10. IMPLICATIONS OF SEDIMENT COMPOSITION ON THE …ite monochromator, automatic divergence slit, and...

Barron, J., Larsen, B., et al., 1991Proceedings of the Ocean Drilling Program, Scientific Results, Vol. 119

10. IMPLICATIONS OF SEDIMENT COMPOSITION ON THE SOUTHERNKERGUELEN PLATEAU FOR PALEOCLIMATE

AND DEPOSITIONAL ENVIRONMENT1

Werner U. Ehrmann2

ABSTRACT

Cretaceous to Quaternary sediments recovered at Leg 119 Sites 738 and 744 on the southern tip of the KerguelenPlateau were studied in order to determine the depositional environment and the paleoceanography of the southern In-dian Ocean and especially the long-term glacial history of East Antarctica. Emphasis is given to bulk-sediment compo-sition, grain-size data, and clay mineralogy.

The sediment sequence at the two sites is generally of a highly pelagic character, with nannofossil oozes, chalks, andlimestones dominant from the Turanian to upper Miocene and diatom oozes dominant within the uppermost Mioceneto Holocene interval.

The first indication of glaciation at sea level is the occurrence of isolated gravel and terrigenous sand grains, whichindicate ice rafting in the middle Eocene interval of 45.0-42.3 Ma. A major intensification of glaciation, probably theonset of continental East Antarctic glaciation, is recorded in sediments of early Oligocene age (36.0 Ma). All major sed-iment parameters document this event. The clay mineralogy changes from smectite-dominated assemblages, typical ofmoderately warm and humid climatic conditions in which chemical weathering processes are prevalent, to illite- andchlorite-dominated assemblages, indicative of cooler climates and physical weathering. Ice-rafted debris of both graveland sand size occurs in large quantities in that interval and coincides with a change in the mode of carbonate deposi-tion. Carbonate contents are relatively high and uniform (90%-95%) in strata younger than early Oligocene; in Oligo-cene to upper Miocene strata they fluctuate between 65% and 95%.

Oligocene and Neogene hiatuses reflect an intensification of oceanic circulation and the increased erosional force ofCircumpolar Deep Water. The long-term Cenozoic cooling trend was interrupted by a phase of early Miocene warmingindicated by maximum Neogene smectite concentrations. Although ice-rafted debris is present only in minor amountsand mainly in the silt fraction of early Oligocene to late Miocene age, it shows that glaciers advanced to the East Ant-arctic shoreline throughout that time. Ice-rafting activity drastically increased in latest Miocene time, when carbonatedeposition decreased and diatom ooze sedimentation started. This suggests a pronounced intensification of Antarcticglaciation combined with a northward migration of the Polar Front.

INTRODUCTION

The Southern Ocean around the Antarctic continent plays acritical role in controlling the global climate and oceanography,because the Antarctic coastal regions with their extensive iceshelves are the principal sources for southern oceanic bottomwaters, which influence the hydrography of large parts of theworld oceans. Changes in the plate tectonic configuration andin the intensity of glaciation in the southern high latitudes havehad great impact on the development not only of regions in theSouthern Hemisphere, but also of far-distant regions. Drillingin the seas around Antarctica consequently has been a main tar-get, both during the early days of the Deep Sea Drilling Project(DSDP Legs 28, 29, 35, and 36) and recently for the OceanDrilling Program (ODP Legs 113, 114, 119, and 120).

One of the principal objectives of Leg 119 was to study thelong-term glacial history of East Antarctica and its implicationsfor the paleoceanography and depositional environment of thesouthern Indian Ocean. For this purpose, two sites (736 and737) were drilled on northern Kerguelen Plateau, two sites (738and 744) on southern Kerguelen Plateau, two sites (745 and 746)in the Australian-Antarctic Basin, and five sites (739-743) inPrydz Bay on the East Antarctic continental shelf (Barron, Lar-sen, et al., 1989). The drilling results from the Prydz Bay sites,

1 Barron, J., Larsen, B., et al., 1991. Proc. ODP, Sci. Results, 119: CollegeStation, TX (Ocean Drilling Program).

2 Alfred Wegener Institute for Polar and Marine Research, Columbusstraße,D 2850 Bremerhaven, FRG.

which were expected to reveal the onset of continent-wide glaci-ation and to document in detail the Cenozoic glacial history ofEast Antarctica, suffered from poor recovery rates and poorstratigraphic control (Hambrey et al., 1989; this volume; Bar-ron, Larsen, et al., 1989). Therefore, this paper concentrates onthe pelagic sediments recovered at Sites 738 and 744 in thesouthernmost Indian Ocean (Fig. 1). The Mesozoic to Quater-nary sedimentary record is investigated in order to decipher, de-scribe, and interpret the depositional environment and to de-duce information concerning the onset of glaciation and thesubsequent Cenozoic glacial history of East Antarctica. Em-phasis is given in this paper to bulk-sediment composition,grain-size data, and clay mineralogy.

The two investigated sites are strategically positioned forhigh-latitude paleoceanographic and paleoclimatic studies onthe southern tip of the Kerguelen Plateau, north of the deep-wa-ter (>35OO m) pathway between Antarctica and the plateau.Site 738 was drilled at 62°42.54' S, 82°47.25' E, in 2253 m waterdepth; Site 744 at 61°34.66'S, 80°35.46'E, in 2307 m waterdepth (Fig. 1). The Kerguelen Plateau is about 2500 km longand 500 km wide, and is the world's largest submarine plateau.It stretches northwest-southeast from about 45 °S to 65 °S andrises 2-4 km above the adjacent Australian-Antarctic Basin tothe east, the Crozet Basin to the north, and the African-Austra-lian Basin to the west. The sites are situated about 1500 kmsouth of the Antarctic Convergence, or Polar Front. This sea-sonally fluctuating water mass boundary separates the coldAntarctic Surface Water in the south from the warmer Subant-arctic Surface Water to the north. Another important hydro-graphic feature, the Antarctic Divergence, is situated south of

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50°E 60c 70c 80c 90c40°S

50c

CrozetBasin

Crozet Island

4000—'Kerguelen Island

Polar Frontal Zone

Australian-Antarctic

BasinAfrican-Antarctic

Basin

ANTARCTICA

Measured at 55°S

• • • • • « Mean Summer Sea Ice Edge

• • Mean Winter Sea Ice Edge

Figure 1. Location map of Leg 119 Sites 738 and 744. Bathymetry (in meters) is from GEBCO (Hayes and Vogel, 1981; Fisher et al., 1982).The position of the Polar Frontal Zone is according to Whitworth (1988), the mean sea ice coverage in summer and winter time according toDietrich and Ulrich (1968).

186

SEDIMENT COMPOSITION ON THE SOUTHERN KERGUELEN PLATEAU

the two sites, at about 65°S. The Antarctic Divergence separatesthe eastward-flowing Antarctic Circumpolar Current (Whitworth,1988) to the north from the westward-flowing Antarctic CoastalCurrent to the south.

Site 738 yielded a 486-m-thick lower Turanian to Quaternarypelagic sequence that overlies volcaniclastic rocks and alteredbasalts (Barron, Larsen, et al., 1989). Biogenic calcareous sedi-ments dominate the Cretaceous and Paleogene section. Theygradually pass upward from limestone into calcareous chalk andnannofossil ooze. The limestone and chalk of Turanian to mid-dle Eocene age contain varying amounts of chert nodules. Thethin Neogene section is separated from the Paleogene section bya hiatus spanning early Oligocene to late Miocene time. The up-per Miocene to Holocene sediments are characterized by diatomooze with minor concentrations of foraminifers, calcareous nan-nofossils, and radiolarians. The recovery was almost 100% inthe upper 130 m of Hole 738B (Fig. 2) and decreased abruptlyto an average of about 45% with the occurrence of chert and thechange to rotary drilling in Hole 738C (Barron, Larsen, et al.,1989).

Because the Neogene section at Site 738 was abnormally thincompared with its seismically estimated thickness in nearbyareas, Site 744 was drilled in order to recover a thicker and morecomplete Neogene section. The drilling was successful and gaineda 176-m-thick upper Eocene to Quaternary sedimentary succes-sion (Fig. 3). Recovery was about 82% in Hole 744A and 100%in Holes 744B and 744C. The latter two holes were drilled to en-hance the recovery from Hole 744A and to provide more mate-rial for special scientific studies (Barron, Larsen, et al., 1989).The sediments recovered at Site 744 are highly pelagic and havea biogenic content of more than 90%. The upper Eocene to up-per Miocene sediments are characterized by nannofossil ooze,and the uppermost Miocene to Holocene sediments by soft dia-tom ooze.

METHODS

Bulk-sediment samples for this study were collected aboardJOIDES Resolution during Leg 119 and later at the ODP Re-pository at Lamont-Doherty Geological Observatory. Sampleswere taken, where possible, at 1.5-3-m intervals from Holes738B, 738C, and 744A. The uppermost 15 m from Site 738 wasnot sampled and analyzed because of a condensed and inter-rupted stratigraphic record and because this stratigraphic inter-val could be studied in more detail at Site 744. Sample size was10-20 cm3. Due to pressure release during the retrieval of sedi-ment cores from the great water depth, recovery sometimes ex-ceeded 100%. Depth information in this paper is corrected forcore recoveries >100%.

Carbonate and Organic Carbon

Total carbon content and inorganic carbon were determinedon freeze-dried and ground bulk subsamples using a Coulomat702 (Ströhlein Instruments). Gravel components and chert frag-ments were removed prior to grinding. Total carbon was set freeby heating the sample to 1100°C. Inorganic carbon was releasedby treatment with 14% phosphoric acid. Organic carbon andcarbonate (CaCO3) contents were calculated as weight percent.Available shipboard data (Barron, Larsen, et al., 1989) were in-corporated into the author's data set.

Grain Size and Biogenic Opal

Bulk samples were disaggregated by freeze-drying and subse-quent treatment with 10% H2O2. Carbonate was removed with10% acetic acid. The residue was washed through 2-mm and 63-µm sieves to separate the gravel and sand fractions. The clayfraction (<2 µm) was separated from the <63-µm fraction bydecantation (settling time based on Stokes' law). In order to

quantify the opal content and to remove the biogenic opalinecomponents from the terrigenous sedimentary particles, such asquartz, feldspar, and rock fragments, a density separation wasperformed on the sand and silt fractions. A sodium-polytung-state solution was used as a heavy liquid, with its density ad-justed to 2.3 g/cm3. The separation was carried out in a centri-fuge. Amorphous silica (mainly diatom debris) was removedfrom the <2-µm fraction quantitatively by a 3-min treatmentwith a boiling 2 molar Na2CO3 leach.

Clay Mineralogy

The clay mineral composition was analyzed by X-ray diffrac-tion (XRD) of the carbonate-free and opal-free <2-µm frac-tion. All samples were washed free of salt. The samples werefixed as texturally oriented aggregates on aluminium sampleholders and solvated with ethylene-glycol vapor at 60°C forabout 18 hr immediately before the analyses. The measurementswere conducted on an automated powder diffractometer systemPW 1700 (Philips) with CoKα radiation (40 kV, 40 mA), graph-ite monochromator, automatic divergence slit, and an auto-matic sample changer. The samples were X-rayed from 2° to 40°20 with a speed of 0.02° 20/s. The individual clay minerals wereidentified by their basal reflections at 15-18 Å (smectite), 10 Å(illite), 7 and 3.57 Å (kaolinite), and 14.2, 7, 4.72, and 3.54 Å(chlorite). Using empirically estimated weighting factors on in-tegrated peak areas of the individual clay mineral reflectionsfrom the glycolated samples, semiquantitative evaluations ofthe clay mineral assemblage were made (Biscaye, 1964, 1965;Lange, 1982). Because the weighting factors were calculated fora fixed slit, using an automatic slit resulted in a slight shift ofthe actual concentrations. However, because only relative changesin the concentration patterns of the individual clay minerals areconsidered in this paper, this inaccuracy seems acceptable.

Age Estimates

All age estimates given in this paper are based on the ship-board stratigraphy (Barron, Larsen, et al., 1989), diatom stra-tigraphy (Baldauf and Barron, this volume), the biostratigraphiosynthesis including the magnetostratigraphy of Site 744 (Barronet al., this volume), and the time scales by Berggren et al. (1985)and Aubry et al. (1988).

RESULTS

Bulk-Sediment Composition

Carbonate

Site 738

The carbonate concentration at Site 738 is generally veryhigh and over long core intervals shows only minor fluctuations(Fig. 2 and Appendix Tables 1 and 2). In the lower part of thesite, the lower Maestrichtian and deeper sediments, the carbon-ate concentrations are 80%-90%. In the upper Maestrichtian tolower Oligocene sediments they are even higher and range from90% to 98%. An abrupt decrease in carbonate content occurs inthe upper Miocene sediments overlying the lower Oligocene/up-per Miocene hiatus at 18 m below seafloor (mbsf). Concentra-tions are almost zero in the lower part of the Pliocene section,increase to 60%-80% in the upper Pliocene and lower Quater-nary section, and then again show a decreasing trend in themost recent sediments (Fig. 2).

Site 744

In the upper Eocene to lowermost Oligocene sediments ofSite 744, the carbonate content is nearly constant and exceeds90% of the bulk sediment (Fig. 3), revealing the same trend as

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o

Carbonate Opal(%)

Nonbiogenic matter

QJ O

0 20 40 60 80 100 0 10 20 30 40 50 0 10 20 30 40 50 60i i i i I i i i i i i i

3 0 0 -

3 5 0 -

400

4 5 0 -

500

Nannofossil ooze

Foraminifer ooze

Diatom ooze

Radiolarian ooze

Basalt A / V \ Hiatus

• • • Chert

Figure 2. Sediment composition at Holes 738B (Cores 1H through 24X) and 738C (Cores 4R through 33R). Car-bonate concentration is based on the author's and shipboard data (Barron, Larsen, et al., 1989).

188

f . 2 s>QJ F O

O

o

s 8 <uO <D O)o tr <


Carbonate Opal Nonbiogenic matter

0 20 40 60 80 100 0 10 20 30 40 50 0 10 20 30 40 50 60

20 -

100 -

120 -

140 -

160 -

Nannofossil ooze

Foraminifer ooze ΛΛΛΛ/

Diatom ooze Basalt / W \ Hiatus

Radiolarian ooze • A A Chert

Figure 3. Sediment composition at Hole 744A. Carbonate concentration is based on the author's and shipboard data(Barron, Larsen, et al., 1989).

the corresponding sediments at Site 738 (Fig. 2). The carbonateconcentration abruptly decreases to 60% at 146 mbsf (Fig. 3).This event has an age of 36.0 Ma, based on the shipboard stra-tigraphy (Barron, Larsen, et al., 1989) and on the studies ofBaldauf and Barron (this volume) and Barron et al. (this vol-ume).

Younger sediments were either not recovered at Site 738 orwere condensed. At Site 744, however, a detailed picture of thedevelopment of carbonate concentrations through Oligoceneand Neogene time can be drawn. After the event at 36.0 Ma thecarbonate concentrations slowly rise again and reach a constantlevel of 90% in the upper part of the lower Oligocene sediments.

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W. U. EHRMANN

This level is maintained through the upper Oligocene and lowerMiocene sediments. Two further weak minima in carbonatecontent occur in the Miocene. The first one comprises the lowermiddle Miocene; the second one the lower part of the upper Mi-ocene. In both cases concentrations fall to 70%-80%.

A distinct fall of carbonate concentrations was initiated inlatest Miocene time and was completed in the lower Pliocene in-terval, where sediments are almost devoid of carbonate (Fig. 3).Subsequently, concentrations rise to 20%-40% in the upper Pli-ocene and to 60% in the middle part of the Quaternary, beforethey decrease again in the youngest sediments. This Pliocene toQuaternary trend can also be seen in a condensed form at Site738 (Fig. 2).

Organic Carbon

The organic carbon content was determined on a few sam-ples from Site 744 only (Appendix Table 1) and is not presentedgraphically. These results and those of the shipboard analyses(Barron, Larsen, et al., 1989) show that the concentration of or-ganic carbon is generally <0.1% in sediments from both Sites738 and 744.

Opal

Site 738

The opaline components at Site 738 consist exclusively of ra-diolarians in the pre-Oligocene lower part of the site and ofmixed diatoms and radiolarians in the upper part. Silicoflagel-lates and sponge spicules also occur, but are quantitatively un-important in the upper part. Several small and isolated maximain the opal concentrations of the bulk sediment are present inthe Cretaceous and Paleogene sediments of Site 738 (Fig. 2 andAppendix Table 1). The concentrations, however, do not exceed7% of the bulk sediment. Most of the maxima correlate with re-ported occurrences of radiolarians (Barron, Larsen, et al.,1989). The most prominent maxima occur in the upper Campa-nian to lower Maestrichtian and in the middle part of the mid-dle Eocene.

Some opaline material is also present in the upper Eocenesediments, and a slight increase in concentrations to 5% occursjust below the Eocene/Oligocene boundary. The latter cannotbe ascribed to radiolarians alone, but also to the small numberof poorly preserved diatoms reported in this interval (Barron,Larsen, et al., 1989). A pronounced increase in opal content,mainly from diatoms, is documented in the upper part of Core119-738B-3H (Fig. 2). It is difficult to date precisely the onset ofdiatom ooze sedimentation, because it occurs within a 2-m-thick core interval containing three hiatuses and probably a con-densed section between the hiatuses (Barron, Larsen, et al.,1989, p. 246). Pure nannofossil ooze accumulation ceased at ahiatus spanning approximately 35-10 Ma. A transition to morediatom-rich sediments occurred between 10 and 8.4 Ma, whenanother hiatus interrupted the sedimentary record. Above thishiatus, at 6.1 Ma, diatom ooze accumulation with opal concen-trations of up to 90% had started, but was interrupted by a fur-ther hiatus spanning >5.8-<4.8 Ma (Barron, Larsen, et al.,1989).

Site 744The opaline material of sediments from Site 744 is composed

mainly of diatoms and, to a minor degree, radiolarians. Silico-flagellates and sponge spicules occur in small quantities. Opalat Site 744 first appears in detectable amounts of about 4% at167 mbsf. The onset of accumulation of biosiliceous compo-nents, even if in small amounts, therefore has a late Eocene ageof 38.4 Ma. Shipboard paleontologists found the oldest diatomsand radiolarians, which were few in number and poorly pre-

served, in the lower Oligocene at 158 mbsf (Barron, Larsen, etal., 1989). Both groups of siliceous microfossils are then consis-tently present and moderately to well preserved throughout theOligocene to Quaternary sediments (Fig. 3 and Appendix Table1).

The first maximum with 5%-20% opaline matter at Site 744occurs in the lower Oligocene (36.0 Ma) and correlates with adiminished carbonate concentration (Fig. 3). In the upper partof the lower Oligocene and lower Miocene, opal concentrationsare low again, at about 5% of the bulk sediment consisting ofdiatoms and radiolarians. In the middle and upper Miocene sec-tion the opal content is similar to that in the lower Oligocene,but fluctuates widely between 5% and 20%. Greatly increasedopal contents occur at Hole 744A above the upper Miocene/Pli-ocene hiatus between Cores 119-744A-4H and 119-744A-3H.Relatively poor recovery rates in this part of Hole 744A hamperan exact dating. In Hole 744B, which had a 100% recovery, thetransition from nannofossil ooze to diatom ooze occurs above ahiatus dated 9.0-6.0 Ma and just below another hiatus dated5.5-4.0 Ma. In Hole 744C the transition occurs approximately 1m below the Miocene/Pliocene boundary, but cannot be datedmore accurately because of the lack of detailed biostratigraphicstudies. The onset of diatom ooze accumulation at Site 744 thusstarted in latest Miocene time, at 5.8-5.6 Ma.

The Pliocene to Holocene sediments are characterized byopal concentrations between 25% and 60%. Estimates based onsmear slide analyses indicate even higher values of up to 80%(Barron, Larsen, et al., 1989).

Nonbiogenic Matter

Site 738Sediments recovered at Site 738 contain approximately 5%

nonbiogenic matter (Fig. 2 and Appendix Table 1). Only two in-tervals are characterized by higher values. Concentrations of upto 13% are recorded in the Turonian to lower Maestrichtian sed-iments, and the nonbiogenic matter consists mainly of clay min-erals (Fig. 6). Concentrations of up to at least 50% characterizethe upper Miocene to Quaternary interval.

Site 744The standard level of 5% nonbiogenic matter at Site 738 is

also found in the upper Eocene to lowermost Oligocene sedi-ments of Site 744 (Fig. 3 and Appendix Table 1). A sharp peak(27%) in the concentration of nonbiogenic matter obviouslymatches the drop in carbonate concentration and the increase inopal content in the lower Oligocene at 36.0 Ma. Concentrationstend to be about 8% in the remaining Oligocene section and inthe lower Miocene sediments, about 10%-15% in the middleMiocene, and about 5% in the upper Miocene. The increase inthe uppermost Miocene and Pliocene to 40% of the bulk sedi-ment correlates with changes in concentration of carbonate andopal. A clear maximum occurs in the upper lower to lower up-per Pliocene interval.

Grain-Size Distribution of Nonbiogenic Matter

Results from Core Descriptions

Site 738The lowermost coarse terrigenous material detected at Site

738 during core description is an angular granite clast in Section119-738B-8H-1. This gravel occurs in sediment that is about40.5 Ma old, but it probably can be ascribed to contaminationdue to hole cave-in (Barron, Larsen, et al., 1989). The same istrue for gravel found along the core liner of Section 119-738B-7H-1 (Fig. 4).

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From Section 119-738B-3H-5, 35 cm, to the top of Hole738B, coarse terrigenous particles are scattered throughout thesediments. The oldest ice-rafted debris belongs to the early Oli-gocene foraminifer Zones PI8-20 and diatom Zones CP15b-16(Barron, Larsen, et al., 1989). Thus, the inferred maximum ageis about 36 Ma. The Neogene ice-rafted debris is composed ofgranitic and gneissic clasts several centimeters in diameter, aswell as of pebble-, granule-, and sand-sized grains consistingmainly of quartz. Most grains are coated with manganese oxideand are angular to subangular (Barron, Larsen, et al., 1989).

Site 744

The lowermost gravel and coarse sand grains visible in a cutcore surface from Site 744 were found scattered through the in-terval from Sections 119-744A-20H-1, 0 cm, to 119-744A-20H-2, 57 cm (Fig. 5). Biostratigraphic data indicate a late Eoceneage of about 38.5 Ma (Barron et al., this volume). This intervalof the core, however, is very disturbed by drilling, and the recov-ered nannofossil ooze has a soupy character. The grains float inthe host sediment mass. The lowermost granule was found inthe transition zone from soupy to firm sediment. The firmerpart of Section 119-744A-20H-2 was X-rayed aboard ship, butno coarse sediment grains could be detected. All coarse materialin this interval is therefore thought to be due to downhole con-tamination.

The next appearance of coarse grains is in Core 119-744A-17H (Fig. 5), which has a core length of only 0.90 m and a re-covery of 0.83 m. The alignment of most of the granules alongthe core liner suggests an origin by downhole contamination.Some other grains, however, occur within a firm host sedimentand are in situ (Barron, Larsen, et al., 1989). One of the ice-rafted grains is 2 cm in diameter and consists of a manganeseoxide-coated gneiss. The other grains comprise numerous dis-seminated quartz granules of 1-4 mm diameter and some smallmanganese oxide-coated pebbles. According to the magneto-stratigraphic interpretation (Barron et al., this volume), Core119-744A-17H is situated just beneath magnetic anomaly 13(35.9-35.3 Ma; Berggren et al., 1985; Aubry et al., 1988). Thisimplies an age of about 36.2-35.9 Ma for the core.

Also found in the lower Oligocene interval from Sections119-744A-16H-7 to 119-744A-16H-CC was a large quantity ofunequivocally ice-rafted granules. The grains are angular andare mainly 1-4 mm in diameter. This occurrence is at the base ofmagnetic anomaly 13, and therefore is 35.9 Ma old. The pulseof early Oligocene ice rafting is assumed to be contemporane-ous with that at Site 738. An exact correlation, however, is notpossible because of the lack of paleomagnetic data from Site738.

In the middle part of the core recovery from Site 744 no ice-rafted pebbles and granules were detected in the cores. The nextones occur in the upper Miocene to Quaternary sediments ofCore 119-744A-1H to Section 119-744A-4H-3, 100 cm (27.2mbsf; late Miocene, 9 Ma). The granules are mostly subangular,and some of the larger pebbles (up to 5 cm in diameter) are sub-rounded. Most of the clasts, which include quartz, granite,gneiss, and amphibolite, are coated by manganese oxide.

Results from Grain-Size Analyses

Site 738

The deepest occurrence of gravel-sized components in sam-ples from Site 738 was in the lower Paleocene at 352 mbsf (Fig. 4and Appendix Table 3). The gravel consists of chert fragmentsprobably created during rotary drilling and transferred down-core. Two samples from 104 and 85 mbsf contained one suban-gular gneissic or granitic rock fragment each. Both samples arefrom core sections that did not show any indications for con-

tamination. Stratigraphically they belong to the middle Eocene(Barron et al., this volume); their absolute age is 43.5-42.3 Ma.The next and youngest gravel occurs in the uppermost sampleinvestigated, in Section 119-738B-3H-1, and has an early Plio-cene age.

The sand-sized nonbiogenic material has a concentrationpattern characterized by numerous, but mainly isolated, peaks(Fig. 4 and Appendix Table 3). Most of the samples, however,did not contain measurable amounts of nonbiogenic sand. Theenhanced sand occurrence in the Paleocene cores is caused bychert fragments generated during rotary drilling and probablysmeared along the liner during coring and along the cut surfaceduring splitting of the cores. The small maxima in Cores 119-738C-7R and 119-738B-16X are due to authigenic minerals. Thelowermost detrital sand grains are found at a depth of 131 mbsf,i.e., deeper than the first terrigenous gravel component. Theyconsist mainly of angular to subangular quartz grains and somefeldspar. According to the biostratigraphy and the derived sedi-mentation rates (Barron et al., this volume), the core intervalhas a middle Eocene age of about 45.0 Ma. The next terrige-nous sand grains occur at 109 and 104 mbsf. These grains com-prise subangular igneous rock fragments and a subrounded red-dish sandstone fragment and have a middle Eocene age of about44-43.5 Ma. In the lower Oligocene, at 20 mbsf, subangularquartz grains are the dominant constituent of the nonbiogenicsand fraction. The highest concentration of the sand-sized frac-tion of the nonbiogenic components of Site 738 was found inthe sample from lower Pliocene Section 119-738B-3H-1.

The silt grain sizes of the nonbiogenic matter (Fig. 4 and Ap-pendix Table 3) in the lower part of Site 738 are controlled bydrilling breccias of chert nodules. In the upper part of the sitethree maxima can be identified, with the first one in the middlepart of the middle Eocene, between 180 and 140 mbsf. The sec-ond maximum comprises the youngest interval of the middleEocene at 90-70 mbsf, and the third one characterizes the top ofthe site and starts in the lower Oligocene sediments.

The clay fraction at Site 738 is the dominant grain-size frac-tion of the nonbiogenic material. Its concentration pattern isopposite that of nonbiogenic sand and silt (Fig. 4 and AppendixTable 3).

Site 744

The deepest occurrence of gravel in samples from Hole 744Ais in Core 119-744A-20H (Fig. 5 and Appendix Table 3). Ac-cording to the preceding discussion, this is due to contamina-tion. The next gravel was found in Samples 119-744A-17H-CC,12-14 cm, and 119-744A-16H-7, 47-49 cm, which have an earlyOligocene age of about 36 Ma. Higher in the core, only Sample119-744A-3H-1, 49-51 cm, which has a Pliocene age, containedgravel.

The sand-sized nonbiogenic material (Fig. 5 and AppendixTable 3) in the Eocene and early Oligocene age sediments showsthe same concentration pattern as the gravel component. At theearly Oligocene peak occurrence, up to 38% of the nonbiogenicmaterial is sand sized. The remainder of the Oligocene and theMiocene part of the sequence recovered at Hole 744A are char-acterized by very low terrigenous sand concentrations. A slightincrease is documented at 26 mbsf (9.0 Ma), but cannot be fol-lowed into younger sediments because of a hiatus. A dramaticincrease in sand-grain size was found in Section 119-744A-4H-1(24 mbsf). which represents the uppermost Miocene, and has anage range of about 5.7-5.8 Ma. From the uppermost Mioceneto Holocene, 10% to 30% of the nonbiogenic sediment compo-nents occurs in the sand fraction.

The nonbiogenic material does not reveal very distinct max-ima and minima in its silt concentration pattern (Fig. 5 and Ap-pendix Table 3). However, after very low values in the Eocene in-

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W. U. EHRMANN

Gravel Sand Silt Clay(% of nonbiogenic component)

0 2 4 6 8 10 0 10 20 30 0 10 20 30 40 0 20 40 60 80 100

500

Nannofossil ooze Diatom ooze Basalt / W \ Hiatus

Foraminifer ooze f * * * * d Radiolarian ooze • A A Chert

Figure 4. Gravel content and grain-size distribution of the nonbiogenic sediment components at Holes 738B (Cores1H through 24X) and 738C (Cores 4R through 33R). Asterisks indicate occurrences of gravel according to visualcore description (Barron, Larsen, et al , 1989). Gravel concentration is the number of pebbles and granules in theinvestigated samples.

192

*Eü£

α>

S i

100 -

120 -

140 -

160 -

o•B


Gravel Sand Silt Clay

(% of nonbiogenic component)

0 2 4 6 8 0 10 20 30 0 10 20 30 40 0 20 40 60 80 100

^ I V x j Nannofossil ooze

Foraminifer ooze

• y y y yDiatom ooze

Radiolarian ooze

Basalt Λ A A Hiatus

A A • Chert

Figure 5. Gravel content and grain-size distribution of the nonbiogenic sediment components at Hole 744A. Aster-isks indicate occurrences of gravel according to visual core description (Barron, Larsen, et al., 1989). Gravel concen-tration is the number of pebbles and granules in the investigated samples.

terval, there is a slightly increasing trend from the lowermostOligocene to the lowermost Miocene (155-93 mbsf). The con-centrations rise upcore from about 5% to 20%. They then de-crease to the base of Core 119-744A-8H. The next maximum innonbiogenic silt-sized material occurs in the lower part of themiddle Miocene interval. In younger sediments the concentra-

tions are lower, with minimum values in the upper middle andlower upper Miocene.

The clay fraction is the dominant grain-size fraction of thenonbiogenic material at Hole 744A. Its concentration pattern isopposite that of the nonbiogenic sand and silt (Fig. 5 and Ap-pendix Table 3).

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Clay Minerals

Site 738

At Site 738 most of the nonbiogenic sediment is in the < 2-µm fraction. Therefore, there is a strong correlation between theconcentration patterns of nonbiogenic matter and clay in thebulk sediment (Figs. 2 and 6).

The clay fractions of the Cretaceous, Paleocene, and Eocenesediments (470-40 mbsf) are dominated by smectites (Fig. 6 andAppendix Table 4). In addition, between about 130 and 470mbsf some randomly interstratified mixed-layer illite-smectitesoccur. The smectite share (including the mixed layers) of theclay mineral assemblage in this lower part at Site 738 is generally>90%. Only at 470-455 and 410-380 mbsf do concentrationsdrop to 7O%-85%. A distinct fall in smectite concentrationstarts in the upper Eocene interval at 42 mbsf and reaches a firstminimum of 59% at 36 mbsf in the uppermost Eocene (38 Ma).Only 4 m higher in the core, the next maximum with 97% smec-tite is reached in the lowermost Oligocene sample, followed bydecreasing values.

Illite concentrations show a strong negative correlation withthe smectite abundance. Chlorites occur at Site 738 in the pre-Oligocene sediments below 40 mbsf, but only in traces. An ex-ception is the interval from 150 to 130 mbsf, which has chloriteconcentrations of 5%-10%. In the upper part of the upper Eo-cene interval, at 40 mbsf, chlorite rises to 15%. After a mini-mum at 32 mbsf the values are > 10% (Fig. 6). The kaolinitecontent shows a pattern similar to that of chlorite, with traceamounts in the deeper part of the site, a small maximum of5%-7% at 150-130 mbsf, and a clear increase to values ofabout 6%-10% at 40 mbsf.

Site 744

The percentage of clay minerals in the bulk sediment of Hole744A is relatively low and reflects the highly pelagic character ofthe depositional environment on southern Kerguelen Plateau(Fig. 7 and Appendix Table 4). From the bottom of the hole at176 mbsf to about 24 mbsf the clay concentrations generallyrange from 3% to 10%. Fluctuations are only minor, but slightmaxima occur at 146 mbsf (early Oligocene, 36 Ma) and at 52-57 mbsf within the middle Miocene. A distinct increase in claycontent can be observed in the uppermost upper Miocene inter-val at 24 mbsf. A maximum of 26% is reached at the boundarybetween the lower and upper Pliocene at 19 mbsf. Generally de-creasing values characterize the overlying sedimentary sequence.

The smectite concentrations in the lower part of Hole 744Aare relatively high (Fig. 7 and Appendix Table 4). Between thebottom of the hole and 150 mbsf, in the upper Eocene and low-ermost Oligocene sediments, they follow a slightly decreasingtrend from about 90% to 70%. A short-lived minimum indi-cated by one sample from Section 119-744A-19H-1 (late Eo-cene, 37.5 Ma) possibly correlates with a short minimum at Site738 (38.0 Ma). A sharp and almost continuous drop to a mini-mum value of 16% occurs in lower Oligocene sediments be-tween 150 and 142 mbsf (36.5-35.5 Ma), and also correlateswith Site 738. The interval 142-95 mbsf is characterized bysmectite concentrations fluctuating at about 40%. A short max-imum of about 50%-60% follows in the lower half of the lowerMiocene interval. Values of about 40% occur again from 81mbsf to the surface, with a slight and indistinct maximum inupper Miocene sediments between 37 and 18 mbsf.

Smectite, together with illite, accounts for 70%-90% of theclay minerals present in the Hole 744A sediments. They show astrong negative correlation (r = —0.93). Changes in the illiteconcentrations are therefore not discussed here. Chlorite con-tents are low (<4 %) in the deepest 10 m at Hole 744A. An in-crease to about 5%-10% occurs at 165 mbsf, and a maximum

of 20% is reached between 147 and 142 mbsf. Further up in thecore the chlorite concentrations fluctuate between 5% and 15%without any significant changes in the distribution pattern (Fig.7). The kaolinite occurrence is restricted to trace amounts in thedeepest part of Hole 744A. The first significant concentrationsappear in the upper Eocene at 165 mbsf. A maximum concen-tration of 26% is found in the lower Oligocene at 146 mbsf. Up-core, the concentrations fluctuate between mainly 2% and 12%with no particularly obvious pattern.

DISCUSSION

Environmental Implications of the SedimentComposition

The latest Miocene to Quaternary sediments at the southernKerguelen Plateau Sites 738 and 744 consist predominantly ofdiatom ooze with a variable carbonate content (Figs. 2 and 3).The older sediments are nannofossil ooze, chalk, and limestonewith a minor biosiliceous and almost negligible nonbiogeniccontent. The opaline components essentially disappear belowthe Eocene/Oligocene boundary, but their former presence andtheir diagenetic alteration are documented by the occurrence ofporcellanite and chert nodules in middle Eocene and older sedi-ments.

The sediments of Site 738 roughly document the subsidencehistory of the southern Kerguelen Plateau (Barron, Larsen, etal., 1989). Algal calciclastic limestones were deposited in a shal-low-water inner shelf setting during Turonian time. A slightdeepening equivalent in depth to a shelf break setting by the endof the Cretaceous Period is indicated by the benthic shellyfauna, trace fossil assemblages, and omission surfaces associ-ated with glauconite. Deepening continued, and deposition ofpelagic oozes took place at or near the present-day water depththroughout most of the Cenozoic Era.

The presence of at least small amounts of carbonate at Sites738 and 744 (Figs. 2 and 3) indicates that the two sites have re-mained above the carbonate compensation depth (CCD) sinceLate Cretaceous time. With the exception of a short interval inearly Pliocene time, the sites also remained above the carbonatecritical depth (CCrD, <10% CaCO3; Kolla et al., 1976). TheCCD in the present-day Indian Ocean is situated in about 5200 mwater depth (Van Andel, 1975), and the CCrD at about 3900 mwater depth (Kolla et al., 1976).

Although the carbonate percentages are very high (>90%)and constant in the Eocene and older sediments, they fluctuateslightly from the lower Oligocene to the upper Miocene andstrongly from the uppermost Miocene to the Holocene (Figs. 2and 3). The transition from consistently high to slightly lowerand fluctuating concentrations occurs abruptly at the base ofCore 119-744A-16H, dated as early Oligocene at about 36 Ma.Compared with older and younger pure calcareous sediments,samples from Core 119-744A-16H are characterized by a muchlower number of planktonic foraminifers and also by a benthicrelict fauna (A. Mackensen, pers. comm., 1989). The reducedcarbonate content is thought therefore to be caused, at least inpart, by dissolution processes. The decrease in carbonate con-centration at 36 Ma correlates with a sharp peak in the concen-trations of nonbiogenic matter of ice-rafted origin and with afirst maximum in opal concentrations (Fig. 3).

Moderate to weak minima in the carbonate content occur atSite 744 in the lower Oligocene (146-130 mbsf; 36-34 Ma), atthe transition between the lower and the middle Miocene (70-60mbsf; 17.5-15.0 Ma), in the middle Miocene (57-48 mbsf; 14-11 Ma), and in the lower upper Miocene (40-35 mbsf; 10.5-10.0Ma). The minima coincide with times of major oxygen isotopeincreases signaling the cooling of Southern Ocean waters (Shack-leton and Kennett, 1975). The cooling could have caused in-

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creased biosiliceous fertility (Barron, Larsen, et al., 1989), andthus a dilution of the calcareous sediment components, or in-creased aggressivity of the bottom water masses and resultingcarbonate dissolution, as indicated for the lower Oligocene sam-ples. At Leg 120 Site 751, approximately 400 km north of Site744, the minima in carbonate concentration are at least in partdue to dissolution (A. Mackensen, pers. comm., 1989). A high-resolution study of the Miocene to recent carbonate content,benthic foraminifer assemblages, and δ 1 8 θ and δ13C isotopicdata will be presented for that site (Mackensen et al., in press).

A transition from almost pure nannofossil ooze to predomi-nantly diatom ooze sedimentation took place in latest Miocenetime (Figs. 2 and 3). On the southernmost Kerguelen Plateau, atSite 738, this transition is dated at >6.1 Ma, and at Site 744 at5.8 Ma. At Site 751 pure diatom ooze started to accumulateabove a hiatus spanning the time interval 5.9-4.7 Ma (Macken-sen et al., in press). It thus seems that the facies change gradu-ally moved northward with time and possibly reflects the north-ward migration of the Polar Front. On the northern KerguelenPlateau, at Site 737, diatom ooze sedimentation started earlier,at about 7 Ma. However, temperate low-latitude species makeup an important part of the diatom assemblages there. A transi-tion to assemblages dominated by Southern Ocean species oc-curred at 3.9 Ma at Site 737 (Barron, Larsen, et al., 1989). Re-placement of a carbonate facies by an almost completely biosili-ceous facies was also recorded in uppermost Miocene sedimentsrecovered at 65 °S on Maud Rise during ODP Leg 113 (Barker,Kennett, et al., 1988).

The decrease in carbonate concentrations in the lower Plio-cene to values < 5 % also can be explained by a drastic shoalingof the CCrD and the CCD. A cooling of the surface waters andthe extension of ice shelves enhanced the production of coldbottom waters, which in turn caused an intensified bottom wa-ter circulation resulting in widespread hiatuses (Davies et al.,1975; Barron and Keller, 1982; Ciesielski et al., 1982). Shoalingof the lysocline and the CCD during the latest Miocene was re-ported for the South Atlantic (Hsü et al., 1984), but so far notfor the southern Indian Ocean. The shallowest Cenozoic level ofthe lysocline and CCD, however, was reported for the middleMiocene.

Ice Rafting on Southern Kerguelen PlateauThe temporal variation in concentration of ice-rafted debris

is the most obvious indicator for strong climatic changes in andaround Antarctica through time. In this paper, ice-rafted debrisis used to date the onset of glaciation at sea level in East Antarc-tica and to reconstruct the Cenozoic glacial history.

Ice-rafted debris, such as that recovered on southern Ker-guelen Plateau, is delivered to the distal marine environmentmainly by icebergs calving from ice streams and outlet glaciersinto the sea (Drewry, 1986). Coarse surface sediments can alsobe picked up on the continental shelf by drifting icebergs, whichperiodically become grounded and then transport the sedimentinto the adjacent seas (Kellogg and Kellogg, 1988).

In contrast, many icebergs calving from larger ice shelvesconsist of clean ice, because of melting at the base of the iceshelf. At the base of an ice shelf most of the terrigenous mate-rial melts out and settles on the shelf area soon after leaving thegrounding line (Kellogg and Kellogg, 1988; Hambrey et al., thisvolume). Debris-laden icebergs calving from ice shelves can beobserved when saline ice freezes onto the base of the ice shelfand thereby protects the debris from early release by meltingprocesses, as is happening under the Amery Ice Shelf (Hambrey,this volume). The debris well inside the body of a calving ice-berg may be released into the sea only when the berg disinte-grates, a long way from the source. Another exception occurswhen a grounded ice sheet, with sediment incorporated in its

base, decouples from its bed, becomes unstable, and breaks intopieces, which then drift away.

Although ice-rafted sediment has all grain-size classes repre-sented in roughly equal parts, most studies have used only theconcentration of terrigenous sand or gravel as a measure of theice-rafting intensity (e.g., Blank and Margolis, 1975; Bornhold,1983; Grobe, 1986). The clay and silt fractions of sediments arenot commonly used as an indicator of ice rafting, because thismaterial could also originate from other sources, such as windtransport, turbidity currents, and bottom currents and becausethey are prone to postdepositional redistribution by current ac-tivity. However, in the highly pelagic setting of the southern Ker-guelen Plateau, an origin for this material other than by ice raft-ing is not likely. Turbidites have not been observed in the sedi-ments drilled at Sites 738 and 744. They are not likely to occurin the morphologic setting presented by these sites, which are el-evated more than 1000 m above the deep-water passage that sep-arates Antarctica from the Kerguelen Plateau (Fig. 1). Windtransport was also probably much reduced when the Antarcticcontinent was glacierized. Transport by bottom currents isthought to have played only a minor role at the two sites, be-cause through all the time represented by sediments, fine-grainedsediment particles continuously accumulated in large amounts.A few hiatuses, however, mark erosive phases.

Origin of the Ice-Rafted Debris

Almost all of the gravel- and sand-sized rock fragments re-covered from sediments at Sites 738 and 744 are of gneissic and,to a minor degree, granitic character. Even the terrigenous sandand silt, which mainly consist of quartz and feldspar, do not in-dicate a different source. The only major exception is a largegrain of reddish sandstone, which was recovered from Section119-738B-13H-1. The ice-rafted debris resembles the materialfound in Prydz Bay (Hambrey et al., this volume) and thatfound just east of the Kerguelen Plateau at Sites 745 and 746,which are situated in about 4000 m water depth in the Austra-lian-Antarctic Basin (Ehrmann and Grobe, this volume; Ehr-mann et al., this volume).

The clast composition is similar to that expected from asource area situated in East Antarctica. The geology of this partof the continent is characterized by widespread occurrences ofPrecambrian gneisses and migmatites, as well as some char-nockites and schists (e.g., Craddock, 1982; Ravich and Fedorov,1982; Trail and McLeod, 1969). The hard sandstone clast maybe derived from an equivalent of Beacon Supergroup strata,which are widespread in various parts of Antarctica. The smallamount of sedimentary rock fragments recovered may mirrorthe areal distribution of sedimentary as opposed to igneous andmetamorphic source rocks in East Antarctica. It may also dem-onstrate that the sedimentary rocks were not sufficiently lithi-fied to survive glacial transport, but were disaggregated intotheir individual mineral components.

The Kerguelen Plateau, with its widespread volcanism andislands, is not a likely source for the sediment particles depos-ited at Sites 744 and 738, which lie on the southern tip of theplateau. Sedimentary calcareous clasts and basaltic components,which could be derived from the plateau, are totally missing.

Ice Rafting Through Time

The amount of ice-rafted debris at the sites is controlled bythe source area of the icebergs, the iceberg frequency, and bythe calving processes. These, in turn, are controlled by sea-levelchanges and the mass budget of the Antarctic ice sheet in re-sponse to climatic changes. Furthermore, the oceanography ofthe Southern Ocean plays an important role in the distributionof icebergs and is a crucial limiting factor in reconstructing theglacial history and the paleoclimate based on the record of ice-

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2 8 ©O Φ D5ü It <

Smectite Illite Chlorite Kaoliniteë> Clay| (%) (%) (%) (%) (%)

0 10 20 0 20 40 60 80 0 20 40 60 0 10 20 0 10 20 30i i | i i i i i I i i i i i i i i i i i i I i i i i i

400

500 -4

' ^ ^ ^ | Nannofossil ooze

Foraminifer ooze

Diatom ooze

Radiolarian ooze

Basalt A A Λ Hiatus

A A • Chert

Figure 6. Amount of terrigenous clay (nonbiogenic, <2 µm) and relative percentage of the individual clay minerals atHoles 738B (Cores 1H through 24X) and 738C (Cores 4R through 33R).

196


nj Ow. O Φo © σ>ü DC <

oClay Smectite Illite Chlorite Kaolinite

0 10 20 0 20 40 60 80 0 20 40 60 0 10 20 0 10 20 30

80

100 -

120 -

140 -

160 ~

'&£$&• Nannofossil ooze

v r ^ ^ g j Foraminifer ooze

Diatom ooze Basalt Λ A A Hiatus

Radiolarian ooze A • A Chert

Figure 7. Amount of terrigenous clay (nonbiogenic, < 2 µm) and relative percentage of the individual clay minerals atHole 744A.

rafted debris. Unfortunately, little is known concerning presentand past oceanic circulation systems in the southern IndianOcean.

At Sites 738 and 744 the intensity of ice rafting, as reflectedin the content of gravel and terrigenous sand, strongly fluctu-ated through time, but correlates between the two sites (Figs. 4

and 5). The first occurrence of coarse terrigenous material, bothgravel and sand sized, is recorded from middle Eocene pelagicsediments at Site 738 (Fig. 4), which have an absolute age of 45-42.3 Ma. It is assumed that this material was transported by ice-bergs, which implies that glaciers already had reached the sea bythat time. The absolute amount of ice-rafted debris, however, is

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W. U. EHRMANN

low, compared with present-day conditions. Continental glacia-tion is therefore not very likely. Some isolated glaciers probablyreached the sea while most of the continent was unglacierized.

The next pulse of input of ice-rafted debris that occurred si-multaneously at Sites 738 and 744 (Figs. 4 and 5) marks an earlyOligocene time interval at about 36 Ma. This pulse is muchmore pronounced than that of middle Eocene time. It correlateswith a change in carbonate deposition and clay mineralogy and,therefore, is thought to represent a major intensification of EastAntarctic glaciation. Whether this intensification represents amajor mountain glaciation with glaciers calving at sea level orthe onset of continental glaciation cannot be concluded fromthe record of ice-rafted debris alone, because of the reasons pre-viously discussed. A joint discussion of all sediment parametersfollows.

In late Oligocene to late Miocene time, ice rafting seems tohave played a minor role. Gravel is totally absent at Sites 738and 744, and terrigenous sand is present in trace amounts only(Figs. 4 and 5). Also, the concentration of terrigenous silt is rel-atively low, but shows slight maxima at the Oligocene/Miocenetransition and in the lower part of the middle Miocene. The de-crease in coarse ice-rafted debris may be due to less intense gla-cial activity, to a reduced availability of detrital material afterthe earliest glaciers had removed all the loose weathering prod-ucts, or to oceanographic changes.

An intensification of ice-rafting activity occurred in late Mi-ocene time at about 9 Ma. Since that time, ice-rafted gravel andsand occur in large amounts at both Sites 738 and 744 (Figs. 4and 5).

Clay Mineral Assemblages as Climatic IndicatorsClay minerals in the ocean basins close to continents are

mainly derived from land and were formed as a result of variousweathering processes, which were ultimately controlled by cli-mate. The clay mineral types and the proportions of the individ-ual clay minerals in marine sediments therefore depend on theclimatic conditions on land and on the nature of the sourcerocks. The distribution of different clay minerals in the present-day oceans reveals a latitudinal zonation that strongly reflectsthe pedogenic zonation and climatic conditions on the adjacentcontinental land masses (Biscaye, 1965; Griffin et al., 1968; Li-sitzin, 1972; Windom, 1976). Clay mineral assemblages in deep-ocean sedimentary sequences are, therefore, useful tools for re-constructing the paleoclimate through time. In high latitudesclay minerals are an especially important tool, because climaticvariations and erosional processes are pronounced. Investiga-tions of clay mineral assemblages supplement the paleoclimaticinformation provided by other indicators (Chamley, 1989). Forexample, coarse-grained ice-rafted debris in marine sedimentsindicates that the adjacent continent was glacierized at sea level(see the preceding), but special clay mineral assemblages trans-ported by currents may indicate the onset of glacial conditionsin the source area and may therefore predate the first ice-rafteddebris. Whereas the ice-rafted debris is difficult to interpret interms of characterizing the intensity of glaciation, the clay sedi-mentology provides important information on the dominantweathering processes, and thus may assist in distinguishing be-tween local mountain glaciation and continent-wide glaciation.

Chlorite and illite are especially common in the marine sedi-ments of high latitudes (Biscaye, 1965; Griffin et al., 1968; Mo-riarty, 1977). These clay minerals are mostly detrital, being theproducts of physical weathering on land or of glacial scouring,particularly of crystalline or metamorphic rocks, such as thosethat are widespread in East Antarctica (e.g., Ravich et al.,1984). Chlorite is a characteristic mineral from low-grade, chlo-rite-bearing metamorphic and basic source rocks, whereas illitetends to be derived from more acidic rocks (Griffin et al., 1968).

Chlorite can only form in the marine environment under ratherrestricted conditions (Griffin et al., 1968). There is no evidencethat illite can form in situ in the open marine environment.

Kaolinite cannot form under glacial conditions. High con-centrations of this clay mineral are normally restricted to be-tween temperate and tropical regions, where intense chemicalweathering, especially of granitic source rocks, and lateritic soilformation occur on land. However, kaolinite occurring in polarregions may be derived from the weathering of older, kaolinite-bearing sediments and the erosion of paleosols (Darby, 1975;Chamley, 1989; Hambrey et al., this volume).

Smectite is derived mainly as a result of submarine chemicalweathering of volcanic material (e.g., Biscaye, 1965; further ref-erences in Kastner, 1981). In contemporary oceans smectite isfound mainly in regions where the volcanic activity that pro-vides the source rocks is high, where sedimentation rates result-ing in long exposure times for volcanic rocks and ashes are low,and where the input of terrigenous material is low and, there-fore, dilution is reduced. However, there is also input of smec-tite from the continents. Cretaceous and older smectites, espe-cially, are often thought to have originated from subaerial weath-ering of basic rocks, such as basalts, under humid to aridclimatic regimes (Chamley, 1979; Chamley et al., 1984).

Origin of Clay MineralsMuch of the East Antarctic craton provides large quantities

of appropriate source rocks for supplying both illites and chlo-rites. This area is composed mainly of migmatites and biotite-quartz-feldspar gneiss and smaller areas of charnockite. Quartz-ites, quartz-mica schists, and some ferruginous schists also oc-cur (Trail and McLeod, 1969; Craddock, 1982; Ravich andFedorov, 1982).

Most, if not all, of the kaolinite found in the sediments atSites 738 and 744 (Figs. 6 and 7) is assumed to be a recyclingproduct. It is derived as a result of the erosion of old kaolinite-bearing sedimentary rocks or soils. The source of the kaolinitemay be the Permian Amery Formation or equivalent rocks thatmay have occupied the Lambert Graben. The Amery Formationcrops out at Beaver Lake, just west of the main stream of theLambert Glacier, close to the present-day coastline. Its exten-sion beneath the ice is not known but is probable. The unit is atleast 500 m thick and is composed of current-bedded feld-spathic sandstones, calcareous in part, made up of angulargrains of quartz, potash feldspar, and minor plagioclase in a ka-olinitic matrix (Trail and McLeod, 1969). At Sites 738 and 744(Figs. 6 and 7), the coincidence of the first occurrence of largeramounts of kaolinite with a distinct increase in the concentra-tion of the detrital clay minerals illite and chlorite provides fur-ther evidence for a detrital source on the Antarctic continent.Found at the Prydz Bay Sites 739 and 742, just in front of theAmery Ice Shelf complex, were even higher kaolinite concentra-tions of about 60% in the lower Oligocene sediments and about20% in the younger sediments, which also indicate a source inthe hinterland (Hambrey et al., this volume).

In several high-latitude settings a correlation between kaolin-ite and smectite has been found and a common source recon-structed. In Baffin Bay, for example, kaolinite and smectiteoriginated from offshore coastal plain sediments of Mesozoic toTertiary age, whilst on the margin of Nova Scotia they are de-rived from Carboniferous to Triassic red beds (Piper and Slatt,1977). For smectites in sediments drilled in the Southern Oceanduring Legs 28, 29, and 113 a continental origin is also assumed(Piper and Pe, 1977; Barker, Kennett, et al., 1988). Evidence ofchemical weathering and smectite formation has been reportedfrom a few soils from Antarctica (Claridge, 1965; compilationsin Campbell and Claridge, 1987, pp. 130 ff.; Chamley, 1989,pp. 26 ff.). A source similar to that of kaolinite is therefore pos-

198


sible for the smectite found on the southern Kerguelen Plateau.However, there is no positive correlation between the smectitecontent and the kaolinite content (Site 744: r = -0.29), whichmakes a common source unlikely. The correlation with chloriteis also poor and negative (Site 744: r = - 0.43). A strong nega-tive correlation exists between the illite and smectite concentra-tions (Site 744: r = -0.93). This is not surprising, becausethese two clay minerals occur in similar proportions and to-gether account for 70%-90% of the clay mineral assemblage. Acontinental source for the smectite in sediments from Site 744therefore seems unlikely. The smectite is derived instead fromsubmarine chemical alteration of volcanic material. For theolder smectite at Site 738, however, a continental source cannotbe excluded.

Paleoclimatic Implications

The clay mineral assemblages recovered from the southernKerguelen Plateau at Sites 738 and 744 reveal only minor fluctu-ations in the Cretaceous to Eocene sediments (Figs. 6 and 7).These assemblages are clearly dominated by smectite-type min-erals, which account for at least 80% of the clay minerals. Thehigh smectite content compared with that of the detrital clayminerals illite and chlorite implies that physical weathering pro-cesses on the nearby continents played only a limited role. Rela-tively stable conditions, with chemical weathering under humidconditions dominant, are therefore assumed. Despite the over-whelming dominance of smectites, minor inputs of detrital clayminerals are documented (Fig. 6). They are restricted to the Co-niacian to lower Santonian (illite), upper Maestrichtian (illite),lower Paleocene (illite and minor chlorite), and a short period inthe middle part of the middle Eocene (chlorite and kaolinite).

The major change in the weathering regime on the Antarcticcontinent occurred near the Eocene/Oligocene boundary. Allfour clay minerals—smectite, illite, chlorite, and kaolinite—re-veal a distinct shift in their concentration patterns. The smectiteconcentrations decrease drastically and the concentrations ofthe detrital clay minerals clearly increase (Figs. 6 and 7).

At Site 738 (Fig. 6) the decrease in smectite content and in-crease in illite content is in Section 119-738B-5H-7, whereas theincrease in chlorite and kaolinite begins upsection, in Section119-738B-5H-5 (no sample from Section 119-738B-5H-6 hasbeen analyzed so far). Both markers occur in the uppermost Eo-cene; the first one gives an approximate age of 38.8 Ma, the lat-ter one 38.4 Ma. The peak minimum and maxima of smectite,illite, chlorite, and kaolinite, respectively, all occur in Section119-738B-5H-3, which corresponds to about 38.0 Ma. The peaksat 38.0 Ma, however, indicate a short excursion only, because inSection 119B-738B-4H-7 (37.2 Ma) the former level is reachedagain. At 37.2 Ma occurs the decrease to persistent lower smec-tite concentrations and higher illite, chlorite, and kaolinite con-centrations. Better age estimates are hampered by the low sedi-mentation rates at Site 738 and an inadequate paleomagnetic re-cord.

Although the upper Eocene clay mineral composition at Site744 is dominated by smectite, smectite abundance decreasesslightly between the upper Eocene and the lower Oligocene (Fig.7). A first, short-lived minimum that occurs in Section 119-744A-19H-2 (37.5 Ma) probably correlates with Site 738. Thesharp kink in the concentration curve occurs in Section 119-744A-18H-2. The illite concentration increases slightly in theuppermost Eocene sediments and abruptly at the point at whichthe smectite concentration decreases. The distribution patternof chlorite is somewhat more diffuse. However, the shift in chlo-rite concentrations seems to start in Section 119-744A-19H-6(38.3 Ma), in uppermost Eocene sediments. Kaolinite first ap-pears in significant amounts also in Section 119-744A-19H-6. It

therefore appears that the change in clay mineral composition atSite 744 begins at about 38.3 Ma.

The peak development in the concentration curve for kaolin-ite occurs in Section 119-744A-16H-7 (146 mbsf), that of chlo-rite in Section 119-744-16H-5 (143 mbsf), and that of smectiteand illite in Section 119-744A-16H-4 (142 mbsf). The corre-sponding ages are 36.0, 35.7, and 35.5 Ma (Barron et al., thisvolume).

The preceding discussion reveals that in latest Eocene time,at 38.8-38.3 Ma, the clay mineralogy started to change signifi-cantly. Before this, the clay mineral assemblages were character-ized by the dominance of smectite-type clay minerals and thelack of the detrital clay minerals illite and chlorite, indicatingchemical weathering under moderately warm, humid climaticconditions. In younger times, the assemblages were dominatedby illite and chlorite, which are indicative of a cooler climateand physical weathering conditions. Thus, cooler, probably gla-cial conditions established in East Antarctica from about 38.5Ma resulted in intensive physical weathering.

In Oligocene to recent sediments the clay minerals chlorite,illite, and kaolinite do not reveal changes in their concentrationpatterns as distinctive as those close to the Eocene/Oligoceneboundary, which indicates that the main shift in the weatheringregime happened then. It does not necessarily imply that themain shift in paleoclimate occurred at that time. Once the conti-nent was under glaciers, the clay mineral composition of thesediments made available by physical weathering would not havechanged significantly in response to a further cooling event. Thetotal amount of detrital clay minerals, however, may react tosuch a climatic change.

No positive or negative correlation among illite, chlorite,and kaolinite has been recognized. This implies that severalsource areas influenced the composition of the clay mineral as-semblages found at Sites 738 and 744. Complex interactions ofprimary availability and distribution processes are probably re-sponsible for the concentration of clay minerals in the deep-seasediments. The data do not allow the different source areas andtheir fluctuating influence on the total clay mineral assemblagesto be defined. However, the sum of the detrital clay minerals il-lite, chlorite, and kaolinite reflects changes in the intensity ofphysical weathering, and therefore probably indicates climaticchanges. For reasons of simplicity, it is not the sum of detritalclay mineral concentrations but the negatively correlated smec-tite concentrations that are used for further paleoclimatic inter-pretations, with high concentrations characteristic of warm cli-mate and reduced input of physical weathering products (Fig. 8).

Following the lower Oligocene minimum in smectite concen-trations, the values rise suddenly, but do not reach their formerconcentrations (Fig. 8). Between the lower Oligocene and thelower part of the lower Miocene, smectite concentrations fluctu-ate at about 40%, with a slight decreasing trend. They indicatesomewhat warmer conditions than in early Oligocene time, butcolder conditions than in the rest of Paleogene and Cretaceoustime. A climatic optimum in the early Miocene is indicated bymaximum Neogene smectite concentrations of about 60%. Thismaximum dates from about 20.5 to 19.0 Ma (Barron et al., thisvolume). Another indication of warmer climate is found in theenhanced smectite concentrations within the uppermost Mio-cene, which are followed by strongly decreasing values indica-tive of a major cooling phase (Fig. 8).

The smectite concentrations on the southern Kerguelen Pla-teau reveal essentially the same features as the oxygen isotopecurve for planktonic foraminifers in the southwest Pacific (Shack-leton and Kennett, 1975; Fig. 8): a decrease in middle Eocenetime, a marked drop at the Eocene/Oligocene boundary, an in-crease after the early Oligocene minimum, relatively constant

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Smectite (%)

20 40 60 80 100

Oxygen isotope data for planktonic foraminifers

To+3 +2 +1 0 -1 -2

Quaternary

PlioceneΛΛΛ

late Miocene

middle Miocene

early Miocene

ΛΛΛ

late OligoceneΛΛΛ

early Oligocene

late Eocene

middle Eocene

early Eocene

Paleocene

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8&CO

Site

279

i277

CO

Pliocenelate Miocene

mid Miocene

early Miocene

mid-lateOligocene

mid Oligocene

early Oligocene

late Eocene

mid Eocene

mid-early Eocene

early Eocene

late Paleocene

This paper Shackleton and Kennett (1975)

Figure 8. Correlation of the oxygen isotope curve for planktonic foraminifers in the southwest Pacific (Shackleton and Kennett, 1975) and smectitecontent on the southern Kerguelen Plateau. The generalized trend is shaded.

values with a slight decreasing trend to early early Miocene, amaximum in early Miocene time, and a distinctive decreasingtrend in the younger sediments. The correlation between the ox-ygen isotope curve and the smectite content is striking in Paleo-cene to early Miocene times, but less convincing in younger sed-iments. The weaker correlation in younger material is probablydue to oceanographic changes connected with the cooling of theSouthern Ocean and to plate tectonic evolution. The opening ofthe Drake Passage as a deep-water pathway probably initiatedthe development of an intensified circumpolar Antarctic currentand thus created a more complicated circulation pattern aroundAntarctica, which also may have influenced the distribution ofclay minerals.

PREGLACIAL CONDITIONS

In Late Cretaceous time the Antarctic continent reached acentral polar position (Smith et al., 1982). Australia was stillconnected to Antarctica. The Cretaceous Period generally isthought to have been characterized by an equitable, warm, andstable world climate that already had governed most of earlierMesozoic time. This general configuration did not change muchduring Late Cretaceous, Paleocene, and early Eocene time (Haq,1981). Oxygen isotope data confirm the relatively high tempera-tures postulated from faunal and floral evidence. They furtherimply low thermal gradients between high and low latitudes andalso low thermal gradients between ocean surface and deep wa-

ters (Bowen, 1966; Kennett, 1982; Shackleton and Kennett,1975; Oberhànsli and Hsü, 1986). The presence of barriers tosurface and deep water circulation around Antarctica prohib-ited the generation of a circumpolar circulation, and, therefore,warm equatorial surface currents were able to influence the highsouthern latitudes and exert a major influence on the global cli-mate and oceanography (Barron and Baldauf, 1989).

Some evidence for Cretaceous glaciation at sea level was pre-sented by Frakes and Francis (1988). They interpreted the occur-rence of outsized "exotic blocks" in Lower Cretaceous strata ofcentral Australia to be due to ice rafting. They further con-cluded that the warm Cretaceous climate allowed at least sea-sonal periglacial or even glacial conditions in southern high lati-tudes. Matthews and Poore (1980) also speculated that thebuildup of Antarctic ice may date back to the Cretaceous Pe-riod. The latest oxygen isotope data obtained from Late Creta-ceous fossils from James Ross Island, Antarctica, also seem toindicate the presence of cool polar regions in Late Cretaceoustime (Pirrie and Marshall, 1990).

The sedimentological data from southern Kerguelen PlateauSite 738 do not support glacial conditions at sea level in Creta-ceous, Paleocene, and early Eocene times. The input of terrige-nous material was low and there is no indication of ice rafting.The terrigenous material present is mainly clay sized, which im-plies transport by suspension or wind. The clay mineral compo-sition is dominated by smectite. The lack of detrital clay miner-

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als indicates that physical weathering processes were of minorimportance on the nearby continents. Chemical weathering pro-cesses and a warm and humid climate are therefore assumed tohave been dominant. The high carbonate concentrations, com-bined with high bulk-sedimentation rates recorded at Site 738,and the high primary content of opal that has been diageneti-cally altered to chert can be regarded as indications of highproductivity.

ONSET OF EAST ANTARCTIC GLACIATION

The oldest isolated gravel and terrigenous sand grains of ice-rafted origin were detected in upper middle Eocene sediments atSite 738. Their age is 45.0-42.3 Ma. These occurrences correlateexactly with slightly enhanced illite concentrations. Chlorite andkaolinite also have short and slight maxima in the upper middleEocene, but they slightly precede the maxima in ice-rafted mate-rial and illite. It is thus obvious that ice had already reached sealevel in late middle Eocene time, but the extent of the glaciationis unknown. In comparison with present-day conditions, and inview of the dominance of smectites over detrital clay minerals, itis assumed that in late middle Eocene time glacierization wasmuch less extensive than that of today. Some glaciers probablyreached sea level, but much of the main East Antarctic conti-nent was ice free.

Prior to Leg 119, Margolis and Kennett (1970) recorded iso-lated and probably glacially derived sand grains in lower Eoceneand upper middle Eocene sediments from the subantarctic SouthPacific Ocean. At least local glacierization in Paleocene-middleEocene time has been postulated by Birkenmajer (in press) fromglaciomarine sediments on King George Island (South Shetland Is-lands). No evidence for ice rafting during Paleocene and earlyEocene time, however, has been found on the Kerguelen Pla-teau.

According to isotopic data, especially from the Atlantic, In-dian, and Southern oceans, a long-term cooling trend began inearliest middle Eocene time at about 52 Ma. This trend ismarked by several major cooling steps separated by intervals ofmore stable conditions and occasional warming (Shackletonand Kennett, 1975; Shackleton, 1986; Miller et al., 1987). Thesurface water temperature in subpolar regions dropped fromabout 19°C in the early Eocene to about 11°C in the late Eo-cene. The bottom waters experienced about the same degree ofcooling (Shackleton and Kennett, 1975). A middle Eocene cool-ing event is also recorded by fossil land plants on the NorthernHemisphere (Wolfe, 1978). This event may also be reflected inthe general decrease in carbonate-accumulation rates in the In-dian Ocean at 44 Ma (Davies and Worsley, 1981) and a migra-tion of high-latitude planktonic microfossil assemblages towardlower latitudes (Oberhànsli and Hsü, 1986).

A major intensification of glaciation, which probably marksthe onset of the continental-scale East Antarctic glaciation, isrecorded in sediments of early Oligocene age. All investigatedsediment parameters at Sites 738 and 744 document this majorevent (Figs. 2-7).

The clay mineral assemblages started to change at about 38.5Ma. Prior to that time, they were dominated by smectite-typeminerals and only trace amounts of the detrital clay minerals il-lite, chlorite, and kaolinite. This assemblage is characteristic ofchemical weathering under humid conditions. At 38.5 Ma thesmectite concentrations began to decrease steadily while the de-trital clay minerals became more important, which indicatesthat the chemical weathering conditions on East Antarcticawere gradually replaced by physical weathering conditions. Thepeak development in the clay mineral distribution (i.e., mini-mum smectite concentrations and maximum illite, chlorite, andkaolinite concentrations) found to occur at 36.0-35.5 Ma im-plies continental rather than mountain glaciation.

A very sharp and intense pulse of ice rafting is recorded atboth Sites 738 and 744 in lower Oligocene sediments with an ageof 36.0 Ma. Early Oligocene ice-rafted debris was also recordedfrom Leg 113 in the Weddell Sea (Barker, Kennett, et al., 1988).These occurrences further support a major intensification ofEast Antarctic glaciation at this time. It is interesting to notethat this maximum occurs later than the initial change in theclay mineralogy, which is explained by the assumption that thefirst glacially derived clay minerals were not transported by icebut by suspension, and that it took about 2.5 Ma from the es-tablishment of an East Antarctic Ice Sheet until ice reached sealevel.

A distinct and sharp change in carbonate deposition also oc-curred at Sites 738 and 744 in early Oligocene time. This changecorrelates with the ice-rafting signal at 36 Ma. Prior to thistime, carbonate content was relatively high and uniform (90%-95%). In Oligocene and younger strata, carbonate concentra-tions fluctuate from 65% to 95%. The diminished carbonateconcentrations are due to dissolution, probably as a result of thegeneration of cold and aggressive bottom waters.

Although the sediments recovered from the southern Ker-guelen Plateau yielded strong evidence of intense glaciation byearly Oligocene time, the dispute of continental vs. mountainglaciation cannot yet be resolved. The most striking evidencefor continent-wide glaciation comes from the Prydz Bay Sites739 and 742. There, massive diamictites deposited as waterlaintills document a position of the grounding line of the Amery IceShelf complex approximately 140 km north of the edge of thepresent-day floating ice shelf, which implies a fully establishedEast Antarctic Ice Sheet (Hambrey et al., 1989; this volume).The sediments recovered during the CIROS-1 drilling operationin the Ross Sea strongly support this hypothesis (Barrett, 1989).A continental glaciation by early Oligocene time (>35 Ma) isalso postulated based on glaciomarine sediments from KingGeorge Island (Gazdzicki, 1989).

Further evidence for the onset of continental glaciation inearly Oligocene time comes from oxygen isotope data. The mostdramatic Cenozoic decrease in δ 1 8θ occurred shortly after theEocene/Oligocene boundary, at about 35.8 Ma (Shackleton andKennett, 1975; Shackleton, 1986; Miller et al., 1987). The cool-ing in East Antarctica and the buildup of continental ice thatreached sea level are responsible for the cooling of the surfacewaters of the Southern Ocean and the formation of sea ice.Cold and saline water could descend to a greater depth and thusform cold bottom water. Surface water temperatures droppedfrom about 11° to 6°-7°C, and bottom water temperaturesdropped from 8°-9° to 5°C. This cooling of deep ocean watersand surface waters occurred within some 100,000 yr (Shackletonand Kennett, 1975; Shackleton, 1986). Prentice and Matthews(1988) postulated that at least as much ice as today has existedin Antarctica since about 42.5 Ma.

OLIGOCENE TO HOLOCENE GLACIAL HISTORY

The steady northward drift of Australia since late early Eo-cene/middle Eocene time (Weissel and Hayes, 1972), the north-ward drift of the Indian plate, and the opening of the DrakePassage allowed the establishment of a circum-Antarctic currentby the earliest Oligocene time (Kennett and Shackleton, 1976).This current had an immense influence on the thermal isolationof Antarctica. The cooling of East Antarctica and of the watermasses of the Southern Ocean also led to an intensification ofoceanic circulation. Sites 738 and 744 are situated within present-day Circumpolar Deep Water, and the several hiatuses clearlyindicate that this water mass has had intensified circulation anda more erosive character since Oligocene time. Hiatuses identi-fied in Site 744 sediments span the time intervals 32.0-28.5,26.5-24.5, 24.0-21.5, 13.5-11.5, 9.0-6.0, and 5.5-4.0 Ma (Bar-

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ron et al., this volume). At Site 738 one large hiatus separatesearly Oligocene from early late Miocene age sediments, span-ning the time interval 35-10 Ma. Further hiatuses comprise theintervals 8.4-6.1 and > 5.8-<4.8 Ma. The sedimentary se-quence between these hiatuses is probably condensed (Barron,Larsen, et al., 1989).

Following the earliest Oligocene minimum in smectite con-centration at 36 Ma, which matches the onset of continentalglaciation in East Antarctica, the values rose again and retainedan almost constant level between about 35 Ma (early Oligocene)and about 20.5 Ma (early Miocene). However, they did notreach their former concentrations. The oxygen isotope datashow a similar trend (Shackleton and Kennett, 1975), implyingsomewhat warmer conditions than in earliest Oligocene time,but colder conditions than in the rest of the Paleogene and Cre-taceous time.

The amount of ice-rafted debris recovered at Sites 738 and744 strongly decreased after its early Oligocene maximum. Thiscould be a result of less intense glaciation. It also could becaused by a lower availability of terrigenous material on thecontinent. The first glaciers that reached the sea eroded largequantities of soils and loose weathering products derived fromboth basement rocks and Permian and younger sedimentarystrata. Later, all the loose, unconsolidated material was removedand the glaciers had to erode unweathered rocks, and thereforewere only able to transport minor amounts of terrigenous mate-rial characterized by a smaller grain size. However, ice-rafted de-bris is documented in all sediments younger than earliest Oligo-cene by the presence of quartz. Although the quartz is foundonly in low quantities, it is more abundant than in the Creta-ceous to Eocene sediments (Bohrmann and Ehrmann, this vol-ume). The quartz is mainly restricted to the silt fraction, but afew sand-sized grains also occur. The presence of quartz in thesesediments indicates that glaciers advanced to the East Antarcticshoreline throughout that time.

The long-term Cenozoic cooling trend was interrupted by anearly Miocene warming, as indicated by maximum Neogenesmectite concentrations of approximately 60% and by isotopicdata (Shackleton and Kennett, 1975; Margolis et al., 1977). Thismaximum is followed by a sharp decrease to values of about40% in the middle Miocene, suggesting a major cooling phase.A worldwide increase in δ 1 8θ of both planktonic and benthicforaminifers occurs at the same time.

Following the late early Oligocene to early late Miocene timeinterval, when ice rafting was of less importance, ice rafting in-tensified in the late Miocene (9.0 Ma) and reached a maximumbetween the latest Miocene and the present (5.8-0 Ma). Thischange in the intensity of ice rafting may indicate an intensifica-tion of glaciation, the buildup of the West Antarctic Ice Sheet,or a cooling of the surface water, allowing further offshoretransport of the debris by icebergs.

The latest Miocene maximum in ice rafting correlates with asharp decrease in carbonate deposition and the onset of intensediatom ooze sedimentation. It suggests a pronounced intensifi-cation of Antarctic glaciation, combined with a northward ex-pansion of cool surface waters and a northward migration ofthe Polar Front, and a drastic shoaling of the CCrD and CCD.This transition from nannofossil ooze to diatom ooze is dated at>6.1 Ma at Site 738 and at 5.8 Ma at Site 744, and also is re-corded in the latest Miocene age sediments at Sites 689 and 690on the Maud Rise (Barker, Kennett, et al., 1988). From thattime on, Sites 738 and 744 have been beneath the high biogenicsiliceous productivity zone.

The late Miocene to Holocene changes in ice-rafting activityand the glacial history of East Antarctica have been deduced inmore detail from the sediments of Sites 745 and 746 in the Aus-tralian-Antarctic Basin (Ehrmann et al., this volume).

ACKNOWLEDGMENTS

This study was generously supported by Grant Fu 119/15 ofthe Deutsche Forschungsgemeinschaft. The author is especiallygrateful to Andreas Mackensen, Michael J. Hambrey, and Mar-tin Melles for fruitful discussions and comments on the manu-script. Anke Hienen, Rita Fröhlking, and Imke Engelbrecht arethanked for technical assistance. The English text was improvedby M. J. Hambrey.

This is Contribution No. 293 of the Alfred Wegener Institutefor Polar and Marine Research.

REFERENCESAubry, M.-R, Berggren, W. A., Kent, D. V., Flynn, J. J., Klitgord, K. D.,

Obradovich, J. D., and Prothero, D. R., 1988. Paleogene geochro-nology: an integrated approach. Paleoceanography, 3:707-742.

Barker, P. E, Kennett, J. P., et al., 1988. Proc. ODP, Init. Repts., 113:College Station, TX (Ocean Drilling Program).

Barrett, P. J. (Ed.), 1989. Antarctic Cenozoic history from the CIROS-1drillhole, McMurdo Sound. DSIR Bull, 245.

Barron, J. A., and Baldauf, J. G., 1989. Tertiary cooling steps and pa-leoproductivity as reflected by diatoms and biosiliceous sediments.In Berger, W. H., Smetacek, V. S., and Wefer, G. (Eds.), Productiv-ity of the Oceans: Present and Past. Life Sci. Res. Rep., 44:341-354.

Barron, J. A., and Keller, G., 1982. Widespread Miocene deep-sea hia-tuses: coincidence with periods of global cooling. Geology, 10:557-581.

Barron, J., Larsen, B., et al., 1989. Proc. ODP, Init. Repts., 119: Col-lege Station, TX (Ocean Drilling Program).

Berggren, W. A., Kent, D. V., Flynn, J. J., and Van Couvering, J. A.,1985. Cenozoic geochronology. Geoi. Soc. Am. Bull., 96:1407-1418.

Birkenmajer, K., in press. Tertiary glacial and interglacial deposits,South Shetland Islands, Antarctica: geochronology versus biostra-tigraphy (a progress report). Bull. Pol. Acad. Sci. Earth Sci.

Biscaye, P. E., 1964. Distinction between kaolinite and chlorite in recentsediments by X-ray diffraction. Am. Miner., 49:1281-1289.

, 1965. Mineralogy and sedimentation of recent deep-sea clayin the Atlantic Ocean and adjacent seas and oceans. Geol. Soc. Am.Bull., 76:803-832.

Blank, R. G., and Margolis, S. V., 1975. Pliocene climatic and glacialhistory of Antarctica as revealed by southeast Indian Ocean deep-seacores. Geol. Soc. Am. Bull., 86:1058-1066.

Bornhold, B. D., 1983. Ice-rafted debris in sediments from Leg 71,southwest Atlantic Ocean. In Ludwig, W. J., Krasheninnikov, V. A.,et al., Init. Repts. DSDP, 71: Washington (U.S. Govt. Printing Of-fice), 307-316.

Bowen, R., 1966. Paleotemperature Analysis: Amsterdam (Elsevier).Campbell, I. B., and Claridge, G.G.C., 1987. Antarctica: Soils, Weath-

ering Processes and Environment: Amsterdam (Elsevier).Chamley, H., 1979. North Atlantic clay sedimentation and paleoenvi-

ronment since the Late Jurassic. In Tàlwani, M., Hay, W., and Ryan,W.B.F. (Eds.), Deep Drilling Results in the Atlantic Ocean. Am.Geophys. Union, Maurice Ewing Sen, 3:342-361.

, 1989. Clay Sedimentology: Berlin (Springer).Chamley, H., Maillot, H., Duée, G., and Robert, C , 1984. Paleoenvi-

ronmental history of Walvis Ridge at the Cretaceous-Tertiary transi-tion, from mineralogical and geochemical investigations. In Moore,T. C , Jr., Rabinowitz, P. D., et al., Init. Repts. DSDP, 74: Washing-ton (U.S. Govt. Printing Office), 685-695.

Ciesielski, P. F , Ledbetter, M. T., and Ellwood, B. B., 1982. The devel-opment of Antarctic glaciation and the Neogene paleoenvironmentof the Maurice Ewing Bank. Mar. Geol., 46:1-51.

Claridge, G.G.C., 1965. The clay mineralogy and chemistry of somesoils from the Ross dependency, Antarctica. N.Z. J. Geol. Geophys.,3:186-220.

Craddock, C , 1982. Geological map of Antarctica. In Craddock, C.(Ed.), Antarctic Geoscience: Madison (Univ. Wisconsin Press).

Darby, D. A., 1975. Kaolinite and other clay minerals in Arctic Oceansediments. J. Sediment. Petrol., 45:272-279.

Davies, T. A., Weser, O. E., Luydendyk, B. P., and Kidd, R. B., 1975.Unconformities in the sediments of the Indian Ocean. Nature, 253:15-19.

202


Davies, T. A., and Worsley, T. R., 1981. Paleoenvironmental implica-tions of oceanic carbonate sedimentation rates. In Warme, J. E.,Douglas, R. G., and Winterer, E. L. (Eds.), The Deep-Sea DrillingProject: A Decade of Progress. Soc. Econ. Paleont. Mineral. Spec.Publ., 32:169-179.

Dietrich, G., and Ulrich, J., 1968. Atlas zur Ozeanographie: Mannheim(Bibliographisches Institut).

Drewry, D. J., 1986. Glacial Geological Processes: London (Arnold).Fisher, R. L., Jantsch, M. Z., and Comer, R. L., 1982. GeneralBathy-

metric Chart of the Oceans (GEBCO), Scale 1:10,000.000: Ottawa(Canadian Hydrographic Service), 5.9.

Frakes, L. A., and Francis, J. E., 1988. A guide to Phanerozoic cold po-lar climates from high-latitude ice rafting in the Cretaceous. Nature,333:547-549.

Gazdzicki, A., 1989. Planktonic foraminifera from the Oligocene Po-lonez Cove Formation of King George Island, West Antarctica. Pol.Polar Res., 10:47-55.

Griffin, J. J., Windom, H., and Goldberg, E. D., 1968. The distribu-tion of clay minerals in the world ocean. Deep-Sea Res. Oceanogr.Abstr., 15:433-459.

Grobe, H., 1986. Spàtpleistozàne Sedimentationsprozesse am Antarkti-schen Kontinentalhang vor Kapp Norvegia, δstliche Weddell See.Repts. Polar Res., Bremerhaven (Alfred Wegener Institute), 27.

Hambrey, M. J., Larsen, B., Ehrmann, W. U, and ODP Leg 119 Ship-board Scientific Party, 1989. Forty million years of Antarctic glacialhistory yielded by Leg 119 of the Ocean Drilling Program. PolarRec, 25:99-106.

Haq, B. U, 1981. Paleogene paleoceanography: early Cenozoic oceansrevisited. Oceanol. Ada, Proc. 26th Int. Geol. Congr., Geol ofOceans Symp., Paris, 4, Suppl.:71-82.

Hayes, D. E., and Vogel, M., 1981. General Bathymetric Chart of theOceans (GEBCO), Scale 1:10,000.000: Ottawa (Canadian Hydro-graphic Service), 5.13.

Hsü, K. J., McKenzie, J. A., Oberhànsli, H., Weissert, H., and Wright,R. C , 1984. South Atlantic Cenozoic paleoceanography. In Hsü, K.J., LaBrecque, J. L., et al., Init. Repts. DSDP, 73: Washington (U.S.Govt. Printing Office), 771-785.

Kastner, M., 1981. Authigenic silicates in sediments: formation and dia-genesis. In Emiliani, C. (Ed.), The Sea (vol. 7): New York (Wiley),915-980.

Kellogg, T. B., and Kellogg, D. E., 1988. Antarctic cryogenic sediments:biotic and inorganic facies of ice shelf and marine-based ice sheetenvironments. Palaeogeogr., Palaeoclimatol., Palaeoecol., 67:51-74.

Kennett, J. P., 1982. Marine Geology: London (Prentice Hall).Kennett, J. P., and Shackleton, N. J., 1976. Oxygen isotopic evidence

for the development of the psychrosphere 38 Myr ago. Nature, 260:513-515.

Kolla, V., Be, A.W.H., and Biscaye, P. E., 1976. Calcium carbonate dis-tribution in the surface sediments of the Indian Ocean. J. Geophys.Res., 81:2605-2616.

Lange, H., 1982. Distribution of chlorite and kaolinite in eastern Atlan-tic sediments off North Africa. Sedimentology, 29:427-432.

Lisitzin, A. P., 1972. Sedimentation in the World Ocean, with Emphasison the Nature, Distribution and Behavior of Marine Suspensions.Soc. Econ. Paleontol. Mineral Spec. Publ., 17.

Mackensen, A., Barrera, E., and Hubberten, H.-W., in press. Neogenecirculation in the southern Indian Ocean: evidence from benthic fo-raminiferal assemblages, carbonate, and stable isotopes (ODP Site751). In Schlich, R., Wise, W. S., et al., Proc. ODP, Sci. Results,120: College Station, TX (Ocean Drilling Program).

Margolis, S. V., and Kennett, J. P., 1970. Antarctic glaciation duringthe Tertiary recorded in sub-Antarctic deep-sea cores. Science, 170:1085-1087.

Margolis, S. V., Kroopnick, P. M., and Goodney, D. E., 1977. Ceno-zoic and late Mesozoic paleoceanography and paleoglacial historyrecorded in circum-Antarctic deep-sea sediments. Mar. Geol., 25:131-147.

Matthews, R. K., and Poore, R. Z., 1980. Tertiary δ 1 8 θ record and gla-cioeustatic sea-level fluctuations. Geology, 8:501-504.

Miller, K. G., Fairbanks, R. G., and Mountain, G. S., 1987. Tertiary ox-ygen isotope synthesis, sea level history, and continental margin ero-sion. Paleoceanography, 2:1-19.

Moriarty, K. C , 1977. Clay minerals in southeast Indian Ocean sedi-ments, transport mechanisms and depositional environments. Mar.Geol., 25:149-174.

Oberhànsli, H., and Hsü, K. J., 1986. Paleocene-Eocene paleoceanog-raphy. In Hsü, K.J. (Ed.), Mesozoic and Cenozoic Oceans: Int.Lithosphere Program, Geodynamics Series, 15:85-100.

Piper, D.J.W., and Pe, G. G., 1977. Cenozoic clay mineralogy fromD.S.D.P. holes on the continental margin of the Australia-New Zea-land sector of Antarctica. N.Z. J. Geol. Geophys., 20:905-917.

Piper, D.J.W., and Slatt, R. M., 1977. Late Quaternary clay mineral dis-tribution on the eastern continental margin of Canada. Geol. Soc.Am. Bull., 88:267-272.

Pirrie, D., and Marshall, J. D., 1990. High-paleolatitude Late Creta-ceous paleotemperatures: new data from James Ross Island, Antarc-tica. Geology, 18:31-34.

Prentice, M. L., and Matthews, R. K., 1988. Cenozoic ice-volume his-tory: development of a composite oxygen isotope record. Geology,16:963-966.

Ravich, M. G., and Fedorov, L. V., 1982. Geologic structure of MacRo-bertson Land and Princess Elizabeth Land, East Antarctica. InCraddock, C. (Ed.): Antarctic Geoscience: Madison (Univ. Wiscon-sin Press), 499-504.

Ravich, M. G., Solov'ev, D. S., and Fedorov, L. V., 1984. GeologicalStructure of Mac. Robertson Land (East Antarctica): New Delhi(Amerind Publishing). (Translation from 1978 Russian text)

Shackleton, N. J., 1986. Paleogene stable isotope events. Palaeogeogr.,Palaeoclimatol., Palaeoecol., 57:91-101.

Shackleton, N. J., and Kennett, J. P., 1975. Paleotemperature history ofthe Cenozoic and the initiation of Antarctic glaciation: oxygen andcarbon isotope analyses in DSDP Sites 277, 279, and 281. In Ken-nett, J. P., Houtz, R. E., et al., Init. Repts. DSDP, 29: Washington(U.S. Govt. Printing Office), 743-755.

Smith, A. G., Hurley, A. M., and Briden, J. C , 1982. Palaokontinen-tale Weltkarten des Phanerozoikums: Stuttgart (Enke).

Trail, D. S., and McLeod, I. R., 1969. Geology of the Lambert Glacierregion. In Craddock, C. (Ed.), Geologic Maps of Antarctica. Am.Geogr. Soc, Antarctic Map Folio Sen, Plate 11.

Van Andel, T. H., 1975. Mesozoic/Cenozoic calcite compensation depthand the global distribution of calcareous sediments. Earth Planet.Sci. Lett., 26:187-194.

Weissel, J. K., and Hayes, D. E., 1972. Magnetic anomalies in thesoutheast Indian Ocean. In Hayes, D. E. (Ed.), Antarctic Oceanol-ogy II: the Australian-New Zealand Sector. Antarct. Res. Sen, 19:165-196.

Whitworth, T, III, 1988. The Antarctic Circumpolar Current. Oceanus,31:53-58.

Windom, H. L., 1976. Lithogenous material in marine sediments. InRiley, J. P., and Chester, R. (Eds.), Chemical Oceanography (vol. 5):New York (Academic Press), 103-135.

Wolfe, J. A., 1978. A paleobotanical interpretation of Tertiary climatesin the Northern Hemisphere. Am. Sci., 66:694-703.

Date of initial receipt: 19 December 1989Date of acceptance: 10 April 1990Ms 119B-121

203

W. U. EHRMANN

APPENDIX TABLE 1.Sediment Composition from Holes 738B, 738C, and 744A

All data are given in weight percent of the bulk sediment. Furthercarbonate concentrations for Holes 738B and 738C are given in Appen-dix Table 2 and by Barron, Larsen, et al. (1989).

Appendix Table 1 (continued).

Core, section,interval (cm)

Depth(mbsf)

Organiccarbon Carbonate Opal Nonbiogenic


119-738B-

3H-1, 125-1293H-5, 48-524H-1, 110-1144H-3, 48-524H-5, 78-824H-7, 30-345H-1, 100-1045H-3, 48-525H-5, 48-525H-7, 30-346H-1, 48-526H-3, 48-526H-5, 48-527H-1, 100-1047H-3, 48-527H-5, 48-528H-3, 48-528H-5, 48-528H-7, 30-349H-1, 80-849H-3, 48-5210H-1, 48-5210H-3, 48-5211H-1, 48-5211H-5, 48-5212H-1, 48-5212H-3, 48-5212H-5, 48-5213H-1, 48-5213H-3, 48-5214X-1, 48-5214X-3, 48-5215X-1, 72-7615X-3, 48-5215X-5, 48-5216X-1, 16-2016X-3, 30-3417X-1, 60-6417X-3, 48-5218X-1, 48-5218X-3, 48-5218X-5, 48-5219X-1, 42-4619X-3, 46-5019X-5, 48-5220X-1, 48-5220X-3, 48-5220X-5, 48-5221X-2, 48-5221X-3, 48-5222X-1, 48-5222X-3, 48-5224X-2, 48-52

119-738C-

4R-3, 48-525R-1, 46-507R-2, 74-7610R-2, 36-3811R-3, 10-1216R-1, 48-5216R-5, 42-4417R-1, 58-6017R-3, 112-11417R-5, 56-6018R-1, 70-7418R-3, 50-5418R-6, 48-5219R-2, 70-7420R-1, 20-24

Depth(mbsf)

14.7519.9824.0626.3529.5231.9433.4635.8538.7341.4442.4845.4848.4852.4754.8557.7964.3567.2970.0371.3073.9880.4883.4885.4891.4894.9897.98

100.98104.47107.17108.68111.68118.52121.28124.28127.66130.80137.80140.68147.28150.28153.28156.92159.96162.98166.68169.68172.68176.28179.26185.88188.88206.58

219.38226.06247.04275.66286.50332.08338.02341.86345.24347.58351.60354.40358.88362.70370.40

Organiccarbon

(<%)Carbonate

(%)

0.0087.3088.8992.2992.7690.4689.8889.2093.6294.0593.1292.1298.1194.0592.7090.6892.5093.4994.9593.5794.0994.0595.3797.4893.3391.8594.2094.9393.6592.0688.4291.2298.1693.1993.5294.8098.1791.5591.0794.5593.3996.0897.9295.8096.2596.9196.7395.2095.8195.8094.9094.9595.32

93.1094.2593.6092.2092.4392.5491.1792.8794.8396.6097.1495.6199.2698.3894.79

Opal(%)

38.986.245.893.762.674.643.816.512.180.710.561.160.491.021.492.070.610.711.140.870.510.320.540.500.802.541.000.781.583.744.572.640.532.291.071.840.583.973.952.352.630.500.240.580.270.280.321.320.520.351.280.620.91

0.700.821.113.922.141.301.682.390.550.380.611.010.190.482.89

Nonbiogenic(%)

61.026.465.223.954.574.906.314.294.205.246.326.721.404.935.817.256.895.803.915.565.405.634.092.025.875.614.804.294.774.207.016.141.314.525.413.361.254.484.983.103.983.421.843.623.482.812.953.483.673.853.824.433.77

6.204.935.293.885.436.167.154.744.623.022.253.380.551.142.32

119-738C-(Cont.)

20R-2, 50-5422R-2, 62-6522R-5, 51-5323R-2, 47-4924R-2, 60-6224R-3, 50-5225R-1, 64-6625R-3, 72-7426R-2, 54-5627R-1, 49-5127R-3, 49-5128R-1, 55-5728R-3, 48-5029R-2, 48-5030R-2, 48-50

119-744A-

1H-1, 49-511H-3, 49-512H-1, 49-512H-4, 49-513H-1,49-513H-4, 49-514H-1, 44-464H-2, 46-504H-3, 40-444H-4, 45-475H-1, 49-505H-2, 40-445H-3, 47-495H-4, 12-156H-1, 49-516H-2, 60-646H-3, 46-506H-4, 49-517H-1, 81-837H-2, 46-507H-3, 46-507H-4, 49-517H-6, 20-248H-1, 49-518H-2, 46-508H-3, 46-508H-4, 49-519H-1, 49-5110H-1, 44-4610H-2, 46-5010H-4, 44-4610H-5, 46-5010H-7, 48-5211H-1, 49-5111H-3, 48-5211H-4, 49-5111H-6, 48-5212H-1, 49-5112H-3, 48-5212H-4, 49-5112H-6, 46-5013H-1, 49-5113H-3, 48-5213H-4, 49-5113H-6, 46-5014H-1, 49-5114H-3, 46-5014H-4, 49-5114H-6, 46-5015H-1, 49-5115H-3, 46-5015H-4, 49-5115H-6, 46-5016H-1, 49-5116H-2, 46-5016H-4, 49-5116H-5, 46-5016H-7, 47-49

372.20391.72396.11401.17411.00412.40419.24422.32430.24438.39441.39448.15451.08459.18468.58

0.493.494.699.19

14.1918.6923.6425.1626.6028.1533.1934.6036.1737.3242.6944.3045.6647.1952.5153.6655.1656.6959.4061.6963.1664.6666.1971.1980.6281.6084.9586.4189.3290.1993.1894.6997.6899.67

102.55104.00106.85109.19112.15113.64116.58118.69121.66123.19126.16128.18131.06132.54135.43137.67139.09142.00143.41146.31

0.160.220.250.040.030.000.00

0.080.00

0.23

0.16

0.110.01

0.26

0.00

0.000.000.01

0.00

0.23

0.00

0.04

0.00

0.00

0.00

0.00

0.16

0.00

0.00

0.00

0.09

0.00

96.1494.2994.9590.7786.4983.1285.2288.3087.7989.1091.5792.9586.6384.6284.61

29.8260.9662.6932.5729.01

3.9169.8789.8691.0181.5191.9581.6072.7573.4389.7686.4495.0591.3868.0482.8779.7377.4793.1382.4278.3282.2484.0093.4579.7190.1588.2292.9390.5288.0589.0390.1583.1884.3089.7688.0489.4791.7989.1589.5991.1987.9394.7591.1188.6587.2582.7478.6988.9684.8287.1477.8282.0161.26

0.831.130.591.276.995.561.130.241.271.240.230.510.492.833.02

51.6627.0526.7646.8126.8957.6413.126.816.29

10.914.52

13.6717.5421.60

8.0210.663.163.42

15.436.038.98

10.512.018.51

14.879.265.823.93

10.296.485.033.543.275.135.404.055.215.315.235.592.002.513.053.312.593.611.812.743.885.979.99

15.256.039.726.72

19.359.78

11.28

3.034.584.467.966.52

11.3213.6511.4610.949.668.206.54

12.8812.5512.37

18.5211.9910.5520.6244.1038.4517.013.332.707.583.534.739.714.972.222.901.795.20

16.5311.1011.2912.024.869.076.818.50

10.182.62

10.003.376.753.536.216.825.575.80

11.6110.395.016.378.535.707.807.106.228.463.446.157.476.787.276.065.015.466.142.838.21

27.46

204




119-744A-(Cont.)

17H-1, 12-1417H-CC, 12-1418H-1, 49-5118H-2, 46-5018H-4, 49-5118H-6, 46-5019H-1, 49-5119H-2, 46-5019H-4, 49-5119H-6, 46-5020H-1, 49-5120H-2, 26-3020H-2, 130-13420H-3, 49-5120H-5, 46-5020H-6, 49-5120H-7, 46-50

Depth(mbsf)

146.82147.29148.08149.52152.49155.40157.59159.06162.09165.06167.00168.29169.29169.96172.81174.28175.70

Organiccarbon

W

0.000.180.00

0.25

0.07

0.13

0.03

0.03

Carbonate(%)

91.5490.0091.8294.6091.5592.1193.7895.0091.5890.7888.6092.4995.6594.0794.6193.4596.89

Opal(%)

3.744.923.322.284.562.232.821.343.814.543.901.390.691.290.901.240.76

Nonbiogenic

W

4.725.084.863.123.895.663.403.664.614.687.506.123.664.644.495.312.35


APPENDIX TABLE 2.Additional Carbonate Concentra-tions of Samples from Holes 738Band 738C


119-738B-

1H-1, 48-501H-3, 48-502H-2, 48-502H-4, 48-502H-6, 48-503H-2, 58-603H-3, 58-603H-4, 58-604H-2, 48-504H-4, 48-504H-6, 48-505H-2, 48-505H-4, 48-505H-6, 48-506H-2, 48-506H-4, 48-506H-6, 48-507H-2, 48-507H-4, 48-507H-6, 48-508H-2, 48-508H-4, 48-508H-6, 48-509H-2, 48-509H-4, 48-5010H-2, 48-5011H-2, 48-5011H-4, 48-5011H-6, 48-5012H-2, 48-5012H-4, 48-5012H-6, 48-5013H-2, 48-5014X-2, 47-4914X-4, 46-48

Depth(mbsf)

0.483.455.928.83

11.7515.5817.0818.5824.9027.7930.6734.4037.2940.1744.0146.5349.5353.4256.3359.2562.9265.8368.7572.4875.4881.9886.9889.9892.9896.4899.48

102.48105.96110.17113.16

Carbonate(%)

63.9971.6452.0039.772.510.91

59.0486.6689.3795.4092.8592.6592.4993.3893.7796.0094.8295.0593.5096.2294.9495.9995.1895.8595.8495.3893.9994.2593.4599.2195.7293.8594.0594.3694.19


119-738B-(Cont.)

15X-2, 48-5015X-4, 48-5015X-6, 48-5016X-2, 47-4917X-2, 47-4917X-4, 48-5018X-2, 46-4818X-4, 46-4819X-2, 48-5019X-4, 48-5019X-6, 48-5020X-2, 48-5020X-4, 48-5020X-6, 48-5021X-2, 48-5021X-4, 44-4622X-2, 48-5022X-4, 48-5023X-1, 40-4224X-1, 48-5024X-3, 48-50

119-738C-

4R-1, 20-224R-2, 4-65R-CC, 0-16R-1, 48-507R-1, 72-747R-3, 19-208R-1, 48-509R-1, 47-5010R-3, 74-7611R-1, 54-5611R-2, 54-5616R-2, 48-5016R-6, 48-5017R-2, 48-5017R-4, 48-5017R-6, 48-5018R-1, 48-5018R-4, 48-5019R-1, 45-4719R-3, 47-492OR-3, 30-3220R-5, 30-3221R-CC, 8-1022R-1, 48-5022R-4, 48-5023R-1, 28-3023R-3, 36-3824R-1, 26-2824R-CC, 1-325R-2, 48-5025R-4, 47-4926R-1, 48-5026R-3, 48-5027R-2, 47-5027R-4, 52-5427R-5, 50-5228R-4, 23-2529R-1, 43-4529R-3, 48-4930R-1, 93-953OR-3, 6-731R-1, 50-5231R-2, 48-5031R-3, 11-13

Depth(mbsf)

119.78122.78125.78128.38139.17142.18148.76151.76158.48161.48164.48168.18171.18174.18177.72179.84187.38190.38195.40205.08208.08

216.10217.44227.88235.68245.52247.99254.88264.57277.54283.94285.44333.58339.58343.19346.07348.94351.38355.88360.95363.97373.50376.50380.67390.08394.58399.48402.56409.16413.26420.58423.57428.68431.68439.87442.92444.40452.33457.63460.68467.53469.66476.80478.28479.41

Carbonate(%)

93.7294.0994.7195.6294.4595.4787.9095.2390.2295.3489.5795.5790.2196.3592.5797.9494.9386.9595.6497.2094.63

95.0194.9595.0594.3790.9889.4796.0588.5993.3990.7969.2491.4593.1093.2494.3894.9395.5495.5087.2592.8292.4785.5894.2096.5496.6393.3688.0384.5688.9889.1987.0583.6184.1590.4092.2586.7787.2794.7577.5894.0493.1088.4582.7243.51

205

W. U. EHRMANN

APPENDIX TABLE 3.Grain-Size Distribution of the Nonbiogenic Sediment Components

at Holes 738B, 738C, and 744A (carbonate-free, opal-free)


119-738B-

3H-1, 125-1293H-5, 48-524H-1, 110-1144H-3, 48-524H-5, 78-824H-7, 30-345H-1, 100-1045H-3, 48-525H-5, 48-525H-7, 30-346H-1, 48-526H-3, 48-526H-5, 48-527H-1, 100-1047H-3, 48-527H-5, 48-528H-3, 48-528H-5, 48-528H-7, 30-349H-1, 80-849H-3, 48-5210H-1, 48-5210H-3, 48-5211H-1, 48-5211H-3, 48-5211H-5, 48-5212H-1, 48-5212H-3, 48-5212H-5, 48-5213H-1, 48-5213H-3, 48-5214X-1, 48-5214X-3, 48-5215X-1, 72-7615X-3, 48-5215X-5, 48-5216X-1, 16-2016X-3, 30-3417X-1, 60-6417X-3, 48-5218X-1, 48-5218X-3, 48-5218X-5, 48-5219X-1, 42-4619X-3, 46-5019X-5, 48-5220X-1, 48-5220X-3, 48-5220X-5, 48-5221X-2, 48-5221X-3, 48-5222X-1, 48-5222X-3, 48-5224X-2, 48-52

119-738C-

4R-3, 48-525R-1, 46-507R-2, 74-7610R-2, 36-3811R-3, 10-1216R-1, 48-5216R-5, 42-4417R-1, 58-6017R-3, 112-11417R-5, 56-6018R-1, 70-7418R-3, 50-5418R-6, 48-5219R-2, 70-7420R-1, 20-2420R-2, 50-5422R-2, 62-65

Depth(mbsf)

14.7519.9824.0626.3529.5231.9433.4635.8538.7341.4442.4845.4848.4852.4754.8557.7964.3567.2970.0371.3073.9880.4883.4885.4888.4891.4894.9897.98

100.98104.47107.17108.68111.68118.52121.28124.28127.66130.80137.80140.68147.28150.28153.28156.92159.96162.98166.68169.68172.68176.28179.26185.88188.88206.58

219.38226.06247.04275.66286.50332.08338.02341.86345.24347.58351.60354.40358.88362.70370.40372.20391.72

>2 mm

(«)

1100000000000000000000001000001000000000000000000000000

00000000003000000

(g)

0.280.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.020.000.000.000.000.000.020.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.00

0.000.000.000.000.000.000.000.000.000.000.030.000.000.000.000.000.00

63 µm-2 mm(%)

27.691.070.180.540.230.000.310.250.290.450.030.190.150.120.000.000.640.030.180.000.000.210.330.110.000.100.000.000.008.810.061.330.140.001.200.040.001.590.200.210.110.160.000.000.140.000.170.000.000.310.050.670.000.07

0.000.001.910.134.145.781.310.05

15.053.570.040.330.020.002.544.250.12

2-63 µm(%)

32.6720.277.015.916.007.294.206.127.436.669.807.876.946.974.439.56

13.4915.329.52

27.2619.6913.1826.3819.5225.705.433.219.892.756.385.95

10.8110.536.123.07

13.1920.09

8.345.23

12.9513.6614.9411.5519.9614.9020.3225.0628.09

6.8021.7518.402.315.27

17.29

10.247.347.73

13.4313.284.755.793.65

21.7518.3418.859.906.748.537.497.31

16.53

<2µm(%)

39.6478.6692.8293.5693.7792.7195.4893.6392.2992.8990.1791.9492.9192.9195.5790.4485.8784.6590.3072.7480.3186.6173.2980.3774.3094.4796.7990.1197.2584.8193.9987.8689.3393.8895.7386.7779.9190.0794.5786.8486.2384.9488.4580.0484.9779.6874.7771.9193.2077.9481.5597.0294.7382.64

89.7692.6690.3686.4482.5889.4792.9096.3063.2078.0981.1189.7793.2491.4789.9788.4483.35

206




U9-738C-(Cont.)

22R-5, 51-5323R-2, 47-4924R-2, 60-6224R-3, 50-5225R-1, 64-6625R-3, 72-7426R-2, 54-5627R-1, 49-5127R-3, 49-5128R-1, 55-5728R-3, 48-5029R-2, 48-5030R-2, 48-50

119-744A-

1H-1, 49-511H-3, 49-512H-1, 49-512H-4, 49-513H-1, 49-513H-4, 49-514H-1, 44-464H-2, 46-504H-3, 40-444H-4, 45-475H-1, 49-505H-2, 40-445H-3, 47-495H-4, 12-156H-1, 49-516H-2, 60-646H-3, 46-506H-4, 49-517H-1, 81-837H-2, 46-507H-3, 46-507H-4, 49-517H-6, 20-248H-1, 49-518H-2, 46-508H-3, 46-508H-4, 49-519H-1, 49-5110H-1, 44-4610H-2, 46-5010H-4, 44-4610H-5, 46-5010H-7, 48-5211H-1, 49-5111H-3, 48-5211H-4, 49-5111H-6, 48-5212H-1, 49-5112H-3, 48-5212H-4, 49-5112H-6, 46-5013H-1, 49-5113H-3, 48-5213H-4, 49-5113H-6, 46-5014H-1, 49-5114H-3, 46-5014H-4, 49-5114H-6, 46-5015H-1, 49-5115H-3, 46-5015H-4, 49-5115H-6, 46-5016H-1, 49-5116H-2, 46-5016H-4, 49-5116H-5, 46-5016H-7, 47-4917H-1, 12-1417H-CC, 12-14

Depth(mbsf)

396.11401.17411.00412.40419.24422.32430.24438.39441.39448.15451.08459.18468.58

0.493.494.699.19

14.1918.6923.6425.1626.6028.1533.1934.6036.1737.3242.6944.3045.6647.1952.5153.6655.1656.6959.4061.6963.1664.6666.1971.1980.6281.6084.9586.4189.3290.1993.1894.6997.6899.67

102.55104.00106.85109.19112.15113.64116.58118.69121.66123.19126.16128.18131.06132.54135.43137.67139.09142.00143.41146.31146.82147.29

>2 mm

(»)

0000000000000

000080000000000000000000000000000000000000000000000000000305

(g)

0.000.000.000.000.000.000.000.000.000.000.000.000.00

0.000.000.000.000.420.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.000.140.000.26

63 µm-2 mm

0.070.110.080.070.160.000.000.110.000.000.000.000.01

9.9625.1417.0816.0238.5318.7016.280.424.180.461.240.170.030.130.190.050.091.310.120.020.010.030.040.080.020.310.110.470.390.030.020.020.080.200.280.100.160.320.020.440.040.310.090.270.020.150.070.260.150.490.140.590.140.840.211.450.05

37.811.465.74

2-63 µm

16.5115.766.685.109.11

20.036.509.905.509.278.04

10.834.42

14.927.259.25

11.0814.2313.6718.799.40

17.573.684.85

12.042.10

11.923.084.837.470.64

15.1819.2924.0519.3218.5013.6469.2591.5213.283.18

11.1110.668.638.24

13.297.03

20.4613.6720.5516.2018.2617.1816.0213.1314.5415.9112.7718.737.95

10.0118.2510.7013.166.959.419.319.01

12.986.52

16.784.085.81

<2 µm

83.4284.1493.2394.8390.7379.9693.5089.9994.5090.7391.9689.1795.57

75.1267.6173.6772.9147.2467.6264.9390.1878.2595.8693.9187.7997.8687.9496.7495.1292.4498.0584.7080.6975.9580.6481.4686.2830.738.17

86.6196.3588.5089.3191.3591.7386.6392.7679.2786.2379.2883.4881.7282.3883.9486.5785.3883.8287.2181.1291.9789.7381.5988.8286.7092.4690.4489.8590.7985.5793.4445.4294.4688.45

207

W. U. EHRMANN



119-744A-(Cont.)

18H-1, 49-5118H-2, 46-5018H-4, 49-5118H-6, 46-5019H-1, 49-5119H-2, 46-5019H-4, 49-5119H-6, 46-5020H-1,49-5120H-2, 26-3020H-2, 130-13420H-3, 49-5120H-5, 46-5020H-6, 49-5120H-7, 46-50

Depth(mbsf)

148.08149.52152.49155.40157.59159.06162.09165.06167.00168.29169.29169.96172.81174.28175.70

>2 mm

(n)

000000001200000

(g)

0.000.000.000.000.000.000.000.000.360.040.000.000.000.000.00

63 µm—2 mm

0.880.400.170.030.580.040.300.210.191.710.030.220.020.580.05

2-63 µm( * )

4.809.274.626.010.624.161.027.269.002.702.691.112.991.557.84

<2 µm

94.3290.3395.2193.9798.8195.8098.6892.5290.8195.5997.2798.6696.9897.8792.11

APPENDIX TABLE 4.Amount of Terrigenous Clay (nonbiogenic, < 2 µm) and Relative Percent-age of the Clay Minerals Smectite, Illite, Chlorite, and Kaolinite at Holes738B, 738C, and 744A


119-738B-

3H-1, 125-1293H-5, 48-524H-1, 110-1144H-3, 48-524H-5, 78-824H-7, 30-345H-1, 100-1045H-3, 48-525H-5, 48-525H-7, 30-346H-1, 48-526H-3, 48-526H-5, 48-527H-1, 100-1047H-3, 48-527H-5, 48-528H-3, 48-528H-5, 48-528H-7, 30-349H-1, 80-849H-3, 48-5210H-1, 48-5210H-3, 48-5211H-1, 48-5211H-3, 48-5211H-5, 48-5212H-1, 48-5212H-3, 48-5212H-5, 48-5213H-1, 48-5213H-3, 48-5214X-1, 48-5214X-3, 48-5215X-1, 72-7615X-3, 48-5215X-5, 48-5216X-1, 16-2016X-3, 30-3417X-1, 60-6417X-3, 48-5218X-1, 48-5218X-3, 48-5218X-5, 48-5219X-1, 42-4619X-3, 46-5019X-5, 48-5220X-1, 48-52

Depth(mbsf)

14.7519.9824.0626.3529.5231.9433.4635.8538.7341.4442.4845.4848.4852.4754.8557.7964.3567.2970.0371.3073.9880.4883.4885.4888.4891.4894.9897.98

100.98104.47107.17108.68111.68118.52121.28124.28127.66130.80137.80140.68147.28150.28153.28156.92159.96162.98166.68

Clay(%)

24.195.084.843.694.284.546.034.023.874.875.706.181.304.585.556.565.594.913.534.054.344.873.001.62

5.555.434.334.174.043.956.165.491.234.334.702.681.124.234.332.673.323.021.473.072.772.10

Smectite(%)

3151376578977059778893938892969290889091

9290918989919196798787959781918894858088829290959394

Illite(%)

513134131018

1511835

106277887

48667773

1810831

146

104232668434

Chlorite(%)

1011241630

13148

:

()

21211

2113212121400411179571102

Kaolinite(%)

8657919

13532121102221

21222011121122221675611122

208




119-738B-(Cont.)

20X-3, 48-5220X-5, 48-5221X-2, 48-5221X-3, 48-5222X-1, 48-5222X-3, 48-5224X-2, 48-52

119-738C-

4R-3, 48-525R-1, 46-507R-2, 74-7610R-2, 36-3811R-3, 10-1216R-1, 48-5216R-5, 42-4417R-1, 58-6017R-3, 112-11417R-5, 56-6018R-1, 70-7418R-3, 50-5418R-6, 48-5219R-2, 70-7420R-1, 20-2420R-2, 50-5420R-5, 89-9022R-2, 62-6522R-5, 51-5323R-2, 47-4924R-2, 60-6224R-3, 50-5225R-1, 64-6625R-3, 72-7426R-2, 54-5627R-1, 49-5127R-3, 49-5128R-1, 55-5728R-3, 48-5029R-2, 48-5030R-2, 48-50

119-744A-

1H-1, 49-511H-3, 49-512H-1, 49-512H-4, 49-513H-1, 49-513H-4, 49-514H-1, 44-464H-2, 46-504H-3, 40-444H-4, 45-475H-1, 49-505H-2, 40-445H-3, 47-495H-4, 12-156H-1, 49-516H-2, 60-646H-4, 49-517H-1, 81-837H-2, 46-507H-3, 46-507H-4, 49-517H-6, 20-248H-1, 49-518H-2, 46-508H-3, 46-508H-4, 49-519H-1, 49-5110H-1, 44-4610H-2, 46-5010H-4, 44-4610H-5, 46-5010H-7, 48-5211H-1, 49-51

Depth(mbsf)

169.68172.68176.28179.26185.88188.88206.58

219.38226.06247.04275.66286.50332.08338.02341.86345.24347.58351.60354.40358.88362.70370.40372.20377.09391.72396.11401.17411.00412.40419.24422.32430.24438.39441.39448.15451.08459.18468.58

0.493.494.699.19

14.1918.6923.6425.1626.6028.1533.1934.6036.1737.3242.6944.3047.1952.5153.6655.1656.6959.4061.6963.1664.6666.1971.1980.6281.6084.9586.4189.3290.19

Clay(%)

2.123.242.863.143.714.193.12

5.574.564.783.364.485.516.644.572.922.361.833.030.511.042.082.68

3.823.726.706.07

10.7412.389.16

10.238.697.755.93

11.8411.1911.82

13.918.117.77

15.0320.8326.0011.043.002.117.273.314.159.504.372.152.765.10

14.008.968.589.693.967.832.090.698.812.528.853.016.173.235.386.33

Smectite(%)

95899794929593

97999294

8398929694969090866794897778689185939292919588888376

463122393548384759252254494414383844453937434342474113346836576145

Illite( * )

2713653

2174

151725298

13185

10191830

511555629

111622

325050414836453719616234324671474542394353404842385378522350362646

ChloriteW

2311112

1001

00111101041123222111111221

14156

11151389

15107

11117

14101353

14898

15947

1268227

Kaolinite(%)

1212101

0001

101111010

121122112112112001

84

239249774

101831639

13428216211365

112

209

W. U. EHRMANN



119-744A-(Cont.)

11H-3, 48-5211H-4, 49-5111H-6, 48-5212H-1, 49-5112H-3, 48-5212H-4, 49-5112H-6, 46-5013H-1, 49-5113H-3, 48-5213H-4, 49-5113H-6, 46-5014H-1, 49-5114H-3, 46-5014H-4, 49-5114H-6, 46-5015H-1, 49-5115H-3, 46-5015H-4, 49-5115H-6, 46-5016H-1, 49-5116H-2, 46-5016H-4, 49-5116H-5, 46-5016H-7, 47-4917H-1, 12-1417H-CC, 12-1418H-1, 49-5118H-2, 46-5018H-4, 49-5118H-6, 46-5019H-1, 49-5119H-2, 46-5019H-4, 49-5119H-6, 46-5020H-1, 49-5120H-2, 26-3020H-5, 46-5020H-6, 49-5120H-7, 46-50

Depth(mbsf)

93.1894.6997.6899.67

102.55104.00106.85109.19112.15113.64116.58118.69121.66123.19126.16128.18131.06132.54135.43137.67139.09142.00143.41146.31146.82147.29148.08149.52152.49155.40157.59159.06162.09165.06167.00168.29172.81174.28175.70

Clay(%)

4.425.009.218.674.095.257.164.936.665.955.426.863.175.526.106.036.305.604.544.915.582.437.67

12.474.464.494.582.823.705.323.363.514.544.336.815.854.355.202.16

Smectite(%)

673946333127523144445132414324404751364748162537544654755577387470738688948469

Illite( * )

17513851525927564042404438416446394045364467522125272611307

46112011984

1226

Chlorite(%)

96

14122

14798

106

1411118

1075

1096

14201511201110136

1156

1242234

Kaolinite(%)

8424

150

145943

10116347498132

2697952

104

103422021

210

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