732809; E-mail: [email protected]

influences the level of the world oceans, the global atmospheric circulation, the global oceanographic circulation and the production of cold and dense bottom waters that penetrate far into the Northern Hemisphere. In order to understand the processes that are responsible for Cenozoic changes in climate and sea level, therefore, the investigation of the long-term history of the Antarctic ice sheet is of fun-

0037-0738/99/$ – see front matter  1999 Elsevier Science B.V. All rights reserved. PII: S 0 0 3 7 - 0 7 3 8 ( 9 9 ) 0 0 0 7 1 - 8

224

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

damental importance. With our current concern for possible future climatic changes, both the short-term and the long-term behaviour of the Antarctic ice is of special interest. In recent years, with the efforts of the Ocean Drilling Program, much progress has been made in reconstructing the Antarctic Cenozoic climate. Most studies concentrated on areas far off the Antarctic continent, because the sedimentary record there is relatively complete, and the dating of the mainly pelagic sediments normally causes no problems. However, it is often difficult to decipher the Antarctic climate and glacial history unambiguously from these records. Thus, the behaviour of the Antarctic ice sheet during Cenozoic time can best be studied from sediments deposited in proximal marine settings, because there direct evidence exists for ice advances and retreats through time, e.g. sediments directly deposited from grounded or floating ice, proximal to distal glaciomarine sediments and glacial erosional features. Onshore outcrops, in contrast, are rare and their sedimentary record is incomplete. The Ross Sea area is especially suitable for studying the Cenozoic glacial history, because large parts of both the present-day West Antarctic ice sheet and the East Antarctic ice sheet discharge into this embayment. The East Antarctic ice sheet drains directly into the Ross Sea by a number of glaciers breaching the Transantarctic Mountains, and indirectly by glaciers feeding the Ross Ice Shelf. Two thirds of the present ice frontage of the Ross Ice Shelf, however, are derived from large ice streams discharging from the West Antarctic ice sheet (Hambrey and Barrett, 1993). Thus, the Ross Sea and McMurdo Sound are situated in a strategic position to document the influence of the ice masses through time. This paper concentrates on the Lower Pliocene to Quaternary heavy mineral assemblages in the drill core CIROS-2, which was recovered very close to the terminus of a present-day glacier discharging through the Transantarctic Mountains. Because the recovered sediments consist mainly of terrestrial components made available by erosion on the Antarctic continent, the heavy mineral assemblages are good indicators for the source areas of the sediments and are therefore a valuable tool for reconstructing the glacial history of the Antarctic hin-

terland and the dynamics of the ice masses (cf. Diekmann and Kuhn, 1999; Polozek and Ehrmann, 1999).

2. Location, stratigraphy and lithology of the drill core CIROS-2 was drilled in 1984 from the sea ice off the Victoria Land Coast on the western shelf of McMurdo Sound at latitude 77º410 S and longitude 163º320 E, in a water depth of 211 m (Fig. 1; Barrett and Scientific Staff, 1985). The drill site is situated in a very proximal setting in respect to the East Antarctic ice sheet: it is only 1.2 km east of the floating tongue of the present-day Ferrar Glacier, which flows from the polar plateau through the Transantarctic Mountains and drains the East Antarctic ice sheet into a fjord. The Ferrar Fjord is a typical U-shaped glacial valley. To the north of Ferrar Fjord are similar valleys, like the Taylor Valley and Wright Valley, that today are ice-free. The valleys were probably cut down by temperate outlet glaciers some time after the early Cenozoic (ca. 50 Ma) uplift phase of the Transantarctic Mountains (Gleadow and Fitzgerald, 1987; Fitzgerald, 1992, 1994). The stratigraphy of the 166 m long CIROS-2 core ranges from Lower Pliocene to Quaternary (Fig. 2). The Pliocene part of the core is dated by wellpreserved marine diatom assemblages. The oldest sediment is ca. 4.5 Ma. The upper 100 m of the core are poorly dated because of the absence of marine diatoms between 98 and 8 mbsf (metres below sea floor), but are considered to be Quaternary (Harwood, 1986; Winter and Harwood, 1997). The CIROS-2 core (Fig. 2) penetrates into crystalline basement. The lower, Pliocene part of the sedimentary sequence (165–100 mbsf) is dominated by thick diamictite beds interpreted mainly as lodgement tills and proximal glaciomarine sediments. They are interbedded with some thin mudstone beds which contain ice-rafted components, indicative of a distal glaciomarine setting. The upper 100 m consist of an alternation of thick sandstones with diamictites. Compared to the diamictites of the lower part, they are thinner, much less abundant and contain less mud but more sand. The diamictites represent periodic ice advances, whereas the sandstones are as-

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

225

Fig. 1. Location of the drill site CIROS-2 on the continental shelf of McMurdo Sound in the Ross Sea, Pacific sector of the Antarctic Ocean. The positions of the drill sites CIROS-1, MSSTS-1 and CRP-1 mentioned in the text are also indicated. Bedrock geology is from Warren (1969).

sumed to represent sedimentation in an ice-dammed lake (Barrett and Scientific Staff, 1985; Barrett and Hambrey, 1992; Hambrey and Barrett, 1993). The lithologies of the CIROS-2 core suggest glacial influence throughout deposition. Glaciers therefore must have reached the nearby coast at all times documented in the core and did not retreat from the coast or even disappear from Antarctica. However, the sediments record a large number of individual ice advances and retreats.

3. Geology of the McMurdo Sound region The geology of the southern and eastern parts of the McMurdo Sound is characterized by alkali volcanic complexes (Fig. 1). They occur in the region of Ross Island, White Island and Black Island. The complex of Brown Peninsula, Mt. Discovery and Mt. Morning is also built up by volcanic rocks (Cole and Ewart, 1968; Kyle, 1981, 1990; LeMasurier and Thomson, 1990). Basaltic shield volcanoes and stratovolcanoes, dominantly trachitic or phonolithic in composition, are general features of this region.

226

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

Fig. 2. Lithology and interpretation of the sedimentary sequence at site CIROS-2 in McMurdo Sound. Simplified after Barrett and Hambrey (1992) and Hambrey and Barrett (1993). Stratigraphy from Harwood (1986) and Winter and Harwood (1997).

The oldest surface samples of this region are 19 Ma old. From 19 to 10 Ma, trachitic rocks were dominant. Most activity, however, occurred in the

last 10 Ma with basanitic rocks being dominant, but phonolitic rocks forming the active volcano of Mount Erebus (Kyle, 1990). According to magnetic

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

surveys, many more volcanic centres of similar size but unknown age exist beneath the present Ross Sea continental shelf, the Ross Ice Shelf and the West Antarctic ice sheet (Behrendt et al., 1994, 1995). In contrast, South Victoria Land comprising the part of the Transantarctic Mountains west of McMurdo Sound consists of a crystalline basement (Fig. 1). The oldest rocks belong to the Late Precambrian to early Palaeozoic Koettlitz Group. They cover wide areas and strike in a NW–SE direction. They consist mainly of amphibolite grade schists, paragneisses and marbles. The sequence is intruded by a variety of early Palaeozoic granitoid plutons, gabbros, and lamprophyre dykes, which are usually referred to as the Granite Harbour Intrusive Complex (Haskell et al., 1965; Warren, 1969; Laird and Bradshaw, 1982; Findlay et al., 1984; Allibone, 1987; Tingey, 1991a; Smillie, 1992). The basement rocks in the Transantarctic Mountains are overlain by flat-lying sedimentary rocks of the Devonian to Triassic Beacon Supergroup. These sediments, however, are restricted to the upper, western part of the valley of Ferrar Glacier and do not crop out within some 20 km of the terminus of the glacier (Fig. 1). The sedimentary rocks of the Beacon Supergroup comprise mainly non-marine friable sandstones, siltstones and quartzites, but also some coal measures and tillites (Nathan and Schulte, 1968; Young and Ryburn, 1968; La Prade, 1982). The sedimentary strata and, less commonly, the basement rocks are intruded by diabase sills and dykes of the Jurassic Ferrar Dolerite. The individual sills are in the range of 100–300 m thick and have an aggregate thickness of ca. 1000 m in South Victoria Land (Gunn, 1962; Fleming et al., 1997). The extrusive equivalent of the Ferrar Dolerite is the Kirkpatrick Basalt, which once may have formed extensive lava plateaus (e.g. Brotzu et al., 1988, 1992). The subaerial outcrops of Kirkpatrick Basalt, which are nearest to the Ferrar Glacier, are in the areas of Allan Hills and Prince Alberts Mountains, some 120 km to the north. The Ferrar Dolerite and Kirkpatrick Basalt are commonly combined to form the Ferrar Group. Scattered small occurrences of subaerially erupted olivine basanite assigned to the McMurdo Volcanic Group comprise only a few percent of the exposed rocks within the Transantarctic Mountains. These

227

basanites form cinder cones of generally <250 m diameter or small flows <10 m thick (Kyle, 1990; Wilch et al., 1993). In Taylor Valley, the dry valley north of the valley occupied by the Ferrar Glacier (Fig. 1), the cones have ages of 4.8–1.6 Ma according to Kyle (1990) or 3.9–1.5 Ma according to Wilch et al. (1993). The East Antarctic craton beyond the Transantarctic Mountains is assumed to consist mainly of Precambrian metamorphic and igneous rocks. Details, however, are unknown because of the extensive ice cover (Tingey, 1991a).

4. Methods Heavy minerals were analysed in the fine sand fraction (63–125 µm) of 37 samples from CIROS-2. They were separated from the light minerals in a centrifuge with a sodium–metatungstate solution (density D 2.83 g=cm3 ) as a heavy liquid. The heavy minerals were fixed with Meltmount (refractive index D 1.68) on glass slides and were identified under a polarising microscope. In each sample at least 300 grains comprising both translucent, opaque and weathered minerals were counted along traverses. The results are presented as grain percentages in the total heavy mineral fraction. For deciphering the silt composition and especially its quartz content, 500 mg of the 2–63 µm fraction was ground and mixed with 100 mg Al2 O3 as an internal standard before being analysed by X-ray diffraction (XRD). The results are presented as mineral=standard ratios. Volcanic glass concentrations have been determined optically in the fraction of 125–250 µm. The composition of the gravel fraction was characterized only semi-quantitatively, because the number of clasts in the samples was too low to gain statistically reliable data. All raw data (Appendix A) are available via the internet from the data bank of the Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany (www.pangaea.de).

228

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

5. Characteristics and distribution of heavy minerals In total, 61 different heavy mineral species have been identified in the fine sand fraction of the CIROS-2 sediments. However, many of these minerals are present in only very minor amounts and do not show distinct variations in their distribution patterns. Therefore they are not discussed in this paper. In general, the heavy mineral assemblages of the drill core are dominated by four main components: pyroxenes, amphiboles, altered minerals and opaque minerals. Within the pyroxene group, clinopyroxenes are most abundant; within the amphibole group, green and brown hornblendes are dominant. Other important heavy minerals present in minor amounts are zircon, titanite, epidote, garnet and apatite (Figs. 3 and 4). 5.1. Pyroxenes The pyroxene grains mainly exhibit a prismatic habit and are characterized by ragged margins. With increasing weathering typical cockscombs occur and the habit becomes more and more skeletal. No, or only minor, cleavage is visible. Most of the pyroxenes show very pale to pale colours and weak pleochroism. The orthopyroxenes have been distinguished from the clinopyroxenes by their straight extinction. However, they were only found in minor amounts and relatively constant concentrations. This paper, therefore, does not treat the individual pyroxenes separately, but concentrates on the pyroxenes as one large mineral group. Titanaugite is the only pyroxene that is discussed separately. Titanaugite is a clinopyroxene occurring as mainly euhedral to subhedral grains with brown to violet-brown mineral colours and anomalous interference colours of leather-brown and blue-grey. Pyroxenes occur with around 10% below 110 mbsf, and about 30–50% above 110 mbsf (Fig. 4). 5.2. Hornblendes Because of their high abundance, the hornblendes have been subdivided according to their colour. Green and brown hornblendes are most common. The pleochroism of the former group ranges from

green to brown, whereas the latter group is characterized by a pleochroism from brown to red. Both green and brown hornblendes occur as elongated or short prismatic mineral grains. Most of them do not exhibit the typical cleavage pattern for amphiboles. The green and brown hornblende grains show evidence of dissolution in different intensity, from ragged margins to skeletal appearance. Some of the brown hornblendes have a rim consisting of glass or chlorite (?), but nevertheless conserve their strong prismatic habit. Other amphiboles observed in minor amounts are blue-green amphibole, tremolite, anthophyllite and actinolite (?). Hornblende comprises up to 60% of the heavy mineral fraction in the lower part of the CIROS-2 core, below 110 mbsf. Concentrations of 20–30% occur in the interval 110–48 mbsf and concentrations of ca. 10% in the uppermost 48 m. This distribution pattern is strongly controlled by the abundance of green hornblende (Fig. 3), with 20–35% below 110 mbsf, ca. 10% in the interval 110–48 mbsf and 5% in the upper 48 mbsf. The brown hornblendes, in contrast, show maximum concentrations of ca. 30% below 137 mbsf, medium values of ca. 15% at 110– 48 mbsf and minimum values of <5% at 137–110 mbsf and 48–0 mbsf (Fig. 4). 5.3. Altered minerals As altered minerals we combined all brown to black mineral grains which were so intensely altered that no assignment to a distinct mineral group was possible. Altered grains that still allowed an identification of the original mineral were included into the individual mineral groups. Altered minerals occur in only minor amounts below 100 mbsf, but with 10– 20% above (Fig. 4). 5.4. Opaque minerals Opaque minerals have not been analysed in detail, but most of them seem to be magnetites. All opaque minerals have been counted as one group. Below 110 mbsf, <10% opaque minerals are present. Above 110 mbsf, an increasing trend to about 20% is obvious (Fig. 4).

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244 229

Fig. 3. Percentage distribution of the main heavy mineral groups of Association 1 (apatite, zircon, titanite, garnet, epidote) and Association 2 (green hornblende) in the CIROS-2 core.

230 W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244 Fig. 4. Percentage distribution of the heavy mineral groups of Association 3 (pyroxene, titanaugite, altered minerals, opaque minerals) and Association 4 (brown hornblende, palagonite) in the CIROS-2 core.

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

5.5. Other heavy minerals The four main heavy mineral groups mentioned above comprise ca. 65–95% of the total assemblage. Other diagnostic heavy minerals are zircon, titanite, epidote, garnet and apatite. They comprise 1–35% of the total assemblage. The zircon grains are typically subangular to rounded. They occur in colourless, red and greybrown varieties. Also the titanite grains are subangular to rounded. They have a typical honey-yellow colour, but some red and brown grains also have been found. The interference colours are anomalous blue and brown. The minerals epidote and clinozoisite have been combined in the epidote group. Epidote minerals typically occur as pistachio-green grains, but some colourless grains are also present. Clinozoisite minerals are colourless and can be best identified by their anomalous shining blue interference colour. Garnet occurs as subangular to rounded isometric grains and angular fragments. They are mainly colourless, but some minor pink grains are also present. The colourless apatite grains are mainly well rounded. Angular components occur in only very minor amounts. All these minerals show similar downcore concentration patterns (Fig. 3). They have minimum concentrations in the upper 110 m of the core and maximum values at 137–110 mbsf. In the lowermost part of the core, apatite and zircon have medium concentrations, whereas titanite, epidote and garnet have minimum concentrations, similar to those in the upper part of the core.

6. Sources for the heavy minerals 6.1. Association 1: apatite–zircon–titanite–garnet–epidote The heavy minerals apatite, zircon, titanite, garnet and epidote show almost identical downcore distribution patterns (Fig. 3). A common source area for these minerals is therefore assumed. Whereas apatite can be found in almost all magmatic rocks, zircons are most common in acid and intermediate magmatic rocks. Titanite is a general accessory mineral in plutonic rocks and pegmatites,

231

but occurs only in very minor amounts in volcanic rocks. Therefore an origin from basement rocks, such as being widespread in the Transantarctic Mountains, is very likely. In fact, all three minerals are well-known accessory constituents of the Granite Harbour Intrusive Complex (Ghent and Henderson, 1968; Skinner and Ricker, 1968a; Smillie, 1992). The zircon and titanite grains are subangular to rounded, most of the apatite grains are even well rounded. The rounded grains could indicate a recycling of sedimentary rocks. The only source for recycled minerals are the sediments of the Beacon Supergroup, which are known to contain both zircon and apatite (Laird and Bradshaw, 1982; Skinner and Ricker, 1968b). However, to our knowledge, titanite is not described as a constituent of the Beacon Supergroup. Within the zircons, the colourless, pink and greybrown varieties show almost identical distribution patterns. Colourless zircons are common in the sediments of the Beacon Supergroup (Laird and Bradshaw, 1982). The pink to red colour of the zircons is due to a long-term influence of radioactive radiation, with the intensity of the colour increasing with radiation and geological age. Therefore, red zircons are generally thought to be derived from Precambrian parent rocks (Zimmerle, 1972), such as the basement rocks in the Transantarctic Mountains. The colour of the brown-grey zircons, the so-called metamict zircons, is caused by the radioactive radiation of the U and Th of the zircons themselves (Mange and Maurer, 1991) and these zircons give no indications of their provenance. Garnet and epidote minerals point to a source characterized by metamorphic rocks, such as the metasediments and orthogneisses of the Koettlitz Group in the Transantarctic Mountains (Riddolls and Hancox, 1968; Allibone, 1987). Some grains have a subrounded to rounded shape. Because both minerals are resistant to transport, they also could be recycled from sandy sedimentary rocks, such as those of the Beacon Supergroup, where they occur as accessory minerals (Skinner and Ricker, 1968b; Laird and Bradshaw, 1982; La Prade, 1982). The subrounded to rounded nature of many grains within Association 1 could also be ascribed to abrasion during transport. Glacial transport can be assumed at least for the diamictites in the lower part

232

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

of CIROS-2 with maximum concentrations of Association 1 minerals. Abrasion during glacial transport is restricted mainly to the basal debris of a glacier or ice sheet and to lodgement tills (e.g. Boulton, 1978). However, because other heavy minerals, such as pyroxenes and amphiboles, as well as the gravel components are mainly angular to subangular, this process was obviously not very effective. The roundness of some of the grains is rather inherited from the host rock and=or from recycling of sedimentary rocks. Because not all heavy mineral grains of Association 1 show some roundness, only a part of the minerals is believed to be derived from the Beacon Supergroup, whereas the remainder part comes probably from basement rocks. To summarize, the heavy minerals of Association 1 indicate a source in the surrounding of Ferrar Glacier in the Transantarctic Mountains (Fig. 1). The Granite Harbour Intrusive Complex or equivalent

rocks probably are responsible for angular apatite, zircon and titanite, and the metamorphic rocks of the Koettlitz Group or equivalent rocks probably are responsible for angular garnet and epidote. The rounded heavy mineral grains of this association are probably recycled from the sediments of the Beacon Supergroup (Table 1). 6.2. Association 2: green hornblende In the hornblende group we distinguished between green and brown hornblendes (Figs. 3 and 4; Table 1). Green hornblendes are ubiquitous components of most intrusive and metamorphic rocks. In the case of the CIROS-2 sediments, the most likely sources for the green hornblende thus are the basement rocks in the Transantarctic Mountains, especially the amphibolites of the Koettlitz Group. Green to green-brown hornblendes were described from the

Table 1 Main source rocks and provenance areas assumed for the heavy minerals in the CIROS-2 sediments Mineral

Source rocks

Provenance

Association

Apatite

Beacon Supergroup Ferrar Group Granite Harbour intrusive complex

TAM

1

Beacon Supergroup, crystalline basement Precambrian basement ?

TAM TAM

1

Granite Harbour intrusive complex

TAM

1

TAM

1

rounded

Koettlitz Group Granite Harbour intrusive complex Beacon Supergroup

Epidote

rounded angular

Beacon Supergroup Koettlitz Group

TAM TAM

1

Hornblende

green

Koettlitz Group Granite Harbour intrusive complex

TAM

2

Pyroxene

McMurdo Volcanic Group

RIS

3

Alterites

omnipresent McMurdo Volcanic Group

RIS

3

Opaques

omnipresent McMurdo Volcanic Group

RIS

3

Palagonite

McMurdo Volcanic Group

TAM

4

hyaloclastitic rocks McMurdo Volcanic Group

TAM

4

Zircon

colourless pink metamict

Titanite Garnet

Hornblende

subangular

brown

TAM

Detailed discussion in the text. TAM D Transantarctic Mountains, RIS D region of the present-day Ross Ice Shelf.

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

Granite Harbour Intrusive Complex of Taylor Valley (Ghent and Henderson, 1968; Skinner and Ricker, 1968a). Light green to colourless amphiboles (tremolite= ?actinolite), in contrast, are typical of the amphibolite schists of the basement (Ghent and Henderson, 1968). In CIROS-2, these minerals occur in only minor amounts of <5%.

233

An interpretation of the provenance of the altered minerals and opaque grains is more difficult, because the minerals could not be identified and because opaque minerals are constituents of all rocks in the hinterland of CIROS-2. However, according to Wimmenauer (1985) high percentages of opaque minerals in general also indicate a volcanic source. 6.4. Association 4: brown hornblende–palagonite

6.3. Association 3: (clino-)pyroxenes–altered minerals–opaque minerals A further correlation in the downcore distribution is obvious between the pyroxenes, altered minerals and opaque minerals (Fig. 4). Their distribution pattern is contrary to the pattern of associations 1 and 2 (Fig. 3). A different source area is therefore very likely. Within the pyroxenes, the clinopyroxenes are dominant, whereas the orthopyroxenes comprise only a few percent of the heavy mineral fraction. Generally, clinopyroxenes are most common in basic to intermediate rocks, and titanaugites are indicative for basic volcanic rocks. Therefore, a source in the southern McMurdo Sound can be assumed, where large areas are occupied by the McMurdo Volcanic Group (Fig. 1). Clinopyroxenes are common constituents of these volcanic rocks (Cole and Ewart, 1968; Kyle, 1981, 1990; LeMasurier and Thomson, 1990). Here they show a poorly developed cleavage or no cleavage at all (Weiblen et al., 1981), comparable with the clinopyroxenes found in CIROS-2. Also titanaugite has been described from rocks of the McMurdo Volcanic Group (Nathan and Schulte, 1968). Theoretically the volcanic rocks of the Ferrar Group are another possible source for clinopyroxenes (Gunn, 1962; Brotzu et al., 1988, 1992; Elliot et al., 1995). However, the Ferrar-derived clinopyroxenes are characterized by a prominent cleavage and numerous exsolution lamellae (Cape Roberts Science Team, 1998; J. Smellie, pers. commun., 1998), features which were not observed in the CIROS-2 samples. Furthermore, the dolerites occupy only minor areas within the crystalline basement and the Beacon Supergroup. Heavy minerals provided by the Ferrar Dolerite, therefore, should be diluted by other heavy minerals and should show a similar distribution trend as basement-derived or Beacon-derived minerals.

In contrast to green hornblende, brown hornblendes like oxyhornblende and basaltic hornblende (kaersutite) are characteristic of volcanic rocks. Oxyhornblendes form from green hornblende by oxidation of the iron. Kaersutite is a Ti-rich amphibole common in tuffs, but less common in the corresponding lavas. Therefore, a volcanic source can be assumed for the brown hornblendes in CIROS-2. Brown hornblendes have been described from the hornblende trachytes, hornblende basalts and trachyandesites of the McMurdo Volcanic Group (Cole and Ewart, 1968; Nathan and Schulte, 1968; Kyle, 1990; LeMasurier and Thomson, 1990). In contrast, no significant occurrences are known within the Ferrar Group. Brown hornblendes occur in pegmatitic lenses within some Ferrar sills (Skinner and Ricker, 1968b), but probably not enough to yield more than a trace as a heavy mineral in the CIROS-2 sediments. An origin of brown hornblende from volcanic rocks is supported by maximum concentrations of brown hornblende below 137 mbsf coinciding with maximum concentrations of palagonite (Fig. 4), which result from the alteration of volcanic glass. Temperature and humidity are probably the most important factors controlling the rate of palagonitization of basaltic glass. Therefore, a reasonable explanation for the formation of palagonite could be intense alteration processes in hydrovolcanic rocks erupted subaerially and beneath the ice or in the sea (Jakobsson, 1978; Wohletz and Sheridan, 1983; Smellie and Skilling, 1994). Such rocks could be connected to the McMurdo Volcanic Group or to the Ferrar Group. Because of the strong correlation between brown hornblende and palagonite, the McMurdo Volcanic Group is the most likely source for the palagonite. Several occurrences of hyaloclastites are known from the

234

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

McMurdo Volcanic Group, which could serve as potential palagonite sources (Kyle, 1981). The Kirkpatrick Basalt includes local hyaloclastites (pillow-palagonite complexes) formed when lavas flowed into lakes or channels, but this type of rock is a very minor component of the lava sequence as a whole (Elliot et al., 1986; D. Elliot, pers. commun., 1998). Also a palagonite source in the dykes and sills of the Ferrar Dolerites is not very likely, because no palagonite and only very minor amounts of glass have been described from these dolerites. In contrast, large amounts of brown glass occur in the Kirkpatrick Basalt (Elliot et al., 1995), but are texturally quite different from the glass derived from the McMurdo Volcanic Group (Tingey, 1991b; J. Smellie, pers. commun., 1998). Thus, it seems that the source for the heavy mineral Association 4 has to be sought in the McMurdo Volcanic Group. However, there is no correlation with McMurdo-derived clinopyroxenes (Fig. 4), which are typical of Association 3. The source of this assemblage therefore will be further discussed in Section 8.

7. Stability of heavy minerals Before discussing the temporal changes in the heavy mineral record of the CIROS-2 drill core in respect to ice dynamics and glacial history, the influence of different mineral stabilities on the distribution patterns has to be considered. Some heavy minerals are unstable during chemical weathering conditions but stable against the influence of physical weathering and transport, whereas other minerals are unstable against physical weathering and transport. Thus, different climates and weathering conditions, as well as different transport mechanisms, may influence the composition of the heavy mineral associations. Samples for investigating the heavy minerals in the core CIROS-2 were taken mainly from diamictites and sandstones, but also from a few mudstones. The downcore distribution patterns of the individual heavy minerals show no dependence on the sediment facies (Figs. 3 and 4). They are also not related to the proximity of the ice, because the assemblages show no major differences between lodgement tills and

glaciomarine sediments. This suggests that the heavy mineral proportions are not influenced by transport, sedimentation or reworking processes. The glacial and glaciomarine=glaciolacustrine nature of all the sediments recovered at CIROS-2 implies that no major changes in the weathering regime occurred during the time represented by the sediments. The sediments document glacial influence throughout (Fig. 2; Barrett and Hambrey, 1992), even if the extent of the Antarctic ice may have varied considerably. Thus, a number of major ice advances and retreats are recorded in the core, but glaciers reached the sea at all times and transported sediment from the continent into the McMurdo Sound. Therefore mainly physical weathering conditions can be assumed for the Antarctic continent, whereas chemical weathering played only a minor role. The heavy mineral record of core CIROS-2 supports the idea of predominantly physical weathering. The heavy minerals comprise minerals that are both stable against chemical weathering processes and minerals that are relatively unstable. Garnet and apatite belong to the latter group (e.g. Boenigk, 1983; Morton, 1985). The two minerals occur throughout the core, in relatively high amounts in the lower part, and in trace amounts in the upper part (Fig. 3). These minerals show no obvious evidence for chemical weathering, like etching. Therefore we assume that chemical weathering was not very active during the time of erosion, transport and deposition of the CIROS-2 sediments. Pyroxenes and amphiboles are even less resistant than garnet and apatite and are prone to both chemical and physical weathering and transport (Morton, 1985; Pettijohn et al., 1987). Indications for their alteration are ragged margins, cockscombs and a skeletal habit, which have been observed in many samples. However, the occurrence of these features seems to be randomly distributed throughout the core. Thus, the intensity of the weathering seems not to have changed significantly during the time represented by the CIROS-2 sediments. To conclude, because neither transport nor weathering can be made responsible for the changes in the distribution patterns of the individual heavy minerals, the main control for the composition of the assemblages has to be the source area.

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

8. Temporal changes in provenance The downcore distribution of the heavy minerals in core CIROS-2 allows a subdivision of the sedimentary sequence into two major units, which can both be further subdivided into two subunits (Figs. 3–6). Unit 1 comprises the lower part of the core, from the bottom of the core at 166 mbsf to about 110 mbsf. Because the changes in the concentration patterns of the individual heavy minerals occur between 110 and 100 mbsf, we placed the boundary at the first distinct change of pyroxene and titanaugite at 110 mbsf (Figs. 3, 4 and 6). According to the stratigraphy by Winter and Harwood (1997), it represents the time from ca. 4.5 Ma to 2.1 Ma. Subunit 1.1 ranges from 166 mbsf to about 137 mbsf, Subunit 1.2 from ca. 137 mbsf to ca. 110 mbsf. The boundary between the two subunits can be dated to ca. 3.1 Ma (Winter and Harwood, 1997). Unit 2 comprises the upper part of the core, from about 110 mbsf to the sea floor. Subunit 2.1 ranges from ca. 110 mbsf to ca. 48 mbsf, Subunit 2.2 from 48 mbsf to 0 mbsf. The boundary between the two subunits can be dated to ca. 0.8 Ma (Winter and Harwood, 1997). As summarized in Fig. 6, Unit 1 is characterized by high concentrations of heavy minerals of Association 1 (apatite, zircon, titanite, garnet, epidote) and Association 2 (green hornblende). In contrast, the sediments are depleted in heavy minerals of Association 3 (clinopyroxene, altered minerals, opaque minerals). This heavy mineral record points to a source area in the Transantarctic Mountains. The Beacon Supergroup possibly supplied the rounded grains of Association 1. The intrusive and metamorphic rocks of the basement probably supplied the angular grains of Associations 1 and the minerals of Association 2 (Table 1). Thus, during the deposition of Unit 1, the ice probably came from the west. It discharged through the Transantarctic Mountains, where it incorporated debris of the Beacon Supergroup and the crystalline basement, and transported this detritus into McMurdo Sound. The discharge system resembled the present-day Ferrar Glacier, but was somewhat extended as indicated by the deposition of lodgement tills and proximal glaciomarine sediments at the CIROS-2 site (Fig. 2). The several advances of the

235

ice documented by the diamictites and the retreats of the ice between the advances had no influence on the composition of the heavy mineral assemblages and concentrations of individual heavy minerals. Further information on the source of the sediments comes from the petrography of the gravel clasts enclosed in the sediments (Fig. 5). Even if the number of gravel clasts that was present in our samples is too small for statistically safe interpretations, the clast data allow a qualitative description of the source. The gravel of Unit 1 contains a basement component throughout, consisting of granite, metamorphic rocks and large angular quartz grains. Sediment clasts of the Beacon Supergroup were detected in only a few samples and in only trace amounts. Probably the poorly lithified nature of large parts of the Beacon Supergroup resulted in a disaggregation of the sediments into their constituents during glacial transport. Also most of the sand grains in the CIROS-2 sediments are derived from the crystalline basement and suggest a provenance from the west (Barrett and Hambrey, 1992). The subdivision of Unit 1 is based on the presence of high amounts of heavy minerals of Association 4 (brown hornblendes, palagonite) in Subunit 1.1 and their almost absence in Subunit 1.2 (Fig. 6). The occurrence of up to 50% of Association 4 minerals in the lowermost part of the core is at the expense of the Association 1 minerals, but obviously does not very much affect the percentages of associations 2 and 3. The brown hornblende and palagonite indicate for this subunit an additional source, most probably composed of hyaloclastitic rocks with a high affinity to the McMurdo Volcanic Group. A volcanic source is supported by the volcanic clasts, some of them resembling hyalotuffs, that we found along with clasts of basement rocks in the gravel fraction of Subunit 1.1, but not of Subunit 1.2 (Fig. 5). Furthermore, the quartz content of the silt fraction (Fig. 5) is somewhat lower in Subunit 1.1 than in Subunit 1.2 and probably suggests a dilution by input of debris derived from a quartz-poor source. One theoretical explanation for the heavy mineral assemblage in Subunit 1.1 could be that the ice masses depositing the sediments at the CIROS-2 site originated in the south. They incorporated the minerals of associations 1 and 2 in the Transantarctic Mountains south or southwest of the present-day

236 W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244 Fig. 5. Relative quartz contents in the fraction 2–63 µm, volcanic glass concentrations in the fraction 125–250 µm and estimates of the abundance of gravel components from different provenances.

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

Fig. 6. Summary of the percentage distribution of the four heavy mineral associations in core CIROS-2. 237

238

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

Ross Ice Shelf. Passing the outcrops of the McMurdo Volcanic Group, they picked up the minerals of Association 4. However, this scenario is not very likely. Because the ice not only would erode hyaloclastites but mainly normal basalts, the sediments of CIROS-2 then also should contain considerable amounts of minerals of Association 3, which are typical for a basaltic source in the McMurdo Volcanic Group. This, however, is not the case (Fig. 6). Another explanation for the mixture of heavy mineral associations in Subunit 1.1 could be that the sediments were delivered by ice coming from the west, by a precursor of the present-day Ferrar Glacier. In this case, the source rocks of the associations 1 and 2 minerals are the well-known Granite Harbour Intrusive Complex, the metamorphic rocks of the Koettlitz Group and the Beacon Supergroup cropping out in the hinterland of Ferrar Glacier (Fig. 1). In contrast, the presence of heavy minerals of Association 4 in this model have to be the result of reworking of sediments deposited by an earlier advance of Ross Sea ice on the western McMurdo shelf. However, this solution is also not likely, because in this case one would expect the co-occurrence of major amounts of minerals of Association 3. The most likely scenario is that the ice came from the west, through the Transantarctic Mountains, and brought debris of the crystalline basement and the Beacon Supergroup. This debris contained the minerals of associations 1 and 2. Source rocks for the Association 4 minerals, however, are not reported in the literature to crop out subaerially in the hinterland of the Ferrar Glacier, but they may be hidden beneath the ice. A possible source could be volcanic rocks, equivalents of which are known as cinder cones from both Wright Valley and Taylor Valley in the close northern neighbourhood of Ferrar Glacier and in the Royal Society Range to the south of Ferrar Glacier (Fig. 1). These Pliocene cinder cones have a close affinity to the McMurdo Volcanic Group (Kyle, 1990; LeMasurier and Thomson, 1990). In fact, several such cinder cones also exist in the valley of the Ferrar Glacier and are very similar in age and composition to the other cinder cones in the Dry Valleys and in the Royal Society Range (T. Wilch, pers. commun., 1998). Because Wright Valley and Taylor Valley were already dry

valleys in the Early Pliocene, the eruptions occurred subaerially (Kyle, 1990; LeMasurier and Thomson, 1990; Wilch et al., 1993). In contrast, the Early Pliocene valley of Ferrar Glacier was filled by ice, as it is today. Therefore possible eruptions in this area mainly happened subglacially and produced hyaloclastites, which could serve as a source for both the brown hornblende and the palagonite. In Subunit 1.2 the brown hornblende and palagonite are missing, because the hyaloclastites were removed by the ice. Unit 2 is characterized by high concentrations of heavy minerals of Association 3 (clinopyroxene, altered minerals, opaque minerals) and low concentrations of Association 1 (apatite, zircon, titanite, garnet, epidote), Association 2 (green hornblende) and Association 4 (brown hornblende, palagonite) (Fig. 6). This heavy mineral record points to a main source area in the south or east, where large areas are occupied by basaltic rocks of the McMurdo Volcanic Group. A volcanic source is supported by the presence of basalt clasts in the gravel fraction and of relatively fresh-looking volcanic glass shards in the sand fraction (Fig. 5). It is striking that high glass concentrations are restricted to sandstones, whereas the diamictites contain only minor amounts or no glass at all. According to Barrett and Hambrey (1992) the sandstones are largely the result of sand being blown onto the surface of lake or sea ice by strong southerly winds. The sand would then have melted through the ice and settled to the lake or sea floor. This idea of primary wind transport would also explain the enrichment of glass that originated from eruptions of the volcanoes in the south of McMurdo Sound and the depletion in coarse-grained basalts. Although we postulate for the sediments of Unit 2 a main source area in the region of the present-day Ross Ice Shelf, the presence of small amounts of minerals of associations 1 and 2 (Fig. 6) indicates an additional, but much less important source in the Transantarctic Mountains. Thus, the ice possibly originated in the Transantarctic Mountains south or southwest of the present-day Ross Ice Shelf where it incorporated the heavy minerals of associations 1 and 2. The main debris load, however, was picked up in the outcrop area of the McMurdo Volcanic Group (Fig. 1). Another possibility is that the ice coming from the McMurdo Volcanic Group on its way to the north reworked associations 1 and 2 minerals on the

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

shelf of McMurdo Sound, where they were deposited during earlier ice advances of glaciers breaching the Transantarctic Mountains. The influence of a crystalline source is also obvious in the high quartz content of the sediments, and especially in the composition of the gravel clasts (Fig. 5). Granules and pebbles of intrusive and metamorphic rocks even dominate over basaltic grains. Probably a large proportion of the basaltic clasts disaggregated during the glacial transport into smaller particles or into their constituents, whereas the basement clasts and the quartz are very stable. A source area in the region of the present-day Ross Ice Shelf was also suggested by Barrett and Hambrey (1992) on the basis of a high content of basaltic grains in the sand fraction of the CIROS-2 sediments. However, according to these authors the volcanic influence started at 100 mbsf, that is about 10 m above the boundary between our Units 1 and 2. Thus, the boundary between Units 1 and 2 marks a significant change in ice flow directions, with ice discharging through the Transantarctic Mountains prior to 2.1 Ma, and ice coming from the south in later times. Several authors postulated that large marine ice sheets covering major parts of the continental shelf existed at least during the last glacial maximum, but probably also during several intervals in earlier Pleistocene and Pliocene times. During these periodic advances, the ice thickened and grounded in the Ross Sea, pushed lobes of grounded ice westward into the fjords of the McMurdo Sound and left a record of ice-dammed lakes as well as several moraines (Kyle, 1981; Stuiver et al., 1981; Denton et al., 1989; Alonso et al., 1992; Anderson et al., 1992; Kellogg et al., 1996). Ice masses coming from the south have been postulated also for two intervals in the Oligocene and one interval in the Early Miocene, based on the clay mineralogy of drill cores CIROS-1 and MSSTS-1, which were recovered some 10 km offshore CIROS-2 (Ehrmann, 1998). Furthermore, the clay minerals of the Lower Miocene and Quaternary sediments of the CRP-1 core (Fig. 1) document three intervals of sediment input from the south (Ehrmann, 1999). The subdivision of Unit 2 into two subunits is based on the decrease in the content of titanaugite and hornblende and the slight increase in the con-

239

centration of opaque minerals at 48 mbsf (Figs. 3 and 4). The minerals of Association 1 only occur in trace amounts in Unit 2 and therefore show no concentration change (Fig. 6). The shift in the pyroxene composition indicates a minor change of the geology of the source area. The decreasing concentration of titanaugite points to a less basic composition of the volcanics providing the glacial debris. The ice probably chose a slightly different path within the outcrop area of the McMurdo Volcanic Group and therewith over-rode volcanic rocks with a slightly different composition. Three major groups of volcanic rocks can be distinguished in the McMurdo Volcanic Group: evolved volcanics (phonolites and trachytes), intermediate volcanics (phono-tephrites and tephro-phonolites) and basic volcanics (basanites and tephrites) (Kyle, 1990; LeMasurier and Thomson, 1990).

9. Conclusions The present study has demonstrated the significance of heavy mineral assemblages for reconstructing source areas and transport directions of the Antarctic ice through time. Although the individual heavy minerals do not allow an unequivocal interpretation, the heavy mineral associations are a valuable tool for reconstructing the glacial history of the area of the southern McMurdo Sound and Ferrar Glacier. Four different associations have been distinguished. Association 1 consists of apatite, zircon, titanite, garnet and epidote. It indicates a provenance from the Transantarctic Mountains, with the Granite Harbour Intrusive Complex, the metamorphic Koettlitz Group and the sediments of the Beacon Supergroup as source rocks. Green hornblende forms Association 2, which also points to a source in the crystalline basement of the Transantarctic Mountains. Association 3 comprises pyroxenes (mostly clinopyroxenes), altered minerals and opaque minerals. A source in the McMurdo Volcanic Group in the southern McMurdo Sound can be assumed for this association. Finally, brown hornblende and palagonite are grouped to form Association 4 and are believed to represent a hyaloclastitic source hidden beneath the ice of Ferrar Glacier.

240

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

The lowermost part of the core CIROS-2, between 166 and 137 mbsf, documents the time from ca. 4.5 to 3.1 Ma. It is characterized by a source in the Transantarctic Mountains as indicated by the dominance of associations 1, 2 and 4. The minerals of associations 1 and 2 are derived from the crystalline basement and the Beacon Supergroup, whereas the minerals of Association 4 show some influence of subglacial volcanism. At 3.1 Ma the volcanic source was exhausted and the crystalline basement and the Beacon Supergroup provided the overwhelming part of the heavy minerals. A major change in the source area and the ice flow happened at ca. 2.1 Ma and is documented in a depth of 110 mbsf. After that date, the heavy minerals of Association 3 indicate a main source in the region of the present-day Ross Ice Shelf, where large areas are occupied by basaltic rocks of the McMurdo Volcanic Group. However, the presence of small amounts of minerals of associations 1 and 2 indicates an additional, but much less important source in the Transantarctic Mountains or a recycling of sediments on the floor of McMurdo Sound. A minor change in the heavy mineral assemblages at ca. 0.8 Ma possibly documents a minor shift in the source area, which, however, remained in the outcrop area of the

McMurdo Volcanic Group. We only can speculate on the causes for the changes at 2.1 and 0.8 Ma. They could be due to uplift of the Transantarctic Mountains, to variations in the intensity of the glaciation, or to changing characteristics of the ice. The heavy mineral data from core CIROS-2 corroborate earlier ideas on temporal variations in the ice dynamic and support the models postulating periodic ice advances from the south, with ice grounded in the Ross Sea and pushing lobes of ice westward into the fjords of the McMurdo Sound.

Acknowledgements All laboratory work for this study was carried out at the Alfred Wegener Institute for Polar and Marine Research in Bremerhaven, Germany. We thank our colleagues at Bremerhaven for their support and good cooperation. Helga Rhodes is appreciated for her technical assistance. We further express our gratitude to David Elliot, John Smellie and Bernhard Diekmann for advice, critical remarks and discussions. The manuscript benefitted from the reviews by A. Lisitzin and D. Pirrie. Financial support was provided by the Deutsche Forschungsgemeinschaft.

Appendix A Relative abundances of individual heavy minerals in the fraction 63–125 µm of sediments from the CIROS-2 core, volcanic glass concentrations in the fraction 125–250 µm, and relative quartz contents in the fraction 2–63 µm

0.01 0.10 0.50 8.49 12.50 16.49 20.48 28.78 31.64 37.80 41.09 42.29 44.26 47.66 49.66 54.33 58.65 60.75 65.68 70.42 70.47 73.30 79.59 82.38 84.85 85.95

0.2

0.0

0.7

0.2

0.9

2.5

25.7

32.2

5.6

0.0

1.6

4.3

18.4

0.3 0.8 0.8 0.3 0.0 1.2 0.3 0.3

0.6 0.8 0.3 0.0 0.0 0.0 0.3 0.0

0.0 0.3 0.0 0.9 0.6 0.9 0.0 0.7

0.0 0.0 0.0 0.0 0.3 0.3 0.0 0.3

0.3 0.0 0.3 1.4 1.2 0.3 1.0 0.7

4.5 7.8 7.9 8.5 5.8 5.5 5.4 8.6

43.0 42.3 33.9 36.8 39.6 48.9 50.8 43.2

20.2 20.7 20.4 19.7 21.1 17.1 12.4 13.9

13.8 14.4 14.6 16.5 15.6 12.5 16.8 20.8

0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0

2.0 1.3 2.4 4.3 5.2 4.0 3.8 2.6

9.3 7.6 14.8 8.6 4.3 5.8 5.4 4.6

0.5 0.6 0.9

0.0 0.0 0.6

0.8 0.0 0.6

0.0 0.3 0.9

1.4 2.4 0.3

8.2 6.3 12.2

44.0 32.6 30.9

16.0 22.8 10.9

16.0 22.5 16.3

0.0 0.0 0.0

3.3 2.4 6.6

4.9 3.0 15.0

1.0

0.5

0.8

0.0

0.3

17.7

28.3

8.5

9.0

0.0

13.1

17.2

0.3

0.0

0.3

0.3

2.5

11.6

21.2

20.3

23.7

0.0

5.1

3.7

0.0 0.0 1.4 1.2 20.3 3.9 7.1 0.5 0.6 16.1 2.6 0.5 1.9 0.0 0.3 15.4

1.8 1.1 1.7 1.2 0.3

1.6 0.5 2.0 0.6 0.3

0.8 0.8 1.4 0.3 1.0

0.0 0.0 0.0 0.3 0.3

0.0 0.8 0.3 0.6 1.0

13.9 15.6 19.8 13.9 11.6

28.3 31.1 30.9 25.8 29.2

13.9 2.6 2.6 12.1 8.3

10.2 16.1 11.3 15.1 9.6

0.0 0.0 0.0 0.0 0.0

7.3 18.5 11.9 13.9 16.3

18.1 4.2 15.6 7.7 18.3

0.0 0.0 11.8 0.6 17.0 0.6

2.1 1.3 1.0 1.1 3.8 3.2 1.1 2.1 1.3 2.3 2.4 2.3 2.7 2.9 1.5 2.3 0.9 2.8 3.5

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

Depth Association 1 Association 2 Association 3 Association 4 % Ti-augite % Glass Qz=std (mbsf) % apatite % zircon % titanite % garnet % epidote % green hbl % pyroxene % opaque % altered min. % palagonite % brown hbl

3.1 2.8 1.7 2.3 1.4

241

242

Appendix A (continued) Depth Association 1 Association 2 Association 3 Association 4 % Ti-augite % Glass Qz=std (mbsf) % apatite % zircon % titanite % garnet % epidote % green hbl % pyroxene % opaque % altered min. % palagonite % brown hbl 0.3

0.3

0.3

0.0

0.6

8.0

31.9

13.2

15.3

0.0

13.5

12.6

0.9 1.0 2.9 3.6 9.2 7.1 17.2 18.1

0.3 0.6 0.0 5.4 7.2 4.9 13.6 10.2

0.9 1.6 1.0 1.8 1.6 2.9 2.9 3.2

0.9 1.0 0.3 1.8 1.3 2.3 1.6 1.3

1.2 0.0 0.6 1.8 3.0 2.6 6.5 4.8

7.0 17.0 17.3 42.3 39.8 42.7 28.1 30.8

29.1 30.0 34.8 14.1 5.3 22.0 8.1 6.3

10.8 4.7 4.5 5.4 7.2 2.3 9.1 8.6

13.4 4.1 4.8 5.1 4.9 3.6 1.6 2.9

0.0 0.0 0.0 0.0 0.0 0.0 0.0 0.0

10.8 18.3 16.6 8.7 6.3 2.9 2.6 1.6

19.5 9.5 13.7 0.9 0.3 0.0 0.0 0.3

0.7 6.7

3.2 5.5

0.3 2.8

0.7 0.6

1.3 4.3

17.2 25.1

27.5 4.3

1.0 22.6

13.9 3.4

0.0 0.0

2.9 0.6

1.0 0.0

7.1 10.4 7.2

5.6 3.6 2.5

1.5 1.5 1.3

0.3 0.0 0.6

1.5 0.9 0.6

33.8 21.1 21.6

9.7 6.0 5.0

5.3 3.0 5.6

3.8 2.7 3.8

0.0 16.7 15.4

28.2 28.3 33.5

0.9 0.0 0.0

5.9 17.1 4.3

2.6 1.6 3.5

1.3 0.6 0.0

0.0 0.0 0.3

2.0 1.3 0.9

31.3 7.9 26.4

12.8 17.7 7.5

1.0 0.6 4.3

5.3 13.9 3.7

3.6 0.0 18.4

28.3 0.6 26.2

0.0 0.0 0.3

5.6

5.9

0.6

0.6

0.6

27.4

10.9

3.7

4.4

16.5

18.4

0.0

6.2 6.1

4.3 6.1

0.0 0.9

0.6 0.0

0.9 1.5

24.5 21.1

6.8 8.2

4.4 5.3

4.4 3.5

5.3 16.4

39.8 25.2

0.0 0.0

23.3 38.6 1.5 0.8 0.0 0.0 0.3 0.0 0.0 0.0 0.3 1.8 1.9 0.0 1.5 0.3 1.9 3.3 1.6 0.3 1.3 1.3

0.8 1.7 2.2 2.3 2.7 4.4 3.2 2.8 5.0 4.4 4.2 3.6 2.5 2.7 2.2 2.1 2.0 2.9 2.4 2.2 2.0 2.4 2.1 1.8 3.2

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

93.22 98.04 99.91 103.93 107.21 112.66 118.24 122.00 126.37 130.62 133.35 135.15 136.47 138.72 139.96 141.16 144.05 149.15 150.81 151.31 153.96 155.94 159.18 161.94 162.93 164.64 165.64

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

References Allibone, A., 1987. Koettlitz Group metasediments and orthogneisses from the mid Taylor Valley and Ferrar Glacier areas. N.Z. Antarct. Rec. 8, 48–60. Alonso, B., Anderson, J.B., Diaz, J.I., Bartek, L.R., 1992. Pliocene–Pleistocene seismic stratigraphy of the Ross Sea: evidence for multiple ice sheet grounding episodes. In: Elliot, D.H. (Ed.), Contributions to Antarctic Research, III. Antarct. Res. Ser. 57, 93–103. Anderson, J.B., Shipp, S.S., Bartek, L.R., Reid, D.E., 1992. Evidence for a grounded ice sheet on the Ross Sea continental shelf during the late Pleistocene and preliminary paleodrainage reconstructions. In: Elliot, D.H. (Ed.), Contributions to Antarctic Research, III. Antarct. Res. Ser. 57, 39–62. Barrett, P.J., Hambrey, M.J., 1992. Plio–Pleistocene sedimentation in Ferrar Fiord, Antarctica. Sedimentology 39, 109–123. Barrett, P.J., Scientific Staff, 1985. Plio–Pleistocene glacial sequence cored at CIROS-2, Ferrar Fjord, McMurdo Sound. N.Z. Antarct. Rec. 6, 8–19. Behrendt, J.C., Blankenship, D.D., Finn, C.A., Bell, R.E., Sweeney, R.E., Hodge, S.R., Brozena, J.M., 1994. Evidence for late Cenozoic flood basalts (?) in the West Antarctic rift system revealed by the CASERTZ Aeromagnetic Survey. Geology 22, 527–530. Behrendt, J.C., Blankenship, D.D., Damaske, D., Cooper, A.K., 1995. Glacial removal of late Cenozoic subglacially emplaced volcanic edifices by the West Antarctic ice sheet. Geology 23, 1111–1114. Boenigk, W., 1983. Schwermineralanalyse. Enke, Stuttgart, 158 pp. Boulton, G.S., 1978. Boulder shapes and grain-size distributions of debris as indicators of transport paths through a glacier and till genesis. Sedimentology 25, 773–799. Brotzu, P., Capaldi, G., Civetta, L., Melluso, L., Orsi, G., 1988. Jurassic Ferrar Dolerites and Kirkpatrick Basalts in Northern Victoria Land (Antarctica): stratigraphy, geochronology and petrology. Mem. Soc. Geol. Ital. 43, 97–116. Brotzu, P., Capaldi, G., Civetta, L., Orsi, G., Gallo, G., Melluso, L., 1992. Geochronology and geochemistry of Ferrar rocks from North Victoria Land, Antarctica. Eur. J. Mineral. 4, 605– 617. Cape Roberts Science Team, 1998. Initial Report on CRP-1, Cape Roberts Project, Antarctica. Terra Antart. 5 (1), 1–187. Cole, J.W., Ewart, A., 1968. Contributions to the volcanic geology of the Black Island, Brown Peninsula, and Cape Bird areas, McMurdo Sound, Antarctica. N.Z. J. Geol. Geophys. 11, 793–828. Denton, G.H., Bockheim, J.G., Wilson, S.C., Stuiver, M., 1989. Late Wisconsin and early Holocene glacial history, inner Ross embayment, Antarctica. Quat. Res. 31, 151–182. Diekmann, B., Kuhn, G., 1999. Provenance and dispersal of glacial-marine surface sediments in the Weddell Sea and adjoining areas, Antarctica: ice-rafting versus current transport. Mar. Geol. 158, 209–231. Ehrmann, W., 1998. Implications of late Eocene to early Miocene clay mineral assemblages in McMurdo Sound (Ross Sea,

243

Antarctica) on paleoclimate and ice dynamics. Palaeogeogr., Palaeoclimatol., Palaeoecol. 139, 213–231. Ehrmann, W., 1999. Lower Miocene and Quaternary clay mineral assemblages from CRP-1. Terra Antart. 5 (3), 613–619. Elliot, D.H., Siders, M.A., Haban, M.A., 1986. Jurassic tholeiites in the region of the upper Rennick Glacier, North Victoria Land. In: Stump, E. (Ed.), Geological Investigations in Northern Victoria Land. Antarct. Res. Ser. 46, 249–265. Elliot, D.H., Fleming, T.H., Haban, M.A., Siders, M.A., 1995. Petrology and mineralogy of the Kirkpatrick Basalt and Ferrar Dolerite, Mesa Range region, North Victoria Land, Antarctica. In: Elliot, D.H., Blaisdell, G.L. (Eds.), Contributions to Antarctic Research, IV. Antarct. Res. Ser. 67, 103–142. Findlay, R.H., Skinner, D.N.B., Craw, D., 1984. Lithostratigraphy and structure of the Koettlitz Group, McMurdo Sound, Antarctica. N.Z. J. Geol. Geophys. 27, 513–536. Fitzgerald, P.G., 1992. The Transantarctic Mountains of southern Victoria Land: the application of apatite fission track analysis to a rift shoulder uplift. Tectonophysics 11, 634–662. Fitzgerald, P.G., 1994. Thermochronology constraints on postPaleozoic tectonic evolution of the central Transantarctic Mountains. Tectonics 13, 818–836. Fleming, T.H., Heimann, A., Foland, K.A., Elliot, D.H., 1997. 40 Ar=39 Ar geochronology of Ferrar Dolerites from the Transantarctic Mountains, Antarctica: implications for the age and origin of the Ferrar magmatic province. Geol. Soc. Am. Bull. 109, 533–546. Ghent, E.D., Henderson, R.A., 1968. Geology of the Mt Falconer Pluton, lower Taylor Valley, South Victoria Land, Antarctica. N.Z. J. Geol. Geophys. 11, 851–879. Gleadow, A.J.W., Fitzgerald, P.G., 1987. Uplift history and structure of the Transantarctic Mountains and new evidence from fission track dating of basement apatites in the dry valleys area, southern Victoria Land. Earth Planet. Sci. Lett. 82, 1–14. Gunn, B.M., 1962. Differentiation in Ferrar Dolerites, Antarctica. N.Z. J. Geol. Geophys. 5, 820–863. Hambrey, M.J., Barrett, P.J., 1993. Cenozoic sedimentary and climatic record from Ross Sea region, Antarctica. In: Kennett, J.P., Warnke, D.A. (Eds.), The Antarctic Paleoenvironment: A Perspective on Global Change, Part Two. An Ocean Drilling Perspective. Antarct. Res. Ser. 60, 91–124. Harwood, D.M., 1986. Diatom Biostratigraphy and Paleoecology with a Cenozoic History of Antarctic Ice Sheets. Ph.D. Thesis, Ohio State University, Columbus, OH, 592 pp. Haskell, T.R., Kennett, J.P., Prebble, W.M., Smith, G., Willis, I.A.G., 1965. The geology of the middle and lower Taylor Valley of South Victoria Land, Antarctica. Trans. R. Soc. N. Z. 2, 169–186. Jakobsson, S.P., 1978. Environmental factors controlling the palagonitization of the Surtsey tephra, Iceland. Bull. Geol. Soc. Denm. 27, 91–105. Kellogg, T.B., Hughes, T., Kellogg, D.E., 1996. Late Pleistocene interactions of East and West Antarctic ice-flow regimes: evidence from the McMurdo Ice Shelf. J. Glaciol. 42, 486–500. Kyle, P.R., 1981. Glacial history of the McMurdo Sound area as indicated by the distribution and nature of McMurdo Volcanic

244

W. Ehrmann, K. Polozek / Sedimentary Geology 128 (1999) 223–244

Group rocks. In: McGinnis, L.D. (Ed.), Dry Valley Drilling Project. Antarct. Res. Ser. 33, 403–412. Kyle, P.R., 1990. McMurdo Volcanic Group, Western Ross Embayment. In: LeMasurier, W.E., Thomson, J.W. (Eds.), Volcanoes of the Antarctic Plate and Southern Oceans. Antarct. Res. Ser. 48, 19–25. Laird, M.G., Bradshaw, J.D., 1982. Uppermost Proterozoic and Lower Paleozoic geology of the Transantarctic Mountains. In: Craddock, C. (Ed.), Antarctic Geoscience. Univ. Wisconsin Press, Madison, pp. 525–533. La Prade, K.E., 1982. Petrology and petrography of the Beacon Supergroup, Shackleton Glacier area, Queen Maud Range, Transantarctic Mountains, Antarctica. In: Craddock, C. (Ed.), Antarctic Geoscience. Univ. of Wisconsin Press, Madison, pp. 581–590. LeMasurier, W.E., Thomson, J.W. (Eds.), 1990. Volcanoes of the Antarctic Plate and Southern Oceans. Antarct. Res. Ser. 48, 488 pp. Mange, M., Maurer, H.F.W., 1991. Schwerminerale in Farbe. Enke, Stuttgart, 148 pp. Morton, A.C., 1985. Heavy minerals in provenance studies. In: Zuffa, G.G. (Ed.), Provenance of Arenites. NATO ASI, Reidel, Dordrecht, pp. 249–277. Nathan, S., Schulte, F.J., 1968. Geology and petrology of the Campbell–Aviator Divide, Northern Victoria Land, Antarctica, Part 1. Post-Paleozoic rocks. N.Z. J. Geol. Geophys. 11, 940– 975. Pettijohn, F.J., Potter, P.E., Siever, R., 1987. Sand and Sandstone. Springer, New York, 2nd ed., 553 pp. Polozek, K., Ehrmann, W., 1999. Distribution of heavy minerals in CRP-1. Terra Antart. 5 (3), 633–638. Riddolls, B.W., Hancox, G.T., 1968. The geology of the upper Mariner Glacier region, North Victoria Land, Antarctica. N.Z. J. Geol. Geophys. 11, 881–899. Skinner, D.N.B., Ricker, J., 1968a. The Geology of the region between the Mawson and Priestley Glaciers, North Victoria Land, Antarctica, Part I. Basement meta-sedimentary and igneous rocks. N.Z. J. Geol. Geophys. 11, 1009–1040. Skinner, D.N.B., Ricker, J., 1968b. The Geology of the region between the Mawson and Priestley Glaciers, North Victoria Land, Antarctica, Part II. Upper Paleozoic to Quaternary geology. N.Z. J. Geol. Geophys. 11, 1041–1075. Smellie, J.L., Skilling, I.P., 1994. Products of subglacial volcanic eruptions under different ice thicknesses: two examples from

Antarctica. Sediment. Geol. 91, 115–129. Smillie, R.W., 1992. Suite subdivision and petrological evolution of granitoids from the Taylor Valley and Ferrar Glacier region, south Victoria Land. Antarct. Sci. 4, 71–87. Stuiver, M., Denton, G.H., Hughes, T.J., Fastook, J.L., 1981. History of the marine ice sheet in West Antarctica during the last glaciation: a working hypothesis. In: Denton, G.H., Hughes, T.J. (Eds.), The Last Great Ice Sheets. Wiley, New York, pp. 319–436. Tingey, R.J., 1991a. Commentary on schematic geological map of Antarctica Scale 1 : 10,000,000. Bureau of Mineral Resources, Australia, Bull. 238, 30 pp. Tingey, R.J., 1991b. Mesozoic tholeiitic igneous rocks in Antarctica — the Ferrar (Super) Group and related rocks. In: Tingey, R.J. (Ed.), Geology of Antarctica. Oxford University Press, Oxford, pp. 153–174. Warren, G., 1969. Geology of Terra Nova Bay–McMurdo Sound Area. Antarctic Map Folio Series 12. Am. Geogr. Soc., New York, Geology, Sheet 14. Weiblen, P.W., Stuckless, J.S., Hunter, W.C., Schulz, K.J., Mudrey, M.G., 1981. Correlation of clinopyroxene compositions with environment of formation based on data from Ross Island volcanic rocks. In: McGinnis, L.D. (Ed.), Dry Valley Drilling Project. Antarct. Res. Ser. 33, 229–246. Wilch, T.I., Lux, D.R., Denton, G.H., McIntosh, W.C., 1993. Minimal Piocene–Pleistocene uplift of the dry valleys sector of the Transantarctic Mountains: a key parameter in ice-sheet reconstructions. Geology 21, 841–844. Wimmenauer, W., 1985. Petrographie der magmatischen und metamorphen Gesteine. Enke, Stuttgart, 382 pp. Winter, D.M., Harwood, D.M., 1997. Integrated diatom biostratigraphy of late Neogene drillholes in southern Victoria Land and correlation to Southern Ocean records. In: Ricci, C.A. (Ed.), The Antarctic Region: Geological Evolution and Processes. Terra Antarctica Publications, Siena, pp. 985–992. Wohletz, K.H., Sheridan, M.F., 1983. Hydrovolcanic explosions, II. Evolution of basaltic tuff rings and tuff cones. Am. J. Sci. 283, 385–413. Young, D.J., Ryburn, R.J., 1968. The geology of Buckley and Darwin Nunataks, Beardmore Glacier, Ross dependency, Antarctica. N.Z. J. Geol. Geophys 11, 922–939. Zimmerle, W., 1972. Sind detritsche Zirkone ro¨tlicher Farbe auch in Mitteleuropa Indikatoren fu¨r pra¨kambrische Liefergebiete? Geol. Rundsch. 61, 116–139.