deep sea drilling project initial reports volume 90 · 2007. 4. 25. · 42. miocene to early...

42. MIOCENE TO EARLY PLIOCENE OXYGEN AND CARBON ISOTOPE STRATIGRAPHY INTHE SOUTHWEST PACIFIC, DEEP SEA DRILLING PROJECT LEG 901

James P. Kennett, Graduate School of Oceanography, University of Rhode Island

ABSTRACT

High-resolution oxygen and carbon isotope stratigraphy is presented for Miocene to early Pliocene sequences atthree DSDP sites from the Lord Howe Rise, southwest Pacific, at water depths ranging from 1,300 to 2,000 m. Site 588is located in the warm subtropics (~ 26°S), whereas Sites 590 and 591 are positioned in transitional (northern temperate)water masses (~ 31°S). Benthic foraminiferal oxygen and carbon isotope analyses were conducted on all sites; plankton-ic foraminiferal isotope data were generated for Site 590 only. Sample resolution in these sequences is on the order of50,000 yr. or better. The chronological framework employed in this study is based largely upon ages assigned to Neo-gene calcareous nannoplankton boundaries.

The benthic oxygen isotope record exhibits several major features during the Neogene. During most of the early Mi-ocene, δ 1 8 θ values were relatively low, reaching minimum values in the late early Miocene (19.5 to 16.5 Ma), and record-ing the climax of Neogene warmth. This was followed by a major increase in benthic δ 1 8 θ values between ~ 16.5 and13.5 Ma, which is interpreted as representing major, permanent accumulation of the East Antarctic ice sheet and cool-ing of bottom waters. During the 3 m.y. 1 8O enrichment, surface waters at these middle latitudes warmed between 16and 14.5 Ma.

During the remainder of the middle and late Miocene, benthic δ 1 8 θ values exhibit distinct fluctuations, but the aver-age value remained unchanged. The isotopic data show two distinct episodes of climatic cooling close to the middle/lateMiocene boundary. The earliest of these events occurred between 12.5 and 11.5 Ma in the latest middle Miocene. Thesecond cooling event occurred from 11 to 9 Ma, and is marked by some of the highest δ 1 8 θ values of the entire Miocene.This was followed by relative warmth during the middle part of the late Miocene. The latest Miocene and earliest Plio-cene (6.2 to 4.5 Ma) were marked by relatively high δ 1 8 θ values, indicating increased cooling and glaciation.

During the middle Pliocene, at about 3.4 Ma, a 0.4‰ increase in benthic δ 1 8 θ documents a net increase in averageglobal ice volume and cooling of bottom waters. During this interval of increased glaciation, surface waters warmed by2-3°C in southern middle-latitude regions. During the late Pliocene, between 2.6 and 2.4 Ma, a further increase in δ 1 8 θoccurred; this has been interpreted by previous workers as heralding the onset of Northern Hemisphere glaciation.

Surface-water warming in the middle latitudes occurred in association with major high-latitude glacial increases inthe early middle Miocene (16-14 Ma), middle Pliocene (-3.5 Ma), and late Pliocene (~2.4 Ma). These intervals werealso marked by increases in the vertical temperature gradient in the open ocean.

Intersite correlation is enhanced by using carbon isotope stratigraphy. The great similarity of the δ1 3C time-series re-cords within and between ocean basins and with water depth clearly indicates that changes in oceanwide average δ 1 3C ofHCOJ in seawater dominated the records, rather than local effects. Broad changes in the Neogene δ 1 3C record werecaused largely by transfer of organic carbon between continental and oceanic reservoirs. These transfers were caused bymarine transgressions and regressions on the continental margins.

The dominant feature of Neogene δ 1 3C stratigraphy is a broad late early to early middle Miocene increase of aboutl‰ between ~ 19 and 14.5 Ma. This trend occurred contemporaneously with a period of maximum coastal onlap(transgression) and maximum Neogene climatic warmth. The δ 1 3C trend terminated during the expansion of the Ant-arctic ice sheet and associated marine regression.

The latest Miocene carbon isotope shift (of up to - O.75‰) at 6.2 Ma is clearly recorded in all sites examined andwas followed by relatively low values during the remainder of the Neogene. This shift was caused by a glacioeustatic sea-level lowering that exposed continental margins via regression and ultimately increased the flux of organic carbon to thedeep sea. An increase in δ 1 3C values during the early Pliocene (~ 5 to 4 Ma) resulted from marine transgression during atime of global warmth.

INTRODUCTION

Cenozoic cooling, which began in the middle Eocene,proceeded gradually, but was punctuated by sudden tran-sitions from one climatic state to another. Global cli-mate underwent major changes during the Miocene Ep-och. The most important change occurred during theearly middle Miocene, between -17 and 14 Ma. Amarked increase in δ 1 8 θ of benthic foraminiferal calciteat this time is commonly interpreted as recording the

Kennett, J. P., von der Borch, C. C , et al., Init. Repts. DSDP, 90: Washington (U.S.Govt. Printing Office).

2 Address: Graduate School of Oceanography, University of Rhode Island, Narragan-sett, R.I. 92882-1197.

rapid growth of the East Antarctic ice sheet and associ-ated cooling of high-latitude surface and deep-ocean wa-ters (Savin et al., 1975; Shackleton and Kennett, 1975b;Kennett, 1977; Schnitker, 1980; Keigwin, 1979; Wood-ruff et al., 1981; Savin et al., 1981). Others have inter-preted this increase in δ 1 8 θ values as reflecting only adecrease in bottom-water temperatures, unaccompaniedby any increase in Antarctic ice (Matthews and Poore,1980).

Oxygen isotope ratios of benthic and planktonic for-aminifers indicate that these events were accompaniedby a marked increase in the equator-to-pole thermal gra-dient and an associated decrease in meridional heat trans-port (Savin et al., 1975, in press). Tropical Pacific near-surface temperatures increased while high-latitude regions

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J. P. KENNETT

cooled as Antarctic ice volume increased during the mid-dle Miocene (Savin et al., in press). Increased latitudinaltemperature gradients are recorded by the progressive di-vergence of tropical planktonic and benthic foraminifer-al δ 1 8θ curves caused by increasing vertical temperaturestratification. The early Miocene was marked by limitedAntarctic glaciation, gentler equator-to-pole and tropi-cal surface-to-bottom temperature gradients, and warm-er high-latitude surface waters. This early Miocene oceancontrasts sharply with that of the late Miocene, whichwas more similar to the modern ocean (Matthews et al.,1980).

The primary objective of DSDP Leg 90 was to studythe Cenozoic paleoceanographic and paleoclimatic evo-lution of the South Pacific, with emphasis on the Neo-gene. A latitudinal traverse of carbonate-rich Neogenesequences was cored during Leg 90 on the Lord HoweRise and New Zealand Platform. Since these sedimenta-ry sequences were obtained using the hydraulic pistoncorer (HPC) and the extended core barrel (XCB), theyare relatively continuous and undisturbed, and are ofgreat value for studies of high-resolution stratigraphyand paleoceanography. HPC sequences in the relativelysoft carbonate ooze that blankets this shallow southwestPacific region were longer than 200 m in each site stud-ied. At Site 588, the HPC sequence is 315 m thick, thelongest ever obtained. This sequence contains a continu-ous, largely uninterrupted record spanning the past 17m.y. from the late early Miocene to the Quaternary.

This paper presents a detailed study of oxygen andcarbon isotope stratigraphy during the Miocene and ear-ly Pliocene in three of the Leg 90 sites: 588, 590, and591 (Fig. 1; Table 1). A number of features of strati-

graphic value were identified in the isotopic records ofthese continuous, high-resolution sequences. The strati-graphic resolution in these sites is much higher than ob-tained in previous investigations and thus allows a moredetailed examination of Miocene paleoceanographic andpaleoclimatic history than was previously possible. Thispaper describes major trends in the isotopic data anddiscusses causal mechanisms underlying the changes. Fu-ture work will consider detailed correlation of the stableisotope records and further examination of cause/effectrelationships.

Site Descriptions

The locations of the three sites are shown in Figure 1,with additional relevant information provided in Table1. The three sequences are composed of relatively uni-form, light-colored, foraminifer-bearing to foraminifer-rich nannofossil oozes to chalks throughout the Mio-cene and Pliocene. Calcium carbonate content is greaterthan 90% throughout, the terrigenous sedimentary com-ponents are virtually nonexistent. Preservation of fora-miniferal assemblages is good to excellent in all sections,and deposition always occurred at depths shallower thanthe foraminiferal lysocline (Kennett and von der Borch,this volume). A noticeable decrease in the quality ofpreservation of planktonic foraminifers occurs in chalksof early Miocene age at Site 590 and of early middle andearly Miocene age at Site 591. Some planktonic forami-niferal specimens from these intervals are slightly recrys-tallized, but benthic foraminifers show no such change.

Rates of sedimentation show marked changes duringthe Neogene (Gardner, Dean, et al., this volume). Eachof the sites exhibits similar trends. Average early Mio-

30°S

40°

<f. /in i i

‰ Ti iTiTi iTiTi 11111 tTi tii~TJV •v, - 5 0 0 m 1 Λ

nil in

•'I.

I I

50°150° 160° 170°W 180° 170°E

Figure 1. Location of Sites 588, 590, and 591 in relation to modern surface-water boundaries and generalcirculation in the southwest Pacific (Knox, 1970; Wyrtki, 1974). Dashed line indicates 500 m isobath.

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MIOCENE TO PLIOCENE ISOTOPE STRATIGRAPHY, SOUTHWEST PACIFIC

Table 1. Data for sites analyzed in this chapter.

Latitude (S)Longitude (E)Water depth (m)Modern water mass

Thickness of Miocenesequence drilled (m)

Depth at which XCB sub-stituted for HPC (m)

Depth of ooze/chalktransition (m)

Site 588

26°06.70'161-13.60'

1533Warm-subtropical

288

248.6

260

Site 590

31° 10.02'163°21.51'

1299Cool-subtropical to

transitional306

250.7

283.1

Site 591

31-35.06'164-26.92'

2131Cool-subtropical to

transitional316

283.1

290

cene (23 to 16 Ma) carbonate accumulation rates for theopen ocean were relatively constant (4.9 g/cm2 per 103

yr.). Accumulation rates were much more variable fromabout 15 to 10 Ma (middle Miocene) reaching a peak(up to 13 g/cm2 per 103 yr.) at approximately 12 Ma.The late Miocene to earliest Pliocene (10 to 4 Ma) wasmarked by fairly constant and average rates of accumu-lation (Gardner, Dean, et al., this volume). At about5 Ma, near the Miocene/Pliocene boundary, accumula-tion rates again increased, and widespread, shortlived,remarkably high accumulation rates (up to 27 g/cm2 per103 yr.) occurred between 4 and 3 Ma (late early Plio-cene). This does not simply reflect an artifact of thetime scale, because a major modification of the timescale would be required substantially to change the pat-tern of increased sedimentation rates. Relatively highrates of sedimentation during this interval have also beenreported elsewhere for areas such as the Columbia Basinand the Caribbean Sea (Prell, Gardner, et al., 1982).These high rates during the late early Pliocene were par-ticularly pronounced at Sites 590 and 591 and were dueto high productivity of calcareous biogenic material (sitechapters, this volume), related to the position of the Sub-tropical Divergence (Tasman Front) that crosses the Tas-man Sea in the vicinity of Sites 590 and 591 (Fig. 1).Sedimentation rates are much lower to the north (Site588) and south (Site 592). Accumulation rates duringthe late Pliocene were much lower than in the early Plio-cene, but remained higher than average Neogene rates.

Oceanographic Setting

The physical oceanography has not been studied inmuch detail in the region of the three sites used in thisstudy (Fig. 1). Knox (1970) and Wyrtki (1974) showedthe position of the major surface-water masses and theirboundaries, and directions of surface-water flow in theregion (Fig. 1). Site 588 lies within the warm-subtropicalwater mass, whereas Sites 590 and 591 lie close to thepresent-day position of the Subtropical Divergence (Tas-man Front) that separates warm-subtropical and tem-perate water masses (Denham and Crook, 1976; Stan-ton, 1979, 1981). The Subtropical Divergence representsthe southern extent of the South Pacific subtropical gyreas the East Australian Current turns to the east, passingto the north of the barrier formed by New Zealand. Up-welling at the Subtropical Divergence is caused by thedivergence of the two surface-water masses and the in-teraction of the eastward-flowing gyral circulation withthe Lord Howe Rise (Stanton, 1981). This has created

high biogenic sedimentation rates, which in the late Ne-ogene ranged up to 131 m/m.y. at Site 591, the site clos-est to the Divergence.

Sites 588 and 590 are located in similar water depths,1500 and 1300 m, respectively, yet bottom-water temper-atures at Site 590 are 0.5°C warmer (Site 588 = 3.75°C;Site 590 = 4.25°C), because of the eastward passageacross this area of the warm waters contained in the sub-tropical gyre.

The sites on the Lord Howe Rise have moved about5° to the north during the Neogene as the result of themovement of the Indo-Australian Plate (Sclater et al.,in press). The northward movement of these sites intowarmer waters during the Neogene should be reflectedby planktonic assemblages indicative of increasingly warmsurface water and by lower planktonic δ 1 8 θ values, un-less the water masses themselves changed latitudinalpositions.

Stratigraphy and Chronology

Rich calcareous microfossil assemblages facilitatedbiostratigraphic subdivision of the sequences. Calcare-ous nannoplankton and planktonic foraminiferal zonalboundaries are from Lohman (this volume) and Jenkinsand Srinivasan (this volume), as summarized in the In-troduction (this volume). The nannoplankton zonalboundaries were determined using uneven sample spac-ing ranging from a few cm to up to 8 m, resulting invarying precision in the placement of individual zonalboundaries. Planktonic foraminiferal zonal boundarieswere determined by examining core-catcher samples on-ly (i.e., at 9 m stratigraphic intervals). Minor adjust-ments were made on zonal boundary positions in thelate Miocene and Pliocene of Site 590, as discussed byElmstrom and Kennett (this volume).

The chronological framework employed here is basedupon ages assigned to Neogene calcareous nannoplank-ton zonal (NN) boundaries (Table 2) by Berggren (1981),and the scheme was uniformly applied by the Leg 90 sci-entific group (Introduction, this volume). I have alsoadopted an age of 6.2 Ma for the NN1 la/1 lb boundary.The chronology of the latest Miocene through early Qua-ternary of Site 590 was that used by Elstrom and Ken-nett (this volume), with some minor modifications fromthe scheme used by the Leg 90 group. The chronological

Table 2. Ages assigned to boundariesbetween nannofossil NN zones toestablish chronological frameworkin this paper.

NN zones

NN 19/2018/1917/1816/1715/1614/1513/1412/13llb/12lla/llb

Ma

0.551.902.302.403.453.654.104.605.206.20

NN zones

NNlO/lla9/108/97/86/75/64/53/42/31/2

Ma

9.5011.0012.0012.2013.0015.0017.0017.8019.0020.50

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J. P. KENNETT

intercalibration used by Berggren (1981) and most otherworkers for the past 10 years is based upon a correlationbetween Chron 9 and Anomaly 5. This scheme placesthe middle/late Miocene boundary at 11 Ma. After Leg90 was completed, a major adjustment to the Neogenetime scale was made by Berggren, Kent, and Van Cou-vering (1985), based upon a correlation between Chron11 and Anomaly 5 (Barron et al., in press). If this corre-lation is accepted, then an age of 8.92 to 10.42 Ma mustbe assigned to Chron 11 in accordance with the Berg-gren, Kent, and Flynn (1985) paleomagnetic time scale.In this scheme, the middle Miocene/late Miocene bound-ary, which is correlated with upper Chron 11, is assignedan age of 9.5 Ma (Barron et al., in press). This changealso requires a shift of about 1.5 to 2 m.y. toward youn-ger ages for much of the late Miocene and late middleMiocene. Nevertheless, to maintain consistency amongthe Leg 90 contributions, the older time scale is employedhere.

All three sites contain continuous sedimentary sequenc-es except for an unconformity at 456 m at Site 590. Thisunconformity contains the interval from the upper partof the Globorotalia miozea Zone (late early Miocene) tothe lower part of the G. fohsi s.l. Zone (early middleMiocene). The calcareous nannoplankton do not indi-cate an unconformity at this level, but suggest insteadthe NN4 and part of NN5 are missing at 470.3 m. In thisstudy I recognize the hiatus at 456 m and place its agebetween about 18 and 16.5 Ma (latest early Miocene).

Sample intervals for the isotopic analyses in the threesites were, for the most part, evenly spaced at 75 cm(Appendix). The quantity of time represented betweensamples varies according to changes in sedimentationrates. For much of the Miocene, sedimentation ratesranged from 1.5 to 3 cm/103 yr., which provides a sam-pling interval of about 25 × 103 yr. In the Pliocene, sed-imentation rates were much higher, ranging from about5 to 13 cm/103 yr. This provides a sampling interval ofabout 5 to 15 × 103 yr. Sampling intervals for isotopeanalyses are smaller than those used to define the posi-tions of the calcareous nannoplankton boundaries andhence the isotope record is of higher resolution than thechronological framework.

METHODS

Samples were disaggregated by soaking in buffered Calgon solu-tion, washed over a 63 µm sieve, and oven-dried at 50°C.

Isotopic analyses were performed on the planktonic foraminifersof the Globigerinoides immaturus-G. quadrilobatus-G. sacculifera lin-eage in Site 590 in the 295-412 µm fraction. The relatively small sizefraction was used to minimize the effect of size on isotopic variability(Berger et al., 1978; Killingley et al., 1981).

Isotopic analyses were performed on four benthic species (> 150 µmsize fraction) because no species extends throughout the entire Mio-cene (Appendix). Cibicidoides coryelli or Oridorsalis umbonatus wereanalyzed in the lower portion of the section, incorporating part of theearly Miocene; C. wuellerstorfi was analyzed in the middle part of thesection and C. kullenbergi in the upper part of the sequence. The mid-dle part of the sequence of Site 591 was analyzed using either C. kul-lenbergi or C. wuellerstorfi. Cibicidoides wuellerstorfi and C. kul-lenbergi have been shown to be isotopically indistinguishable withinanalytical error. The isotopic systematics of Cibicidoides spp. and Ori-dorsalis spp. have been discussed by Savin et al. (1981) and Loutit,Kennett, et al., (1983). C. wuellerstorfi deposits its test near carbonisotopic equilibrium, but Oridorsalis is 1.06‰ lower in δ 1 3C than C.

wuellerstorfi (Savin et al., 1981). With respect to δ 1 8 θ , C. wuellerstor-fi deposits its test out of oxygen isotopic equilibrium by 0.72‰,whereas Oridorsalis is closer to equilibrium (-0.03‰) (Savin et al.,1981).

Almost all analyses in this study were performed on Cibicidoidesspp. No adjustments were made to the data for disequilibrium effectsknown for this genus. Because only a few measurements were made onOridorsalis spp., the values of this form were adjusted relative to Cibi-cidoides by adding + 1.06‰ for δ 1 3C and subtracting 0.72 for δ 1 8 θ assuggested by Savin et al. (1981).

Specimens selected for isotopic analysis were ultrasonically cleanedin reagent-grade methanol and roasted in vacuo at 400°C for 1 hr. Thesamples were then reacted in 100% orthophosphoric acid at 50°C.The evolved CO 2 was purified by distilling four times and was ana-lyzed for 1 8 O/ 1 6 O and 1 3 C/ 1 2 C ratios using an on-line VG Micromass602D mass spectrometer. The analytical precision of the mass spec-trometer is ±O.l‰ for both δ 1 8 θ and δ 1 3C, which is based upon theanalyses of a laboratory standard of powdered B-l (Orchardo, in theAppendix to Murphy and Kennett, this volume). This standard wasanalyzed at the beginning and end of each day. All isotopic results areexpressed as per mil difference from PDB.

RESULTS

All isotopic results are presented in the Appendix. Foreach of the sites, the isotopic data were plotted againstboth depth and age and a three-point running averageof the isotopic data was plotted against age. For Site588, the data are plotted in Figures 2, 3, and 4; for Site590 in Figures 5, 6, and 7; and for Site 591 in Figures 8,9, and 10.

Site 588

Benthic Foraminiferal Oxygen Isotopes

The main features of Miocene benthic oxygen isotoperecords are exhibited by Figures 3 and 4. During the ear-ly Miocene, a slow, progressive decrease in δ 1 8 θ values(apparent warming) occurred, reaching several minima(maximum isotopic temperatures) during the late earlyMiocene between about 19.5 and 17 Ma. Within this in-terval of optimal Neogene warmth, an episode of slight-ly higher δ 1 8 θ values (apparent cooling) occurred be-tween 19 and 18 Ma.

From 17 Ma (latest early Miocene) to about 13.5 Ma(middle middle Miocene), δ 1 8 θ values increased by about1.5%o. This increase, recognized in most deep-sea sedi-mentary sections, represents a shift in average values ofabout l%o from the early Miocene to the late Miocene.

From about 13,5 Ma (middle middle Miocene) to about4 Ma (early Pliocene), the benthic δ 1 8 θ values remainedpermanently offset, but exhibit no apparent long-termtrends. There were, however, significant fluctuations dur-ing this interval. Within this 13.5 to 4 Ma interval, epi-sodes interpreted as having been warmer and with lesspolar ice occurred between about 13 and 12 Ma (late mid-dle Miocene), 7.5 and 6.2 (middle late Miocene), and5.2 and 4 Ma (early Pliocene). Episodes interpreted ashaving been cooler and with more polar ice occurredfrom about 11 to 10 Ma (earliest late Miocene) and 6.2to 5.2 Ma (latest Miocene).

Benthic Foraminiferal Carbon Isotopes

The Miocene benthic carbon isotope record (Figs. 3and 4) exhibits both short-term and long-term variabili-

1386

δ18O°/ooPDB


δ13C°/ooPDB

100

150 —

200 -

E 250 -

300 -

350 -

400

Figure 2. Benthic δ 1 8 θ and δ13C for Holes 588, 588A, and 588C plotted against depth at Site 588 andbiostratigraphic zonations for calcareous nannoplankton (Lohman, this volume) and planktonicforaminifers (Jenkins and Srinivasan, this volume).

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J. P. KENNETT

22 -

Figure 3. Benthic δ 1 8 θ and δ13C for Holes 588, 588A, and 588C plotted against age atSite 588 and biostratigraphic zonations for calcareous nannoplankton (Lohman, thisvolume) and planktonic foraminifers (Jenkins and Srinivasan, this volume).

1388


δ 1 8 O °/oo PDB

21 -

22 -

Figure 4. Smoothed benthic δ 1 8 θ and δ13C data (three-point running aver-ages) for Holes 588, 588A, and 588C and biostratigraphic zonations forSite 588, plotted against age.

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J. P. KENNETT

δ 1 8 O°/oo PDBBenthics Planktonics

δ 1 3 C °/oo PDB

200 -

G. quadrilobatus—sacculifer complex

Figure 5. Benthic and planktonic δ 1 8 θ and δ13C for Holes 590A and 590B plotted against depth in Site 590 and bio-stratigraphic zonations for calcareous nannoplankton (Lohman, this volume) and planktonic foraminifers (Jen-kins and Srinivasan, this volume).

1390


ty. Relatively high δ1 3C values mark the earliest Miocenebetween about 23 and 22 Ma, followed by lower valuesbetween 22 and 19.5 Ma. This, in turn, was followed bythe most conspicuous feature of the Neogene benthiccarbon isotope record: a period of increased δ1 3C valuesfrom the late early Miocene to the early middle Mio-cene, between 19.5 and 15 Ma. The increase leading tothese high δ1 3C values consisted of a shift of about l%othat occurred between -19.8 and 18.8 Ma. Althoughoverall values were high during this interval, distinct fluc-tuations occurred, including minima centered at 18.0,17.0, and 15.8 Ma. The return to lower values is alsodistinctive, consisting of a - l%o shift between 15.4 and14.2 in the early middle Miocene. Following this majormid-Neogene episode, there were no such major fluctu-ations in δ1 3C during the remainder of the middle andlate Miocene except for a negative peak centered at about11.2 Ma near the middle/late Miocene boundary. Fol-lowing this, the most distinctive change in the Neogenebenthic carbon isotope record occurred during the lateMiocene at about 6.2 Ma, representing a marked shift inδ 1 3C values of about - O.7%o. This represents the well-known latest Miocene carbon isotope shift (see Vincentet al., 1980, for review). Generally lower values markthe remainder of the latest Miocene and the early Plio-cene, although there was a trend toward slightly highervalues during the early Pliocene (5.2 to 4 Ma). The rec-ord from Site 588 suggests that the most importantevents in δ1 3C occurred at - 2 2 to 21.6 Ma; 19.8 to 19Ma; 15.4 to 14.2 Ma; and 6.2 Ma.

Site 590


Sedimentation at Site 590 was continuous except foran unconformity centered in the latest early Miocene be-tween about 18 and 16.5 Ma. The early Miocene ex-hibits relatively low δ 1 8 θ values, with the lowest valuesfor the entire Neogene probably dated between about 19and 18 Ma. From the latest early Miocene (about 16.5Ma) until about 13 Ma, δ 1 8 θ values increased irregularlyby about l%o.

During the remainder of the Miocene and early Plio-cene, average values remained unchanged, although im-portant oscillations occurred. These include an episodeof lower values (warmer temperatures) centered at about12.5 Ma followed rapidly by a brief interval of highervalues centered at 12.2 Ma. An interval of high values isalso centered at about 4.8 Ma in the earliest Pliocene.

Important positive shifts (~O.75%o) occurred duringthe Pliocene between about 3.5 Ma and 2.2 Ma; thesechanges are discussed in detail by Elmstrom and Ken-nett (this volume).

Planktonic Foraminiferal Oxygen Isotopes

The planktonic oxygen isotope record (Figs. 6 and 7)generally exhibits higher variability than the benthicrecord, especially during the early Miocene, latest Mio-cene, and Pliocene. The high variability in the early Mi-ocene, between 20 and 18 Ma, of up to 1.3‰, is con-sidered to have resulted from recrystallization of plank-tonic foraminiferal tests.

The early middle Miocene exhibits distinctly lower val-ues (warmer isotopic temperatures) between 16 and 14.5Ma. This was followed by the well-known δ 1 8 θ increase( of about l%o), which continued in the δ 1 8 θ planktonicrecord at this site until about 12.4 Ma. Following this,no apparent long-term trends occurred during the latestmiddle and late Miocene in planktonic δ 1 8 θ until 6 Ma.At that time, a distinct positive shift occurred of aboutO.5%o (cooler temperatures). Average latest Miocene toPliocene values are almost O.5%o higher than during theearlier part of the late Miocene. After 6 Ma, several dis-tinct planktonic δ 1 8 θ oscillations occurred, including anincrease between about 4.6 Ma and 3.7 Ma (0.5‰) andanother increase beginning at 3 Ma and continuing tothe Pleistocene (at least to 1.8 Ma). Intervals marked bylower values occurred between 5.4 and 5.2, between 4.8and 4.6, and between 3.7 and 3 Ma. Detailed descrip-tions of these changes have been made by Elmstrom andKennett (this volume).

Benthic and Planktonic Foraminiferal Carbon Isotopes

Distinct parallelism occurs between benthic and plank-tonic carbon isotope records (Figs. 6 and 7), includingmany of the short-term oscillations, although the ampli-tude of change often differs. Because of their similarityof change, the records are described as one. The earliestMiocene sequence (20.8 to 18 Ma), preceding the uncon-formity, exhibits relatively low δ1 3C values. A major in-crease in δ1 3C (~ 1.2%o) occurred during the latest earlyMiocene to early middle Miocene, with highest valuescentered at about 15 Ma. This is the most distinct car-bon isotope episode of the Neogene. Following this broadearly middle Miocene peak, there was a somewhat irreg-ular decrease in δ1 3C values, which reached particularlylow values between about 11.5 and 10.5 Ma, close to themiddle/late Miocene boundary. Values again increasedtoward a maximum in the middle late Miocene betweenabout 9.5 and 8 Ma. In the latest Miocene at 6.2, a dis-tinct, rapid, negative carbon isotope shift occurred thatis synchronous with the Chron 6 carbon isotope shift.The carbon isotope change in the planktonic record(~O.9%o) is almost double that exhibited in the benthicrecord (~O.4%o). Relatively low δ1 3C values mark the re-mainder of the late Neogene sequence. Following theMiocene/Pliocene boundary, at 5.2 Ma, a temporary in-crease in δ1 3C occurred. This, in turn, was followed by areturn to generally lower values typical of the middleand late Pliocene. Details of these changes are describedin Elmstrom and Kennett (this volume).

Site 591


The stable isotope record at Site 591 extends from thelate early Miocene (17 Ma) to the middle Pliocene (4Ma) (Figs. 8-10). The benthic δ 1 8 θ record exhibits thelowest values (warmest isotopic temperatures) at about17 Ma. From 17 Ma (latest early Miocene) to about 14.5Ma, an irregular increase in δ 1 8 θ values of about 1.25‰occurred, synchronous with the early middle Mioceneshift observed in other sites. The time of maximumchange occurred between —16 and 15.2 Ma. Although

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J. P. KENNETT

δ1 8O°/oo PDB

Benthicsin o in o m o in

20 -

/ rG. quadrilobatus—immaturis—sacculifer complex

Figure 6. Benthic and planktonic δ 1 8 θ and δ13C for Holes 590A and 590B plotted against age in Site 590 and biostratigraphic zonationfor calcareous nannoplankton (Lohman, this volume) and planktonic foraminifers (Jenkins and Srinivasan, this volume). Note un-conformity between 16.5 and 18 Ma.

1392


δ 1 8 o Voo PDBBenthics Planktonics

δ 1 3 C ° / o o P D BPlanktonics Benthics

p o

20 -

G. quadrilobatus—immaturis—sacculifer complex

Figure 7. Smoothed benthic and planktonic δ 1 8 θ and δ13C data (three-point running averages) for Holes 59OAand 590B and biostratigraphic zonations for Site 590, plotted against age.

'

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J. P. KENNETT

δ18O°/ooPDB

Figure 8. Benthic δ 1 8θ and δ13C for Holes 591 and 59IB plotted against depth in Site 591 and biostratigraphic zo-nation for calcareous nannoplankton (Lohman, this volume) and planktonic foraminifers (Introduction, thisvolume). Stable isotope data for Oridorsalis umbonatus have been adjusted to that of Cibicidoides by adding+ 1.06‰ for δ13C and subtracting 0.72 for δ 1 8 θ (see text for discussion).

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δ18O°/ooPDB

Figure 9. Benthic δ 1 8 θ and δ13C for Holes 591 and 591B plotted against age in Site 591 and biostratigraphic zonation for calcar-eous nannoplankton (Lohman, this volume) and planktonic foraminifers (Introduction, this volume). Stable isotope datafor Oridorsalis umbonatus have been adjusted to that of Cibicidoides by adding + 1.06‰ for δ13C and subtracting 0.72 forδ 1 8 θ (see text for discussion).

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δ1 80°/oo PDB

Figure 10. Smoothed benthic δ 1 8 θ and δ13C data (three-point running averages) for Holes 591 and 591B and biostrati-graphic zonations for Site 591, plotted against age.

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stratigraphic resolution during the middle Miocene atthis site is not high, some trends seem to be evident.From 14.5 to 11.5 Ma, the δ 1 8 θ values remained muchthe same or perhaps slightly increased, reaching a max-imum close to the middle/late Miocene boundary at-11.5 Ma. During the late Miocene and the Pliocene,there was relatively little change in δ 1 8 θ in this section.

Benthic Foraminiferal Carbon Isotopes

The benthic carbon isotope record at Site 591 (Figs.8-10) is similar to that of Site 588 and exhibits consider-able variability. As at Site 588, the most conspicuousfeature of the record is the interval of increased δ13C val-ues during the late early to early middle Miocene (baseof sequence at 17 Ma to 14.5 Ma). During the remain-der of the Miocene the δ13C values generally decreased,reaching a minimum between about 11.6 and 10.6 Ma,associated with the middle/late Miocene boundary. Thiswas followed by a δ13C maximum at 10.0 Ma and thengradually decreasing values in the late Miocene, culmi-nating in a distinct shift toward low δ13C values fromabout 7.5 to 6 Ma. However, unlike the record at theother two sites, this latest Miocene carbon isotope shiftwas more gradual, rather than sharply defined at 6.2Ma. A gradual depletion in 13C of 0.4‰ occurred be-tween 7.5 and 6.2, with an additional shift of O.5%o cen-tered at 6.2 Ma. During the early Pliocene between 5and 4 Ma, average δ13C values were about O.25%o higherthan latest Miocene values.

DISCUSSION

Oxygen Isotopes

The oxygen isotope composition of calcite is con-trolled by both the oxygen isotope composition of ambi-ent seawater and its temperature (Epstein et al., 1953).The isotopic composition of the oceans is strongly af-fected by changes in continental ice volume because 16Ois preferentially stored in such ice. An increase in conti-nental ice volume causes an increase in δ 1 8 θ values ofthe ocean. A record of ice volume is provided by benthicforaminiferal oxygen isotope changes during intervalswhen the temperatures of ocean bottom waters remainedconstant. Planktonic foraminiferal δ 1 8 θ is controlled bylocal and global changes in the isotopic composition ofthe oceans and temperature of surface waters. It is diffi-cult to differentiate the effects of temperature from thosedue to ice volume. The detailed relations provided herebetween benthic and planktonic δ 1 8 θ changes at differ-ent latitudes may help differentiate between changes thatresulted from ice volume and those that resulted fromtemperature.

This chapter has no statistically based correlations be-tween the isotope stratigraphies of the individual sites.Nevertheless, visual comparison of the curves shows thatmany events are easily recognizable in the three sequenc-es. Intersite correlation of these isotopic events also re-veals that the assigned ages of individual events, basedupon correlations using the nannoplankton biostratigra-phy, are not identical in each of the sequences.. Individ-ual peaks are often offset by 0.1 to 0.3 m.y. It is clear

that intersite correlation could be enhanced by the com-bined use of δ13C and δ 1 8 θ stratigraphy. More detailedcomposite isotopic curves will be established in the fu-ture by applying quantitative techniques of signal corre-lation to these and other existing records. In the mean-time, visual comparisons of the isotope stratigraphy forthe three sites are adequate for interpretation of pale-oceanographic history. In the following sections, a syn-thesis is provided of this history.

Early Miocene

The earliest Miocene (oldest than 19.5 Ma) was markedby relatively high benthic δ 1 8 θ values, although the val-ues were lower than those of the late Miocene. This indi-cates warmer intermediate water and less continental iceduring the earliest Miocene than during the late Mio-cene. A slow, general warming of intermediate water oc-curred between 23 and 19.5 Ma, leading to the lowestδ 1 8 θ values for the entire Neogene between about 19.5and 16.5 Ma (late early Miocene; late NN2 to middleNN5). This interval represents the optimum of Neogenewarmth. Haq (1980) identified a particularly warm in-terval between 17 and 16 Ma based on mid- to high-lati-tude Atlantic calcareous nannoplankton biogeography.A minor, cool episode occurred within this warm periodbetween 19 and 18 Ma.

Middle Miocene

The late early Miocene climatic optimum at 16.5 Maimmediately preceded the major middle Miocene episodeof increasing δ 1 8 θ values, which I interpret to representmajor, permanent accumulation of the East Antarcticice sheet and cooling of bottom waters (Shackleton andKennett, 1975b; Savin et al., 1975). Thus, sea-surfacetemperatures at high latitude were relatively warm im-mediately preceding the buildup of the Antarctic ice sheet,a necessary prerequisite according to Mercer and Suter(1982).

The early middle Miocene δ 1 8 θ increase began about16.5 to 16 Ma, and continued until about 13.5 to 13.2Ma; the total duration of this event was about 3 m.y.Barron et al. (in press) previously assigned an age of15.0 to 13.5 Ma to this shift, which encompasses ZonesN9 to lower N12. The maximum enrichment in 18O dur-ing this event ranged between 1.25 and 1.5O%o. This 18Oenrichment event has been globally recognized in marinesedimentary sections by Savin et al. (1975); Shackletonand Kennett (1975b), Keigwin (1979), Margolis et al.(1975), Loutit (1981), Loutit, Kennett, et al. (1983); Vin-cent and Berger (1985); Vincent et al. (in press); Savin etal. (1981), Woodruff et al. (1981), Savin et al. (in press),and many other workers. Haq (1980) identified a cool-ing trend in mid- to high-latitude Atlantic calcareousnannoplankton assemblages at 15 Ma that he correlatedwith this isotopic event.

The middle Miocene event has generally been accept-ed as representing permanent, major buildup of the EastAntarctic ice sheet (Savin et al., 1975; Shackleton andKennett, 1975b) and an associated decrease in bottom-water temperatures related to enhance polar glaciation(Savin et al., 1975). This interpretation was challenged

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by Matthews and Poore (1980), who suggested that allof the isotopic change was due to the decreasing temper-ature of intermediate and bottom waters. They attrib-uted none of the isotopic shift to ice accumulation, sincethey believed that most of the Antarctic ice sheet hadformed by at least the earliest Oligocene, and used assupporting evidence the major sea-level fall during themiddle Oligocene (Vail and Hardenbol, 1979).

Tropical planktonic δ 1 8 θ values decrease, show nochange, or exhibit an increase that is smaller in magni-tude than the shift observed in high-latitude sequences.Savin et al. (1981, in press), Loutit, Kennett, et al. (1983)and others interpreted these observations as reflecting awarming of tropical sea-surface temperatures. Warmingof tropical surface waters during the middle Miocenemay have resulted from the progressive thermal isolationof Antarctica and the contraction of the tropical-warmsubtropical gyres toward lower latitudes, as the polar andsubpolar waters expanded toward the equator. A majorquestion remains, however, as to how much of the δ 1 8 θshift during the early middle Miocene was due to in-creased ice volume and how much to decreased tempera-ture.

Comparison of planktonic and benthic δ 1 8 θ recordsat Site 590 (Figs. 6 and 7) may assist in separating theseeffects, although interpretations are not unequivocal. Site590 is currently located at 31 °S, in proximity to the Sub-tropical Divergence that separates the warm-subtropicaland temperate water masses. During the early middleMiocene, this site was located in the present temperatelatitudes of about 34°S (Sclater et al., in press). At Site590, the shift toward higher benthic δ 1 8 θ values beganat 16 Ma (Figs. 6 and 7) and continued until 13.2 Ma,with a total shift of about 1.25%o. In contrast, the plank-tonic δ 1 8 θ values actually decreased at 16 Ma by aboutO.25%o and remained low until 14.6 Ma. This indicates adistinct warming of surface waters during the early stageof glaciation. The data of Vincent et al. (in press) fromthe tropical Indian Ocean also exhibit surface-water warm-ing during the first half of the middle Miocene benthicδ 1 8 θ increase. At Site 590, planktonic δ 1 8 θ values in-creased irregularly by about O.65%o between 14.6 and13.2 Ma (Figs. 6 and 7). The magnitude of the plank-tonic δ 1 8 θ change was only half that exhibited by ben-thic δ 1 8 θ. Based upon these observations, we suggestthat the early stage of the middle Miocene δ 1 8 θ shift(between 16 and 14.6 Ma) resulted largely from high-lat-itude cooling, which was followed by a phase in whichthe benthic isotope record largely reflected ice growthbetween 14.6 and 13.2 Ma (or 13.6 Ma, as indicated inSite 588). The smaller magnitude of the increase in plank-tonic δ 1 8 θ than benthic δ 1 8 θ suggests that surface wa-ters warmed at these sites near the northern edge of thetemperate zone during the middle Miocene. Savin et al.(1975), Savin et al. (in press), and Loutit, Kennett, et al.(1983) suggested previously that sea-surface temperaturesin the tropical Pacific warmed during the middle Mio-cene. Assuming an adjustment in ice volume of O.5%oand no major changes in the regional variation of the18O/16O ratio of surface waters, then tropical Pacific sur-

face waters typically warmed by 2 to 5°C during the mid-dle Miocene (Savin et al., in press).

The middle Miocene benthic δ 1 8 θ shift exhibits nu-merous oscillations of up to 0.5‰, although these aresmaller than the l%o changes reported at Site 289 in thewestern tropical Pacific (Woodruff et al., 1981), wheresampling density was higher. A brief but well-definedepisode of more positive δ 1 8 θ values is centered at 15(Site 590) or 15.5 Ma (Sites 588 and 591), within themain δ 1 8 θ shift, and may be equivalent to the same epi-sode dated at 15 Ma by Savin et al. (1981).

The isotopic data show two more episodes of espe-cially cool surface and deep water close to the middle/late Miocene boundary. The earliest of these, between12.5 and 11.5 Ma, seems to be equivalent to that datedby Savin et al. (1981) between 13.5 and 12.5 Ma. Thisevent apparently correlates with the early part of paleo-magnetic Chron 11 (-11.7 Ma) according to Burckle etal. (1982). Following a period of warming, a second de-cline of isotopic temperatures (not well defined in Site590) occurred between 11 and 9 Ma (earliest late Mio-cene), but this event was dated by Savin et al. (1981) asoccurring between 11.8 and 10.5 Ma. This event wascorrelated with the early part of paleomagnetic Chron10 (11.4-11 Ma) by Burckle et al. (1982). Some of thehighest δ 1 8 θ values (coldest isotopic temperatures) forthe entire Miocene occur within this interval.

Late Miocene

The late Miocene was marked by large, but variablebenthic δ 1 8 θ values reflecting a period of prolonged coolclimates. Intermediate waters were coldest near the be-ginning (11 to 9 Ma) and the end of the late Miocene(-6.5 to 4.5 Ma).

During much of the middle late Miocene, betweenabout 9 and 6.5 Ma, average δ 1 8 θ values were slightlylower, indicating a general period of relative climaticwarmth (this is not observed at Site 590). Haq (1980) al-so noted warm climate conditions during the middle lateMiocene, between 9 and 7 Ma, based upon Atlantic cal-careous nannoplankton biogeography.

Further cooling occurred during the latest Miocene,beginning at about 6.2 Ma (shown as 5.5 Ma in Site590), and continued until about 5 to 4.5 Ma. Similarhigh δ 1 8 θ values were reported by Shackleton and Ken-nett (1975a) from Site 284, and interpreted to representa major expansion of the Antarctic ice sheet. This ter-minal Miocene event was marked by widespread cool-ing. Moderate to severe cooling of the ocean surface inmiddle to high latitudes is well documented over wideareas during the latest Miocene (Ingle, 1967, 1973; Kel-ler, 1979; Kennett, 1967; Kennett and Vella, 1975; Loutitand Kennett, 1979; Bandy et al., 1971; Cita and Ryan,1979; Poore, 1981; Haq, 1980). In Argentina, Patagoni-an glaciers extended beyond the Andean Mountain front,apparently for the first time, between about 7 and 5.2Ma (Mercer and Sutter, 1982). In Sites 588 (Fig. 3) and591 (Fig. 9), a number of brief intervals of high δ 1 8 θvalues between 5.4 and 4.9 Ma are interpreted to repre-sent further increases in Antarctic ice volume (Hodell,

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Elmstrom, and Kennett, unpublished data). The termi-nal Miocene event was also marked by marine regression(Kennett, 1967; Adams et al., 1977), growth of the WestAntarctic ice sheet (Ciesielski et al., 1982), intensifica-tion of oceanic circulation (Brewster, 1980; Savin et al.,in press), high sediment accumulation rates (Davies etal., 1977), and the Messinian Salinity Crisis (Hsu et al.,1973; Cita, 1976). The late Miocene was also the time ofthe distinct latest Miocene carbon isotope shift (Keig-win, 1979).

During the late Miocene until 6.2 Ma, planktonic δ 1 8 θvalues exhibited large, rapid oscillations about an un-changing average (Figs. 6 and 7). At 6.2 Ma, a markedincrease in δ 1 8 θ of about O.5%o coincided exactly withthe late Miocene Chron 6 carbon shift. This representssignificant and rapid cooling of surface waters. In Site588, the carbon shift also coincided with a shift towardhigher benthic δ 1 8 θ values.

Pliocene

During the early Pliocene, benthic δ 1 8 θ values oscil-lated about an average similar to that of the latest Mio-cene. At about 3.4 Ma, a 0.4‰ increase (Elmstrom andKennett, this volume) is documented in the benthic δ 1 8 θrecord of Site 590 (Figs. 6 and 7). An oxygen isotopeshift similar to this has been reported from the equatori-al Pacific (Shackleton and Opdyke, 1977; Keigwin, 1979;Prell, 1984), the South Atlantic (Weissert et al., 1984),the North Atlantic (Shackleton and Cita, 1979), and theMediterranean (Keigwin and Thunell, 1980). This shiftalmost certainly documents a cooling of intermediatewaters (Prell, 1984), but possibly also some additionalgrowth of Antarctic ice (Weissert et al., 1984). Increasedglaciation of Antarctica at about 3.5 Ma created a north-ward shift in the deposition of circum-Antarctic glacialmarine sediments (Ciesielski and Weaver, 1974). A gla-cial increase at about 3.5 Ma is also consistent with theobservation by Keany (1978) of cooling of the SouthernOcean at 3.6 Ma, with cooling in temperate areas (Ken-nett and Vella, 1975; Kennett et al., 1979), and with in-creased Patagonian glaciation (Mercer, 1976). Significantsediment ice-rafting in the North Atlantic did not seemto occur until 2.5 Ma. Evidence of ice-rafted materialprior to 2.4 Ma is negligible and may be due to glacia-tion unrelated to major Northern Hemisphere ice sheets(Backman, 1979; Shackleton et al., 1984).

The planktonic δ 1 8 θ record during the Pliocene is dis-tinctly different from the benthic δ 1 8 θ record. First, theplanktonic isotopic record is more variable, reflectingrapidly changing surface-water conditions. Secondly,planktonic and benthic records are not parallel. Higherplanktonic δ 1 8 θ values, reflecting cooler surface waters,occurred between 4.6 and 3.6 Ma. In contrast, benthicδ 1 8 θ values show no such trend during this interval. Thetransition from warm early Pliocene to cooler middlePliocene intermediate waters is not reflected in theplanktonic δ 1 8 θ record from Site 590. From 3.4 to 2.2Ma, the average δ 1 8 θ values in the planktonic recordwere lower (Fig. 6), whereas benthic δ 1 8 θ values were in-creasing. Weissert et al. (1984) also noted this for se-quences at similar latitudes (~ 30°S) in the South Atlan-

tic. I concur with their interpretation that surface waterswarmed by at least 2°C at these southern latitudes, whenhigh-latitude glaciers were growing and bottom waterswere cooling. Between 3.6 and 2.2 Ma at Site 590, a pos-itive shift of O.4%o in benthic δ 1 8 θ generally coincidedwith a negative shift of about O.4%o in planktonic δ 1 8 θ.Surface waters could have warmed by as much as 3°C ifit is assumed that the oxygen isotope composition ofocean waters increased by O.4%o because of ice-volumeaccumulation. However, since the δ 1 8 θ change in ben-thics was almost certainly caused, in part, by cooling ofintermediate waters, the increase in surface waters wouldhave probably been closer to 2°C. Surface waters seemto have warmed over wide latitudinal areas extendingfrom equatorial regions (Vergnaud Grazzini, 1980) tothe warm subtropics (Weissert et al., 1984) between 3.4and 2.2 Ma.

A further increase of about O.4%o in benthic δ 1 8 θ oc-curred in the late Pliocene between 2.6 and 2.4 Ma. Thisisotopic shift heralds the beginning of Northern Hemi-sphere ice accumulation, as discussed previously (Back-man, 1979; Prell, 1984; Shackleton et al., 1984).

The late Pliocene oxygen isotope trends are some-what comparable to those that occurred during the earlymiddle Miocene between 16 and 14 Ma; that is, they in-dicate a warming of surface waters at these middle lati-tudes at times of major ice accumulation at high lati-tudes.

It is perhaps significant that the transitions betweenthe early and middle Miocene, the middle and late Mio-cene, the Miocene and Pliocene, and the preglacial andpostglacial Pliocene were all marked by episodes of dis-tinct climatic cooling. These stratigraphic boundarieswere initially recognized in the classic, shallow-marinesequences exposed on land, on the basis of changes inlithofacies and shallow-water fossils, all susceptible tosea-level change (Loutit and Kennett, 1981) and large-scale climatic change.

Water Column Thermal Gradients

The difference between benthic and planktonic δ 1 8 θvalues in the same samples (Δ18O benthics - plankton-ics) is a measure of the temperature difference betweensurface and bottom waters, or the vertical temperaturegradient. For Site 590 (Figs. 11 and 12), the average ver-tical δ 1 8 θ gradient changed from about 1.75 during theearly Miocene to about 2.25 during the middle and lateMiocene and for much of the early Pliocene (a shift ofO.5%o) This is equivalent to about a 2°C increase in thevertical temperature gradient between 1300 m and theocean surface (assuming no salinity change occurred insurface waters). This increase in the southwest Pacificvertical temperature gradient occurred between 16 and15 Ma, during the early middle Miocene (Fig. 12). Dur-ing the intervals before and after this shift, the averagethermal gradient did not change significantly, althoughthere were large temporary oscillations, such as between13 and 12 Ma. These results concur with those of Loutit,Kennett, et al. (1983), who suggested that much of thechange that occurred in the thermal gradient in the Mio-cene occurred during the middle Miocene.

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200 -

Figure 11. δ 1 8 θ benthics - planktonics ( = Δ1 8O) for Holes 590A and590B plotted against depth for Site 590 and biostratigraphic zona-tions (see caption for Fig. 5). Δ 1 8O represents the δ 1 8 θ gradient be-tween bottom waters (water depth = 1300 m) and the ocean sur-face at Site 590.

Figure 12. δ 1 8 θ benthics - planktonics ( = Δ1 8O) for Holes 590A and590B plotted against age for Site 590 and biostratigraphic zona-tions (see caption for Fig. 5). Δ 1 8O represents the δ 1 8 θ gradient be-tween bottom waters (water depth = 1300 m) and the ocean sur-face at Site 590. Note unconformity between 16.5 and 18 Ma.

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Further distinct changes in the vertical temperaturegradient occurred during the middle Pliocene at 3.6 Ma,with the average thermal gradient increasing by anotherO.5%o (~2°C), and during the late Pliocene ait 2.4 Ma,with a further shift of 0.25 to O.5%o (~ 1-2°C).

The records therefore document three permanent, rap-id shifts in the vertical temperature gradient of the Neo-gene: (1) between 16 and 15 Ma during the time of ma-jor accumulation of the large, permanent East Antarc-tic ice sheet; (2) at ~ 3.5 Ma during a time of further in-creased polar glaciation especially in the Antarctic; and(3) at —2.4 Ma when Northern Hemisphere ice sheetsbegan to form. There was little change in the averagegradient between these three Neogene episodes. The ver-tical temperature gradient of the open ocean during theNeogene changed permanently only during times of large-scale polar glacial development. As in middle latitudesof the South Atlantic (Hsü et al., 1984), middle-latitudeSouth Pacific locations also record an increase in thevertical temperature gradient during the Neogene.

CARBON ISOTOPES

Work on Leg 90 sites revealed two major characteris-tics of carbon isotope stratigraphy. First, considerablesimilarity exists among the δ1 3C stratigraphies of all threesites and with those of other regions, especially with re-spect to the longer-term oscillations (Figs. 3, 6, and 9).Second, the planktonic and benthic δ1 3C changes are al-most completely parallel, as shown in Site 590 (Fig. 6).This parallelism indicates that both surface and bottomwaters are equally affected by changes in δ1 3C of oceanicHCO 3, although average planktonic δ1 3C values are al-ways about 1.25‰ higher than benthic δ 1 3C. The simi-larity of δ1 3C stratigraphies in the Neogene between dif-ferent sites throughout the ocean basins has been rec-ognized previously (Douglas and Savin, 1973; Loutit,Pisias, et al., 1983; Woodruff and Savin, 1985; Vincentand Berger, 1985). Shackleton and Kennett (1975b) noteda strong parallelism between South Pacific benthic andplanktonic δ1 3C records during the middle and late Ce-nozoic, as did Hsü et al. (1984) for the South Atlantic,Barrera et al. (in press) for the northeast Pacific, Vin-cent et al. (in press) for the tropical Indian Ocean, andKeigwin et al. (in press) for the North Atlantic. The sim-ilarity of the δ1 3C time-series records within and be-tween ocean basins and between surface and bottom wa-ters clearly indicates that oceanwide changes in averageδ1 3C of HCO 3 in seawater, rather than local effects,dominated the records.

A major question concerns the origin of the changesin the benthic and planktonic δ1 3C records. Proposedcauses of δ1 3C change through time include changes insea level, continental hypsometry, continental biomass,vital effects, and oceanic circulation (Shackleton, 1977;Bender and Keigwin, 1979; Broecker, 1982; Loutit, Pisi-as, et al., 1983; Vincent et al., in press). Summarizing,Berger (1982) noted that the average oceanic δ1 3C com-position will change as a result of several processes: theaccumulation or release of organic carbon from tempo-rary reservoirs on the continents and in the ocean ba-

sins; changes in forest biomass; changes in ocean strati-fication and fertility; and changes in deep circulationand oxygenation. Oceanic δ1 3C values can be changedmost rapidly by changing the flux of carbon from majorcarbon reservoirs: terrigenous carbon having δ1 3C val-ues of about — 25%o and marine carbon near - 2 0 ‰(Craig, 1953; Sackett, 1964; Wedepohl, 1969). Broecker(1971) suggested that transgressions and regressions acrossthe continental shelf could produce significant changesin open-ocean δ1 3C values. During transgressions, theshelves act as sinks for organic carbon. Conversely, dur-ing regressions, eroded sediment containing organic ma-terial depleted in 1 3C is supplied to the oceans.

My preferred interpretation of the Neogene δ13C chang-es is that they were largely caused by average oceaniccompositional changes resulting from fractionation be-tween continental and oceanic organic carbon reservoirs.What caused this fractionation? Loutit, Pisias, et al.(1983) and Woodruff and Savin (1985) have shown thata correlation exists between changes in Neogene carbonisotopic stratigraphy and sea level as inferred from theonlap/offlap curve of Vail and Hardenbol (1979). Dur-ing marine transgressions, organic carbon accumulatedon the continental shelves, and ocean waters becamemore enriched in 1 3C by the extraction of organic car-bon (enriched in 1 2C). Conversely, during marine regres-sion, organic carbon was delivered to the ocean basinsand thus the oceans became more depleted in 1 3 C. Thetransgressions and regressions were probably of largemagnitude to account for the large changes in oceanicδ1 3C values and were controlled by fluctuations in sealevel. The magnitude of the δ1 3C change depended uponthe areas of shallow shelves covered or exposed by agiven change in sea level. This, in turn, is directly relatedto the hypsometry of the continental margins and themagnitude of the sea-level change. A marine transgres-sion that covered wide continental shelves provided alarger area for accumulation of organic carbon, andhence a larger increase in oceanic δ1 3C values.

Paleoclimatic change reinforced the δ13C changes. Glo-bal warming probably enhanced the accumulation oforganic carbon in the continental reservoir by expansionof the biosphere through greater supply of moisture.Conversely, cooling led to greater aridity and reductionin size of the continental biosphere (Berger, 1982).

Superimposed upon the average δ1 3C of the oceansare a number of regional and local effects that have cre-ated gradients from region to region and with waterdepth. For example, Kroopnick (1974, 1980, 1985) doc-umented the fact that different water masses have di-ferent δ1 3C signatures depending upon source and addi-tion of metabolic CO 2 along the flow path. The mecha-nisms that controlled this variation in the past are lesswell understood, but include local and regional differ-ences in biological productivity of surface waters, changesin deep-water circulation patterns, changes in the δ1 3Cof HCO 3 in the source regions of bottom-water forma-tion, and the amount of organic carbon that has beenoxidized to form 12C-enriched HCO 3 in deep waters fol-lowing its formation at high latitudes (Loutit, Pisias, etal., 1983).

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Early Neogene Carbon Isotope Excursion

The dominant feature of Neogene carbon isotopestratigraphy is the broad late early to early middle Mio-cene maximum of about l%o which Vincent et al. (inpress) and Vincent and Berger (1985) have termed the"Monterey Carbon Isotope Excursion." This excursionis best exhibited at Site 588 (at Site 590, part of the ex-cursion is missing because of the hiatus), where it beganas a shift toward increased δ1 3C values at about 19.5 Ma(within Zone NN2). The excursion consists of broad peakscentered at 19 Ma, 16.5 Ma, and about 15 to 15.5 Ma.As shown in Sites 588 and 590, the decline from the highvalues began within the middle Miocene between 15 and14 Ma and continued with little interruption until 13Ma. This negative shift began about midway throughthe distinct, early middle Miocene increase in benthicδ 1 8 θ values, and is contemporaneous with the trend to-ward more positive planktonic δ 1 8 θ values (Figs. 3 and6).

As several previous workers have proposed, I concurthat the origin of this carbon isotopic excursion was dueto global marine transgression (Loutit, Pisias, et al.,1983; Vincent et al., in press; Woodruff and Savin, 1985).According to Vincent et al. (in press) and Vincent andBerger (1985) widespread transgression during this ex-cursion occurred throughout the circum-Pacific regionand created major sinks for organic carbon. Organic andphosphate-rich deposits accumulated within the Monte-rey Formation of California and other contemporane-ous deposits throughout the Pacific margin. Vail andHardenbol (1979) documented a period of coastal onlapthat began at about 19 Ma and terminated about 14 Ma,and was contemporaneous with the interval containingthe especially positive δ1 3C values (Figs. 3, 6, and 9).

This excursion also corresponded to the maximumocean temperatures of the Neogene as reflected in ben-thic and planktonic δ 1 8 θ values. Because the late earlyMiocene was the time of optimum climatic warming dur-ing the Neogene, tropical forests were probably wide-spread and the continental organic biosphere would havebeen more extensive than at other times. The high conti-nental biomass at this time would have contributed tothe positive δ13C values during this interval. Global cool-ing during the middle Miocene probably led to greateraridity and reduction of the terrestrial biomass, result-ing in a lowering of oceanic δ1 3C values during the mid-dle Miocene.

At Site 590, the warming of surface waters during theearly stage of the benthic δ 1 8 θ shift suggests that theearly part of that shift may have resulted largely fromdecreasing temperatures at high latitudes rather than fromincreasing ice accumulation. If little ice had accumu-lated at that point, then glacioeustatic regression wouldnot have occurred. The δ1 3C (benthic and planktonic)began to decrease, suggesting regression, at about 14.5Ma, when the δ 1 8 θ shift occurred in both the benthicand planktonic records. It is possible that this synchro-nous enrichment in both planktonic and benthic δ 1 8 θrepresents the major ice-growth phase of East Antarcti-ca. If this scenario is correct, then major ice accumula-

tion and resulting regression occurred between 14.5 Maand 13 Ma.

Following the carbon isotope excursion, during thelatest middle Miocene, further changes in δ13C occurred,including a trend toward lower values that was centeredin the earliest late Miocene at 11 Ma. A return to mark-edly increased benthic δ1 3C values occurred during themiddle part of the late Miocene, especially between 10and 7 Ma.

Latest Miocene Carbon Isotope Shift

The latest Miocene at 6.2 Ma is marked by a distinct,rapid δ1 3C shift of up to - 0.75‰ in all of the southwestPacific sites, both in the benthic and planktonic records.This depletion in 1 3C was followed by markedly low val-ues during much of the remainder of the late Neogene,which lead to the suggestion that it was "permanent"(Keigwin, 1979). This well-known late Miocene carbonisotope shift is recorded throughout the Indo-Pacific andSouth Atlantic (Keigwin, 1979; Bender and Keigwin, 1979;Loutit and Kennett, 1979; Keigwin and Shackleton, 1980;Vincent et al., 1980; Haq et al., 1980; Hodell and Ken-nett, 1984). This shift, which seems to have occurredwithin a period of not much more than 100,000 yr. (Vin-cent et al., 1980), was assigned an age of ~ 6 Ma by Haqet al. (1980) and of 6.2 Ma by Loutit and Kennett (1979),and was recently estimated at 6.4 Ma (Barron et al., inpress). The magnitude of the shift throughout the Indo-Pacific region is generally reported to be about - 1 ‰(Haq et al., 1980); in the South Atlantic a shift of 0.7%0

has been estimated by Hodell, Elmstrom, and Kennett(unpublished data). Numerous climatic and oceanographicevents closely associated with the carbon shift have beensummarized by Loutit and Keigwin (1982) (for addition-al references see therein) and include the following: low-ering of sea level; cooling of high- and middle-latitudesurface waters; isolation and dessication of the Mediter-ranean Basin; increase in bottom-water circulation andfertility of the oceans: increase in biogenic siliceous de-posits in the Southern Ocean; decrease of biogenic sili-ceous sedimentation rates in the eastern equatorial Pa-cific; shoaling of the Isthmus of Panama; a change, inthe North Atlantic, from a calcite compensation depth(CCD) shallower than that in the Pacific to one deeperthan the Pacific CCD; and an increase in deep-sea sedi-mentation rates. During the latest Miocene, the CCDexperienced a sharp fall of at least several hundred me-ters (van Andel et al., 1977), or even up to 1 km in theSouth Atlantic (Hsü et al., 1984).

Numerous hypotheses have been proposed for the causeof the carbon isotope shift, including a change in therate of organic carbon buried in the oceans and in therate of upwelling and ocean fertility (Bender and Keig-win, 1979). The carbon shift is believed by some to havebeen caused by an increase in the flux of organic matterfrom coastal lowlands and continental shelves exposedby regression (Loutit and Keigwin, 1982; Vincent et al.,1980; Loutit, Pisias, et al., 1983), reinforced by increasedoceanic fertility (Vincent et al., 1980). The close tempo-ral association of the carbon shift with global regression(Kennett, 1967; Adams et al., 1977; Vail and Harden-

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bol, 1979), which is inferred to have resulted from gla-cioeustatic lowering of sea level (Kennett, 1967; Loutitand Kennett, 1979; Mercer and Sutter, 1982) in the latestMiocene, makes this theory very compelling. Estimatesof the magnitude of the marine regression are similar,ranging from 40 to 80 m (Loutit and Kennett, 1979;Berggren and Haq, 1976; Cita and Ryan, 1979).

The drop in the CCD throughout the oceans has anorigin similar to that of the carbon isotope shift, havingresulted from an increase in the fractionation of carbon-ate and organic carbon entering the deep basins fromthe continental shelves during the latest Miocene regres-sion (Berger, 1970; Ciesielski et al., 1982). This sea-levelfall was also sufficient to isolate the Mediterranean andinitiate the "Messinian Salinity Crisis" (Cita, 1982) be-tween about 5.7 and 5.2 Ma. Previous work has indi-cated that the sea-level fall was largely glacioeustatic be-cause of the rapidity of the fall and its association withsurface-water cooling at middle and high latitudes (Ken-nett, 1967; Kennett and Watkins, 1974; Loutit and Ken-nett, 1979; Loutit, 1981) and with a period of relativelyhigh δ 1 8 θ values (Shackleton and Kennett, 1975a; Ken-nett et al., 1979; Loutit and Kennett, 1979; Ciesielski etal., 1982; McKenzie et al., 1984).

Early Pliocene Events

During the Pliocene (as shown by Site 590), averagebenthic δ13C values were lower by about 0.25‰, plank-tonic δ13C by about O.75%o, than during the Miocenepreceding the carbon shift. An interval of relatively in-creased values occurred in the early Pliocene, betweenabout 5 and 4 Ma. This interval corresponds with a peri-od of relative global warmth (Kennett, 1967; Loutit andKennett, 1979; Ingle, 1967; McKenzie et al., 1984); globalmarine transgression (Kennett, 1967; Kennett and Wat-kins, 1974; Vail and Hardenbol, 1979); relatively lowδ 1 8 θ values in benthic foraminifers as shown at Site 590and in records elsewhere (McKenzie et al., 1984; Hodell,Elmstrom, and Kennett, unpublished data), reflecting adecrease in global ice volume; and a marine transgres-sion in the Mediterranean with restoration of p>elagic sed-imentation. I suggest that the early Pliocene interval ofhigh δ13C values resulted from marine transgression. De-tails for this interval are provided in Elmstrom and Ken-nett (this volume).

CONCLUSIONS

1. A high-resolution study of benthic and plank-tonic oxygen and carbon isotope stratigraphy was con-ducted for the Miocene to early Pliocene in high-qualityDSDP Sites 588, 590, and 591, from the southwest Pa-cific (warm-subtropical to transitional latitudes). Sam-ple resolution for much of the Miocene is between-25,000 and 50,000 yr. and for the Pliocene between5,000 and 15,000 yr.

2. The climax of Neogene warmth, marked by thelowest δ 1 8 θ values of the Neogene, occurred during thelate early Miocene, between ~ 19.5 and 16.5 Ma.

3. This climatic warming was immediately followedby a major increase in δ 1 8 θ values between -16.5 and13.5 Ma. This event is interpreted as representing major,

permanent accumulation of the East Antarctic ice sheetand cooling of bottom waters; it lasted 3 m.y. Surfacewaters at these middle latitudes warmed during the firsthalf of the ice volume increase, but cooled at higher lati-tudes.

4. The late Miocene began, in the earliest part (11to 9 Ma), with a distinct climatic cooling that is chroni-cled by the highest δ 1 8 θ for the Miocene. This was fol-lowed by relative warmth during the middle part of thelate Miocene, between - 9 and 6.5 Ma.

5. The latest Miocene and early Pliocene (6.2 to 4.5Ma) were marked by increased δ 1 8 θ values, indicatingincreased cooling and glacial development in the South-ern Hemisphere.

6. During the middle Pliocene at about 3.4 Ma, aO.4%o increase in benthic δ 1 8 θ documents a net changein average global ice volume and cooling of bottom wa-ters. During this increase in Southern Hemisphere glaci-ation, surface waters warmed by 2-3 °C in middle-lati-tude regions of the Southern Hemisphere.

7. During the late Pliocene between 2.6 and 2.4, afurther increase in benthic δ 1 8 θ heralded the develop-ment of large ice sheets on the Northern Hemisphere.

8. Surface-water warming in the middle latitudes oc-curred in association with major high-latitude glacial in-creases in the early middle Miocene (16-14 Ma), middlePliocene (~3.5 Ma), and late Pliocene (-2.4 Ma). Theseintervals were also marked by increases in the verticaltemperature gradient in the open ocean. At interveningtimes, the gradient remained unchanged.

9. Transitions between the early and middle Mio-cene, the middle and late Miocene, the Miocene and Pli-ocene, and the "preglacial" and "postglacial" Pliocenewere marked by distinct episodes of climatic cooling.

10. A high similarity of the δ13C time-series recordswithin and between ocean basins and between surfaceand bottom waters clearly indicates that ocean widechanges in the average δ13C of HCO 3 in seawater domi-nated the records, rather than local effects. Broad chang-es in the Neogene δ13C record were caused largely bytransfers of organic carbon between continental and oce-anic reservoirs. These transfers resulted from marine re-gressions and transgressions on the continental margins.During transgressions, organic carbon, which has lowδ13C values, accumulated on the continental shelves, andocean waters thus increased in δ13C. Conversely, duringmarine regressions, organic carbon was delivered to theocean basins and the oceans became lower in δ13C. Pa-leoclimatic changes reinforced these δ13C changes by in-fluencing the size of continental biomass.

11. The dominant feature of Neogene δ13C stratigra-phy is a broad maximum of about l%o between —19and 14.5 Ma. This excursion occurred contemporane-ously with a period of maximum coastal onlap (trans-gression) and maximum Neogene climatic warmth. Theexcursion terminated during the expansion of the Ant-arctic ice sheet and associated marine regression in theearly middle Miocene (15.5-14.5 Ma).

12. The latest Miocene carbon isotope shift (of up to- O.75%o) at 6.2 Ma is clearly recorded in the sites exam-ined, and was followed by generally low δ13C values dur-

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ing the remainder of the Neogene. This shift may havebeen caused by an increase in the flux of organic matterfrom continental marginal areas exposed by regression,which was, in turn, triggered by glacioeustatic loweringof the sea level.

13. An increase in δ1 3C values during the early Plio-cene (~ 5 to 4 Ma) resulted from marine transgressionduring a time of global warmth.

ACKNOWLEDGMENTS

I wish to thank J. Orchardo for his meticulous care in producingthe stable isotope data. Valuable conversations about this work withD. Hodell and N. Shackleton are appreciated. I thank D. Hodell forhis most valuable constructive criticism of the manuscript, which wastyped by N. Meader. The manuscript also benefited from constructivereviews by L. Keigwin, S. Savin, T. Loutit, and N. Shackleton. Thanksare also due to R. Ross, K. Nelson, and R. Rice for picking foramini-fers for the analyses and to K. Nelson for producing the data plots andfor general assistance. J. Peck also lent useful general assistance.

This research was supported by the U.S. National Science Founda-tion Grant No. OCE82-14937 (Submarine Geology and Geophysics).

REFERENCES

Adams, C. G., Benson, R. H., Kidd, R. B., Ryan, W. B. R, andWright, R. C , 1977. The Messinian salinity crisis and evidence oflate Miocene eustatic changes in the world ocean. Nature, 269:383-386.

Backman, J., 1979. Pliocene biostratigraphy of DSDP Sites 111 and116 from the North Atlantic Ocean and the age of Northern Hemi-sphere glaciation. Stockholm Contrib. GeoL, 32(3):115-135.

Bandy, O. L., Casey, R. E., and Wright, R. C , 1971. Late NeogenePlanktonic Zonation, Magnetic Reversals, and Radiometric Dates,Antarctic to the Tropics. Am. Geophys. Un. Antarctic Sen, 15.

Barrera, E., Keller, G., and Savin, S. M., in press. Evolution of theMiocene Ocean in the eastern North Pacific as inferred from oxy-gen and carbon isotopic ratios of foraminifera. In Kennett, J. P.(Ed.), The Miocene Ocean: Paleoceanography and Biogeography.Geol. Soc. Am. Mem., 163.

Barron, J. A., Keller, G., and Dunn, D. A., in press. A multiple mi-crofossil biochronology for the Miocene. In Kennett, J. P. (Ed.),The Miocene Ocean: Paleoceanography and Biogeography. Geol.Soc. Am. Mem., 163.

Bender, M. L., and Keigwin, L. D., Jr., 1979. Speculations about theupper Miocene change in abyssal Pacific dissolved bicarbonate δ 1 3C.Earth Planet. Sci. Lett., 45:383-393.

Berger, W. H., 1970. Planktonic Foraminifera in selective solution andthe lysocline. Mar. Geol., 8:111-183.

, 1982. Deep-sea stratigraphy: Cenozoic climate steps and thesearch for chemo-climatic feedback. In Einsele, G., and Seilacher,A. (Eds.), Cyclic and Event Stratification: Berlin (Springer-Ver-lag), pp. 121-157.

Berger, W. H., Killingley, J. S., and Vincent, E., 1978. Stable isotopesin deep-sea carbonates: Box Core ERDC-92, West Equatorial Pa-cific. Oceanol. Acta, 1:203-216.

Berggren, W. A., 1981. Correlation of Atlantic, Mediterranean andIndo-Pacific Neogene stratigraphies: geochronology and chrono-stratigraphy. Proc, IGCP-114 Int. Workshop Pacific Neogene Bio-stratigraphy, Nov. 25-29, 1981, Osaka, Japan, pp. 29-60.

Berggren, W. A., and Haq, B. U , 1976. The Andalusian Stage (lateMiocene): biostratigraphy, biochronology, and paleoecology. Pa-laeogeogr., PalaeoclimatoL, Palaeoecol., 20:67-129.

Berggren, W. A., Kent, D. V, and Flynn, J. J., 1985. Paleogene geo-chronology and chronostratigraphy. In Snelling, N. J. (Ed.), Geo-chronology and the Geological Record. Geol. Soc. London, Spec.Pap.

Berggren, W. A., Kent, D. V, and Van Couvering, J. A., 1985. Neo-gene geochronology and chronostratigraphy. In Snelling, N. J.(Ed.), Geochronology and the Geologic Time Scale. Geol. Soc.London, Spec. Pap.

Brewster, N. A., 1980. Cenozoic biogenic silica sedimentation in theAntarctic Ocean, based on two Deep Sea Drilling Project sites.Geol. Soc. Am. Bull., 91:337-347.

Broecker, W. S., 1971. A kinetic model for the chemical compositionof seawater. Quat. Res., 1:188-207.

, 1982. Ocean chemistry during glacial time. Geochim. Cos-mochim. Acta, 46:1689-1705.

Burckle, L. H., Keigwin, L. D., and Opdyke, N. D., 1982. Middleand late Miocene stable isotope stratigraphy: correlation to the pa-leomagnetic reversal record. Micropaleontology, 28(4):329-334.

Ciesielski, P. F., Ledbetter, M. T., and Ellwood, B. B., 1982. The de-velopment of Antarctic glaciation and the Neogene paleoenviron-ment of the Maurice Ewing Bank. Mar. Geol., 46:1-52.

Ciesielski, P. F., and Weaver, F. M., 1974. Early Pliocene temperaturechanges in the Antarctic seas. Geology, 2:511-515.

Cita, M. B., 1976. Biodynamic effects of the Messinian salinity crisison the evolution of planktonic foraminifers in the Mediterranean.Palaeogeogr., PalaeoclimatoL, Palaeoecol., 20:23-42.

, 1982. The Messinian salinity crisis in the Mediterranean: Areview. Alpine Mediterranean Geodynamics Series (Vol. 7): Wash-ington (Am. Geophys. Un.), 113-140.

Cita, M. B., and Ryan, W. B. R, 1979. Late Neogene environmentalevolution. In von Rad, U., Ryan, W. B. R, et al., Init. Repts.DSDP, 47 (Pt. 1): Washington (U.S. Govt. Printing Office), 447-459.

Craig, H., 1953. The geochemistry of the stable carbon isotopes. Geo-chim. Cosmochim. Acta, 3:53-92.

Davies, T. A., Hay, W. W , Southam, J. R., and Worsley, T. R., 1977.Estimates of Cenozoic oceanic sedimentation rates. Science, 197:53-55.

Denham, R. N., and Crook, F. G., 1976. The Tasman Front. N.Z.J.Mar. Freshwater Res., 10:15-30.

Douglas, R. G., and Savin, S. M., 1973. Oxygen and carbon isotopeanalyses of Cretaceous and Tertiary Foraminifera from the centralnorth Pacific. In Winterer, E. L., Ewing, J. I., et al., Init. Repts.DSDP, 17: Washington (U.S. Govt. Printing Office), 591-605.

Epstein, S., Buchsbaum, R., Lowenstam, H. A., and Urey, H. C ,1953. Revised carbonate-water isotopic temperature scale. Geol.Soc. Am. Bull., 64:1315-1326.

Haq, B. U , 1980. Biogeographic history of Miocene calcareous nan-noplankton and paleoceanography of the Atlantic Ocean. Micro-paleontology, 26(4) :414-443.

Haq, B. U , Worsley, T. R., Burckle, L. H., Douglas, R. G., Keigwin,L. D., Jr., Opdyke, N. D., Savin, S. M., Sommer, M. A. II, Vin-cent, E., and Woodruff, R, 1980. Late Miocene marine carbon-isotopic shift and synchroneity of some phytoplanktonic biostrati-graphic events. Geology, 8:427-431.

Hodell, D., and Kennett, J. P., 1984. Late Miocene carbon shift inDSDP Site 516A, western south Atlantic. GSA Abstr. Progr., 16(6):540.

Hsü, K. J., Cita, M. B., and Ryan, W. B. R, 1973. The origin of theMediterranean evaporites. In Ryan, W. B. R, Hsü, K. J., et al., In-it. Repts. DSDP, 13(Pt. 2): Washington (U.S. Govt. Printing Of-fice), 1203-1231.

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.

Ingle, J. C , 1967. Foraminiferal biofacies variation and the Miocene-Pliocene boundary in Southern California. Bull. Am. Paleontol.,52:217-394.

, 1973. Neogene foraminifera from the northeastern PacificOcean, Leg 18, Deep Sea Drilling Project. In Kulm, L. D., vonHuene, R., et al., Init. Repts. DSDP, 18: Washington (U.S. Govt.Printing Office), 517-567.

Keany, J., 1978. Paleoclimatic trends in early and middle Pliocene deep-sea sediments of the Antarctic. Mar. Micropaleontol., 3:35-49.

Keigwin, L. D., Jr., 1976. Late Cenozoic planktonic foraminiferal bio-stratigraphy and paleoceanography of the Panama Basin. Micro-paleontology, 22:419-440.

, 1979. Late Cenozoic stable isotope stratigraphy and pale-oceanography of DSDP sites from the east equatorial and centralnorth Pacific Ocean. Earth Planet. Sci. Lett., 45:361-382.

Keigwin, L. D., Aubry, M.-P., and Kent, D. V, in press. Upper Mio-cene stable isotopic stratigraphy, biostratigraphy, and magnetostra-tigraphy of North Atlantic DSDP Sites. In Ruddiman, W. R, Kidd,R. B., et al., Init. Repts. DSDP, 94: Washington (U.S. Govt. Print-ing Office).

1404


Keigwin, L. D., Jr., and Shackleton, N. J., 1980. Uppermost Miocenecarbon isotope stratigraphy of a piston core in the equatorial Pa-cific. Nature, 284(5757):613-614.

Keigwin, L. D., Jr., and Thunell, R. C , 1980. Middle Pliocene climat-ic change from faunal and oxygen isotopic trends: western Medi-terranean. Nature, 282(5736):294-296.

Keller, G., 1979. Late Neogene planktonic foraminiferal biostratigra-phy and paleoceanography of the Northwest Pacific Site 296. Pa-laeogeogr., Palaeoclimatol., Palaeoecol., 27:129-154.

Kennett, J. P., 1967. Recognition and correlation of the Kapitean Stage(upper Miocene, New Zealand). N.Z.J. Geol. Geophys., 10:1051-1063.

, 1977. Cenozoic evolution of Antarctic glaciation, the cir-cum-Antarctic ocean and their impact on global paleoceanogra-phy. J. Geophys. Res., 82(27):3843-3859.

Kennett, J. P., Keller, G., and Srinivasan, M. S., in press. Mioceneplanktonic foraminiferal biogeography and paleoceanographic de-velopment of the Indo-Pacific region. In Kennett, J. P. (Ed.), TheMiocene Ocean: Paleoceanography and Biogeography. Geol. Soc.Am. Mem., 163.

Kennett, J. P., Shackleton, N. J., Margolis, S. V., Goodney, D. E.,Dudley, W. C , and Kroopnick, P. M., 1979. Late Cenozoic oxygenand carbon isotopic history and volcanic ash stratigraphy: DSDPSite 284, South Pacific. Am. J. Sci., 279:52-69.

Kennett, J. P., and Vella, P., 1975. Late Cenozoic planktonic forami-nifera and paleoceanography at DSDP Site 284 in the cool sub-tropical South Pacific. In Kennett, J. P., Houtz, R. E., et al., Init.Repts. DSDP, 29: Washington (U.S. Govt. Printing Office), 769-799.

Kennett, J. P., and Watkins, N. D., 1974. Late Miocene-early Plio-cene paleomagnetic stratigraphy, paleoclimatology, and biostratig-raphy in New Zealand. Geol. Soc. Am. Bull., 85:1385-1398.

Killingley, J. S., Johnson, R. R, and Berger, W. H., 1981. Oxygen andcarbon isotopes of individual shells of planktonic foraminifera fromOntong-Java plateau, equatorial Pacific. Palaeogeogr., Palaeocli-matol., Palaeoecol., 33:193-204.

Knox, G. A., 1970. Biological oceanography of the South Pacific. InWooster, W. A. (Ed.), Scientific Exploration of the South Pacific:Washington (National Academy of Sciences), pp. 155-182.

Kroopnick, P., 1974. The dissolved O2-CO2-13C system in the eastern

equatorial Pacific. Deep-Sea Res., 21:211-227., 1980. The distribution of δ13C in the Atlantic Ocean. Earth

Planet. Sci. Lett., 49:469-484., 1985. The distribution of 13C of ECO2 in the world oceans.

Deep-Sea Res., 32:57-84.Loutit, T. S., 1981. Late Miocene paleoclimatology: subantarctic wa-

ter mass, southwest Pacific. Mar. Micropaleontol., 6:1-27.Loutit, T. S., and Keigwin, L. D., Jr., 1982. Stable isotopic evidence

for latest Miocene sea-level fall in the Mediterranean region. Na-ture, 300(5888): 163-166.

Loutit, T. S., and Kennett, J. P., 1979. Application of carbon isotopestratigraphy to late Miocene shallow marine sediments, New Zea-land. Science, 204:1196-1199.

, 1981. New Zealand and Australian Cenozoic sedimentarycycles and global sea-level changes. Am. Assoc. Pet. Geol. Bull.,65:1586-1601.

Loutit, T. S., Kennett, J. P., and Savin, S. M., 1983. Miocene equato-rial and southwest Pacific paleoceanography from stable isotopeevidence. Mar. Micropaleontol., 8:215-233.

Loutit, T. S., Pisias, N. G., and Kennett, J. P., 1983. Pacific Miocenecarbon isotope stratigraphy using benthic foraminifera. Earth Plan-et. Sci. Lett., 66:48-62.

McKenzie, J. A., Weissert, H., Poore, R. Z., Wright, R. C , Percival,S. E, Jr., Oberhànsli, H., and Casey, M., 1984. Paleoceanographicimplications of stable-isotope data from upper Miocene-lower Pli-ocene sediments from the southeast Atlantic (Deep Sea DrillingProject Site 519). In Hsü, K. J., LaBrecque, J. L., et al., Init.Repts. DSDP, 73: Washington (U.S. Govt. Printing Office), 717-724.

Margolis, S. V., Kroopnick, P. M., Goodney, D. E., Dudley, W. C ,and Mahoney, M. E., 1975. Oxygen and carbon isotopes from cal-careous nannofossils as paleoceanographic indicators. Science, 189:555-557.

Matthews, R. K., Curry, W. B., Lohmann, K. C , and Sommer, M.A., 1980. Late Miocene palaeo-oceanography of the Atlantic: oxy-

gen isotope data on planktonic and benthic Foraminifera. Nature,283:555-557.

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

Mercer, J. H., 1976. Glacial history of southernmost South America.Quat. Res., 6:125-166.

Mercer, J. H., and Sutter, J. F., 1982. Late Miocene-earliest Plioceneglaciation in southern Argentina: implications for global ice-sheethistory. Palaeogeogr., Palaeoclimatol., Palaeoecol., 38:185-206.

Poore, R. Z., 1981. Late Miocene biogeography and paleoclimatologyof the central North Atlantic. Mar. Micropaleontol., 6:599-616.

Prell, W. L., 1984. Covariance patterns of foraminifera δ 1 8 θ: an eval-uation of Pliocene ice volume changes near 3.2 million years ago.Science, 206:692-693.

Prell, W. L., Gardner, J. V., et al., 1982. Init. Repts. DSDP, 68: Wash-ington (U.S. Govt. Printing Office).

Sackett, W. M., 1964. The depositional history and isotopic organiccarbon composition of marine sediments. Mar. Geol., 2:173-185.

Savin, S. M., Abel, L., Barrera, E., Hodell, D., Keller, G., Kennett, J.P., Killingley, J., Murphy, M., and Vincent, E., in press. The evo-lution of Miocene surface and near-surface marine temperatures:oxygen isotopic evidence. In Kennett, J. P. (Ed.), The MioceneOcean: Palaeoceanography and Biogeography. Geol. Soc. Am.Mem., 163.

Savin, S. M., Douglas, R. G., Keller, G., Killingley, J. S., Shaughnes-sy, L., Sommer, M. A., Vincent, E., and Woodruff, F., 1981. Mio-cene benthic foraminiferal isotope records: A synthesis. Mar. Mi-cropaleontol. , 6:423-450.

Savin, S. M., Douglas, R. G., and Stehli, F. G., 1975. Tertiary marinepaleotemperatures. Geol. Soc. Am. Bull., 86:1499-1510.

Schnitker, D., 1980. North Atlantic oceanography as possible cause ofAntarctic glaciation and eutrophication. Nature, 287:615-616.

Sclater, J., Meinke, L., Bennett, A., and Murphy, C , in press. Thedepth of the ocean through the Neogene. In Kennett, J. P. (Ed.),The Miocene Ocean: Paleoceanography and Biogeography. Geol.Soc. Am. Mem., 163.

Shackleton, N. J., 1977. Carbon-13 in Uvigerina: tropical rain foresthistory and the equatorial Pacific carbonate dissolution cycles. InAnderson, N. R., and Malahoff, A. (Eds.), The Fate of Fossil FuelCO2 in the Oceans: New York (Plenum), pp. 401-427.

Shackleton, N. J., Backman, J., Zimmerman, H., et al., 1984. Oxy-gen isotope calibration of the onset of ice rafting and history ofglaciation in the North Atlantic region. Nature, 307(5952):620-623.

Shackleton, N. J., and Cita, M. B., 1979. oxygen and carbon isotopestratigraphy of benthic foraminifers at Site 597: detailed history ofclimatic change during the late Neogene. In von Rad, U., Ryan, W.B. R, et al., Init. Repts. DSDP, 47(Pt.l): Washington (U.S. Govt.Printing Office), 433-445.

Shackleton, N. J., and Kennett, J. P., 1975a. Late Cenozoic oxygenand carbon isotopic changes at DSDP Site 284: Implications forglacial history of the Northern Hemisphere and Antarctica. In Ken-nett, J. P., Houtz, R. E., et al., Init. Repts. DSDP, 29: Washing-ton (U.S. Govt. Printing Office), 801-806.

, 1975b. Paleotemperature history of the Cenozoic and theinitiation of Antarctic glaciation: Oxygen and carbon isotope anal-yses in DSDP Sites 277, 279, and 281. In Kennett, J. P., Houtz, R.E., et al., Init. Repts. DSDP, 29: Washington (U.S. Govt. PrintingOffice), 743-755.

Shackleton, N. J., and Opdyke, N. D., 1977. Oxygen isotopic and pa-leomagnetic evidence for early Northern Hemisphere glaciation.Nature, 270:216-219.

Stanton, B. R., 1979. The Tasman Front. N.Z.J. Mar. Freshwater Res.,138:201-214.

, 1981. An oceanographic survey of the Tasman Front. N.Z.J.Mar. Freshwater Res., 15:289-297.

Vail, P R . , and Hardenbol, J., 1979. Sea-level changes during the Ter-tiary. Oceanus, 22:71-79.

van Andel, T. H., Thiede, J., Sclater, J. G., and Hay, W. W., 1977.Depositional history of the South Atlantic ocean during the last125 million years. J. Geol., 85:651-698.

Vergnaud Grazzini, C , and Rabussier-Lointier, D., 1980. Essai de cor-relation stratigraphique par le moyen des iotopes de 1'oxygène etdu carbone. Soc. Geol. France, 22(7):719-730.

Vincent, E., and Berger, W. H., 1985. Carbon dioxide and polar cool-ing in the Miocene: The Monterey hypothesis. In Sundquist, E. T.,

1405

J. P. KENNETT

and Broecker, W. S. (Ed.), The Carbon Cycle and AtmosphericCO2: natural Variations Archean to Present: Am. Geophys. UnionMonogr. Ser., 32.

Vincent, E., Killingley, J. S., and Berger, W. H., 1980. The magneticEpoch-6 carbon shift: a change in the ocean's 13C/12C ratio 6.2million years ago. Mar. Micropaleontol., 5:185-203.

, in press. Miocene oxygen and carbon isotope stratigraphyof the tropical Indian Ocean. In Kennett, J. P. (Ed.), The MioceneOcean: Paleoceanography and Biogeography. Geol. Soc. Am.Mem., 163.

Wedepohl, K. H. (Ed.), 1969. Handbook of Geochemistry: New York(Springer-Verlag), p. 11-1(6).

Weissert, H. J., McKenzie, J. A., Wright, R. C , Clark, M., Oberhánsli,H., and Casey, M., 1984. Paleoclimatic record of the Pliocene atDeep Sea Drilling Project Sites 519, 521, 522, and 523 (CentralSouth Atlantic). In Hsü, K. J., LaBrecque, J. L., et al., Init. Repts.DSDP, 73: Washington (U.S. Govt. Printing Office), 701-715.

Woodruff, E, and Douglas, R. G., 1981. Response of deep-sea ben-thic foraminifera to Miocene paleoclimatic events, DSDP Site 289.Mar. Micropaleontol., 6:617-632.

Woodruff, F., and Savin, S. M., 1985. δ13C values of Miocene Pacificbenthic foraminifera: correlations with sea level and biological pro-ductivity. Geology, 13(2):119-122.

Woodruff, E, Savin, S. M., and Douglas, R. G., 1981. Miocene stableisotope record: a detailed deep Pacific Ocean study and its paleo-climatic implications. Science, 212:665-668.

Wyrtki, K., 1974. The Dynamic Topography of the Pacific Ocean andIts Fluctuations. Hawaii Inst. Geophys. Rep. HIG 74-5.

Date of Initial Receipt: 10 May 1985Date of Acceptance: 17 May 1985

APPENDIXOxygen and Carbon Isotope Composition (‰ PDB) of Benthic and

Planktonic Foraminifers, DSDP Sites 588, 590 and 591'

Core-Section(interval in cm)

Hole 588 a

10-1, 35-3710-2, 35-3710-3, 35-3710-4, 35-3710-5, 35-3710-6, 35-3711-1, 35-3711-2, 35-3711-3, 35-3711-4, 35-3711-5, 35-3711-6, 35-3712-1, 35-3712-2, 35-3712-3, 35-3712-4, 35-3712-5, 35-3712-6, 35-3712-7, 35-3713-1, 35-3713-2, 35-3713-3, 35-3713-4, 110-11113-5, 110-11113-6, 35-3713-7, 35-37

SITE 588Sub-bottom

depth(m)

82.884.385.887.388.890.392.493.995.496.998.499.9

101.9103.4104.9106.4107.9109.4110.9111.6113.1114.6116.8118.3119.1120.6

Age(Ma)

4.344.384.424.464.504.544.604.704.794.884.985.075.125.165.205.235.255.285.305.325.345.375.415.435.455.47

δ ^ O

(°/00

1.972.071.931.821.991.901.711.861.912.132.061.971.882.032.132.072.042.042.062.472.282.252.041.792.092.29

δ 1 3 C

PDB)

0.750.810.700.880.830.640.840.850.920.841.040.630.900.810.830.960.980.901.120.730.630.650.690.620.790.56

adbicidoides kullenbergi (benthic) used except in Sample588-13-1, 35 — 37 cm, where C. wuellerstorfi was used.

Stable isotopic data for Oridorsalis umbonatus (Hole 591B) have been adjusted to thatof Cibicidoides by adding + 1.06‰ for δ 1 3 C and subtracting 0.72 for δ 1 8 θ (see text for dis-cussion).


Sub-bottomdepth

(m)

Hole 588 (Cont.)

14-1, 35-3714-2, 35-3714-3, 35-3714-4, 35-3714-5, 35-3714-6, 35-3714-7, 35-3715-1, 35-3715-2, 35-3715-3, 35-3715-4, 35-3715-5, 35-3715-6, 35-3716-1,35-3716-2, 35-3716-3, 35-3716-4, 35-3717-1,35-3717-2, 35-3717-3, 35-3717-4, 35-3717-5, 35-3717-6, 35-3717-7, 35-3718-1,35-3718-2, 35-3718-3, 35-3718-4,35-3718-5, 35-3718-6, 35-3719-1, 50-5119-2, 35-3719-3, 35-3719-4, 35-3719-5, 35-3719-6, 35-3719-7, 35-3720-2, 35-3720-3, 35-3720-4, 35-3720-6, 35-3720-7, 35-3721-1, 35-3721-2, 35-3721-3, 35-3721-4, 35-3721-5, 35-3721-6, 35-3721-7,35-3722-1,35-3722-2, 35-3722-3, 35-3722-4, 35-3722-5, 35-3722-6, 35-3722-7, 35-3723-1,35-3723-2, 35-3723-3, 35-3723-4, 35-2723-5, 35-3723-6, 35-3723-7, 35-3724-1,35-3724-2, 35-3724-3, 35-3724-4, 35-3724-5, 35-3724-6, 35-3724-7, 35-3725-1,35-3725-2, 35-3725-3, 35-3725-4, 35-3725-5, 35-3725-6, 35-3725-7, 35-37

121.1122.6124.1125.6127.1128.6130.1130.8132.3133.8135.3136.8138.3140.4141.9143.4144.9150.0151.5153.0154.5156.0157.5159.0160.0161.0162.5164.0165.5167.0169.3172.1173.6175.1176.6178.1179.6180.3181.8183.3186.3187.8188.4189.9191.4

192.9194.4195.9197.4198.0199.5201.0202.5204.0205.0207.0207.5209.5210.5212.0213.5215.0216.5217.1218.6220.1221.6223.1224.6226.1226.8228.3229.8231.3232.8234.3235.8

Age(Ma)

5.485.515.535.565.585.615.645.655.675.705.735.75 :5.785.815.845.875.895.986.016.03 :6.066.086.116.146.156.176.206.226.256.506.887.367.617.878.12 ;8.37 :8.638.75 :9.009.259.579.649.679.749.8i ;9.889.96 :

10.03 :10.10 :io.i3 ;10.20 ;10.27 ;10.34 ;10.4110.48 ;10.55 :10.57 ;10.67 ;

io.7i ;10.79 ;10.8610.9311.0011.0411.1411.2411.3411.45 ;11.55 111.65 1

(°/oo

2.292.25> . l l.86.96

>.O7.99.97.94

2.052.032.03.99.91

2.002.092.032.012.04>.02.89.78.86

2.052.05.88.77.73.84.75.81.85.79.89

2.012.032.162.13.91.89.85.74.81.85

!.2O2.15>.26>.O7>.09>.26!.192.15!.21.99

..07>.26!.4O>.15!.18'.08.96.79.78.82.79.62.99.13.91.89

11.70 2.0511.80 111.90 112.00 112.04 112.08 112.11 1

.94

.91

.75

.61

.99

.91

δ 1 3 C

PDB)

1.071.100.750.900.820.930.740.750.850.400.770.830.810.820.81

0.670.600.550.910.740.620.821.000.781.021.001.051.061.181.191.241.201.011.231.421.481.451.300.941.021.201.171.431.181.121.311.391.491.211.351.491.411.421.091.101.181.081.091.000.880.720.780.750.800.770.921.131.211.181.341.321.161.181.100.851.071.26

1406



Hole 588Ab

1-1, 35-37

1-2,35-371-3, 35-371-4, 35-371-5, 35-371-6, 35-371-7,35-372-1, 35-372-2, 35-372-3, 35-372-4, 35-373-1, 35-373-2, 35-373-3,35-374-1, 35-374-2, 35-374-3, 35-375-1,35-375-2,35-375-3, 35-376-1, 35-376-2, 35-376-3, 35-376-4, 35-377-1,35-377-2, 35-377-3, 35-377-4, 35-378-1,35-378-2, 35-378-3, 35-378-4, 35-379-1,35-379-2, 35-379-3, 35-3710-1, 35-3710-2, 35-3710-3, 35-3710-3, 35-3710-4, 35-3711-1,35-3711-2,35-3711-3,35-3712-1,35-3712-2, 35-3712-3,35-3712-4, 35-3713-2,35-3713-3, 35-3713-4, 35-3714-1,35-3714-2, 35-3714-3, 35-3715-1, 35-3715-2, 35-3715-3, 35-3715-4, 35-37

Hole 588CC

1-1,35-37

1-2, 35-371-2, 110-1111-3, 35-371-3, 110-1111-5,35-371-5, 110-1111-6,35-371-6, 110-1112-1, 110-1112-2, 35-372-2, 110-1112-3, 35-37

Sub-bottomdepth

(m)

236.4237.9239.4240.9242.4243.9245.4246.0247.5249.0250.5251.0252.5254.0256.0257.5259.0261.0262.5264.0266.0267.5269.0270.5271.0272.5274.0275.5276.0277.5279.0280.5281.0282.5284.0286.0287.5289.0289.0289.5291.0292.5294.0296.0297.5299.0300.5302.5304.0305.5306.0307.5309.0311.0312.5314.0315.5

306.0307.5308.3309.0309.8312.0312.8313.5314.3316.4317.2317.9318.7

Age(Ma)

12.1312.1612.2012.2912.3912.4812.5712.6112.7012.7912.8012.9113.0013.1413.3313.47116113.8013.9314.0714.2614.4014.5414.6814.7314.8715.0015.1215.1615.2715.3915.5115.5515.6615.7815.9316.0516.1716.1716.2016.3316.4616.5816.7516.8817.0017.1017.2417.3417.4417.4717.5717.6717.8118.0718.3318.59

17.4717.5717.6217.6717.7217.9818.1218.2418.3818.7418.8819.0019.05

1 8 n

(°/oo

1.621.501.521.531.461.621.541.731.641.491.741.721.711.881.791.941.971.921.981.921.691.851.711.701.571.621.441.511.251.371.601.611.541.331.131.211.040.910.911.120.870.730.890.930.930.890.951.030.581.141.231.201.180.970.831.171.12

0.721.031.090.711.081.551.140.501.301.131.111.221.41

δ 1 3 Co v~-

PDB)

0.960.950.921.301.391.251.241.361.171.130.920.771.171.211.131.101.161.081.251.031.151.181.261.411.411.501.591.541.481.481.961.981.841.650.891.511.681.651.651.691.711.401.091.431.251.341.441.501.071.781.922.041.861.351.151.241.25

1.321.661.541.341.631.261.891.771.952.032.122.131.67


Sub-bottomdepth(m)

Hole 588C (Cont.)2-3, 110-1112-4, 35-372-4, 110-1112-5, 35-372-5, 110-1112-6, 35-372-6, 110-1113-1,35-373-1, 110-1113-2, 35-373-2, 110-1113-3, 35-373-3, 110-1113-4, 35-373-4, 110-1113-5, 110-1113-6, 35-373-6, 110-1114-2, 110-1114-3, 35-374-3, 110-1114-4, 35-374-5, 35-374-5, 110-1115-1, 35-375-1, 110-1115-2, 35-375-2, 110-1115-3, 110-1115-4, 35-375-5, 35-375-5, 110-1115-6, 110-1116-1,35-376-2, 35-376-3, 110-1116-4, 110-1116-5, 35-376-6, 110-1116-7, 35-377-1,35-377-1, 110-1117-2, 35-377-4, 110-1118-2, 35-378-2, 110-1118-3, 35-378-4, 110-1118-5, 35-378-5, 110-1118-6, 35-378-6, 110-1118-7, 35-379-2, 35-379-3, 35-379-3, 110-1119-4, 35-379-5, 35-379-6, 3 5 - 3 79-6, 110-111

319.4320.2320.9321.7322.4323.2323.9325.3326.0326.8327.5328.3329.0329.8330.5332.8335.6336.3337.1337.8338.6339.3340.8341.6344.5345.2346.0346.7348.2348.9350.4351.2352.7354.0355.5357.8359.3360.0362.3363.0363.7364.4365.2368.9374.8375.5376.3378.5379.3380.0380.8381.5382.3384.3385.8386.6387.3388.8390.3391.1

Age(Ma)

19.0819.1319.1719.2119.2519.3019.3419.4219.4619.5019.5419.5919.6319.6719.7119.8420.0020.0420.0820.1220.1720.2120.2920.3420.5020.5420.5820.6220.7020.7320.8120.8520.9321.0021.0821.1921.2721.3121.4321.4721.5021.5421.5821.7722.0822.1222.1622.2722.3122.3522.3922.4322.4722.5722.6522.6922.7322.8122.8922.93

δ1 8θ δ 1 3 C

(°/oo PDB)

0.850.890.850.920.790.630.730.690.910.720.641.221.201.041.211.261.221.180.971.371.311.151.17

.061.401.301.591.431.301.241.141.151.031.091.581.301.231.361.521.451.311.131.271.111.37

.51

.26

.04

.61

.16

.17

.28

.56

.18

.27

.35

.34

.81

.511.02

1.671.921.661.511.381.321.551.62

1.301.311.201.651.411.021.351.140.981.020.941.041.241.131.000.791.191.171.511.341.220.931.370.710.991.061.271.001.061.101.241.001.070.681.191.361.521.751.941.301.621.311.331.341.541.191.260.991.431.571.570.87

Samples 588A-1-1, 3 5 - 3 7 cm to 588A-10-4, 3 5 - 3 7cm, C. kullenbergi; Samples 588A-11-3, 3 5 - 3 7 cm and588A-11-2, 35-37 cm, C. wuellerstorfi; Samples 588A-1-

J, 35-37 cm to 588A-15-4, 35-37 cm, C. coryelli."All samples, C. coryelli.

Core-Section

(interval in cm)

Hole 590Aa

1-6, 35-372-1, 35-372-3, 35-37

SITE 590Sub-bottom

depth

(m)

34.136.239.2

Age

(Ma)

1.811.902.02

Benthic(°/oo PDB)

δ 1 8 θ

2.592.462.63

δ 1 3 C

0.740.850.64

Planktonic(°/oo PDB)

δ 1 8 0 δ 1 3 c

-0.35 1.90-0.45 0.95-0.49 1.42

All benthic results measured using C. kullenbergi; all planktonic results,Globigerinoides sacculifer.

1407

J. P. KENNETT

Sub-bottomCore-Section

(interval in cm)

Hole 590A (Cont2-5, 35-372-6, 35-373-1, 35-373-2, 35-373-3, 35-373-4, 35-373-5, 35-373-6, 35-374-1, 35-374-2, 35-374-3, 35-374-4, 35-374-5, 35-374-6, 35-375-1, 110-1125-3, 110-1125-5, 110-1126-1, 55-576-3, 55-576-5, 55-577-1, 35-377-3, 35-377-5, 35-378-1, 103-1048-3, 35-378-5, 35-378-6, 35-3710-1, 35-3710-3, 35-3710-5, 35-3711-1, 35-3711-3, 35-3711-4,35-3711-6,35-3712-1, 35-3712-3, 35-3712-5, 35-3712-6, 35-3713-1, 35-3713-3, 35-3713-5, 35-3713-6, 35-3714-1, 35-3714-2, 35-3714-3, 35-3714-4, 35-3714-5, 35-3714-6, 35-3715-1,35-3715-2, 35-3715-3, 35-3715-4, 35-3715-5, 35-3715-6, 35-3716-1, 35-3716-2, 35-3716-3, 35-3716-4, 35-3716-5, 35-3717-1, 35-3717-2, 35-3717-3, 121-12217-5, 35-3718-1, 35-3718-2, 35-3718-4,35-3718-5, 35-3718-6, 35-3719-1, 35-3619-2, 35-3719-3, 35-3719-4, 35-3719-5, 35-3719-6, 35-3720-1, 35-3720-2, 35-3720-4, 35-37

depth

(m)

.)

42.243.745.847.348.850.351.853.355.356.858.359.861.362.864.967.970.974.677.680.684.187.190.193.896.899.8

101.3112.9115.9118.9122.6125.6127.1130.0132.1135.1138.1139.6141.8144.8147.8149.3151.4152.9154.4155.9157.4158.9161.0162.5164.0165.5167.0168.5170.6172.1173.6175.1176.6180.1183.1184.0186.1189.9191.3194.3195.8197.3199.4200.9202.4203.9205.4206.9209.0210.5213.5

Age

(Ma)

2.152.212.302.332.372.402.442.482.532.572.612.652.692.732.792.862.943.043.123.203.293.373.453.473.483.493.503.553.563.583.593.613.613.633.643.653.703.733.773.823.873.903.933.963.994.014.044.064.104.144.184.224.264.304.354.394.434.474.514.604.704.734.804.924.975.075.125.175.225.245.265.285.305.325.355.375.41

Benthic(°/oo

δ1 8o

2.572.523.002.51

—2.60

—

2.392.382.292.412.342.291.972.252.251.981.982.171.872.15

2.012.22

—

2.042.111.981.931.911.89

..

1.92

-–1.841.64

—

1.881.781.87

..1.80

—

-–1.81..

1.76

-–1.76—

1.95-

1.73. .

1.801.921.85

-–1.751.97

--–2.131.931.911.841.791.972.031.831.972.101.861.831.941.761.76

PDB)

δ 1 3 C

0.750.940.440.79

—0.61

. .

0.730.780.730.650.850.560.700.980.810.740.830.760.880.93

..0.870.59

—0.990.800.730.690.760.91

-–0.87—

0.870.75

. .

1.060.940.86

..0.87

—

-–0.77—

0.68—

0.93—

0.63-

1.01..

0.961.001.03

. .0.821.15

-

0.861.091.100.780.880.470.680.840.690.590.880.700.510.660.88

Planktonic(°/oo

δ 1 8 θ

-0.45-0.69-0.57-0.72-0.090.11

-0.55-0.50-0.56-0.55-0.60-0.70-0.48-0.52-0.77-0.31-0.96-0.60-0.98-0.58-0.67-0.39-0.63-0.28-0.35-0.78-0.90-0.74-0.85-0.73-0.46-0.45-0.68-0.85-0.59-0.31-0.54-0.66-0.22-0.32-0.23-0.20-0.31-0.38-0.36-0.53-0.66-0.26-0.29-0.45-0.26-0.29-0.16-0.11-0.31-0.35-0.49-0.22-0.35-0.73-0.802.17

-0.31-0.59-0.62-0.31-0.49-0.13-0.60-0.55-0.37-0.75-0.76-0.67-0.45-0.50-0.69

PDB)

δ 1 3 C

1.501.581.571.661.391.421.551.671.521.741.241.571.441.361.391.561.491.591.501.751.581.451.721.741.341.591.551.371.471.641.481.391.731.371.571.531.711.471.681.521.701.321.541.531.081.120.831.451.581.461.511.351.591.581.751.761.791.881.311.851.670.721.432.012.061.851.831.471.631.931.651.541.782.011.281.592.11

Core-Section

(interval in cm)

Sub-bottomdepth

(m)

Hole 590A (Cont.)20-6, 35-3721-1, 35-3721-3, 35-3721-5, 35-3721-7, 35-3722-2, 35-3723-1, 35-3723-3, 35-3723-5, 35-3724-1, 35-3724-3, 35-3724-5, 35-3724-6, 35-3724-7, 35-3725-1, 121-122

Hole 590Bb

28-6, 35-3729-1, 35-3729-2, 35-3729-3, 35-3729-4, 35-3729-5, 35-3729-6, 35-3730-1, 35-3730-2, 35-3730-3, 35-3730-4, 35-3730-5, 35-3730-6, 35-3731-1,35-3831-2, 35-3731-3, 35-3731-4, 35-3731-5, 35-3732-1, 35-3732-2, 35-3732-3, 35-3732-4, 35-3732-5, 35-3732-6, 35-3733-1, 35-3733-2, 35-3733-3, 35-3733-4, 35-3733-5, 35-3733-6, 35-3733-7, 35-3734-1, 35-3734-2, 35-3734-3, 35-3734-4, 35-3734-5, 35-3734-6, 35-3734,CC35-1, 35-3735-2, 35-3735-3, 35-3735-4, 35-3735-5, 35-3736-1, 35-3736-2, 35-3736-3, 35-3736-4, 35-3736-5, 35-3737-1, 35-3737-2, 35-3737-3, 35-3737-4, 35-3737-5, 35-3737-6, 35-3738-1, 35-3738-2, 35-37

216.5218.6221.6224.6227.6229.7237.8240.8244.6247.4250.4253.4254.9256.4257.0

258.5259.5261.5262.5264.0265.5267.4269.0270.0272.0273.5275.0276.2278.7280.7281.7283.2284.7288.3289.8291.3292.8294.3295.8297.8299.3300.8301.3302.8304.3305.8307.5309.0310.5312.0313.5315.0316.6317.0318.5320.0321.5323.0326.7328.2329.7331.2332.7336.3337.8339.3340.8342.3343.8345.9347.4

Age

(Ma)

5.455.485.525.565.605.635.745.785.835.875.915.965.986.006.00

6.026.046.076.086.106.126.156.176.186.216.236.256.326.486.606.666.756.847.067.157.257.347.437.527.647.737.837.867.958.048.138.248.338.428.518.608.698.798.828.919.009.099.189.419.509.589.669.739.92

10.0010.0810.1610.2310.3110.4210.50

Benthic(°/oo

δ 1 8 θ

1.671.61

.811.801.841.771.871.901.851.96

.84

.871.641.802.02

-1.80

--–1.71

-

1.90---

1.85

.76

.73

.78

.86

.69

.65

.881.771.66

.811.631.661.491.461.551.711.421.691.701.851.781.821.581.681.821.991.811.92

.851.982.001.781.591.751.541.701.681.711.751.931.401.50

PDB)

δ 1 3 C

0.650.800.500.930.710.670.690.750.700.480.450.460.600.540.69

—-–0.77

---

0.79-–0.70

-

-–-

0.82-.02.11

0.931.101.06

.120.930.981.020.971.261.191.281.291.371.321.001.151.181.171.331.240.93

.00

.14

.14

.38

.43

.43

.37

.52

.34

.19

.49

.011.091.111.041.011.120.760.98

Planktonic( % o

δ 1 8 θ

-0.45-0.80-0.64-0.44-0.45-0.55-0.47-0.50-0.72-0.60-0.38-0.38-0.32-0.40-0.13

-0.39-0.29-0.33-0.26-0.55-0.20-0.70-0.19-0.36-0.30-0.53-0.47-0.43-0.56-0.61-0.71-0.43-0.65-0.62-0.34-0.64-0.49-0.60-0.69-0.69-0.67-0.48-0.62-0.74-0.25-0.38-0.71-0.35-0.56-0.62-0.48-0.22-0.17-0.61-0.69-0.43-0.27-0.20-0.53-0.49-0.43-0.32

—

-0.70-0.51-0.74-0.43-0.37-0.52-0.51-0.44

PDB)

δ 1 3 C

1.591.841.591.971.762.011.731.681.571.261.241.621.851.561.92

2.001.941.591.782.131.872.162.092.302.142.122.202.072.112.242.152.112.152.182.232.272.132.492.352.482.142.542.672.572.562.072.002.042.022.292.051.952.201.931.941.502.272.252.292.352.262.22

-1.912.012.141.831.952.201.851.90

All benthic results measured using C. kullenbergi except for samplesmarked with asterisk, where C. coryelli was measured. All planktonicresults, Globigerinoides quadrilobatus.

1408


Core-Section

(interval in cm)

Sub-bottomdepth

(m)

Age

(Ma)

<y>

Benthic(°/oo PDB)

18 0 δ 1 3 C


δ 1 8 o δ 1 3 C

Hole 590B (Cont.)

Core-Section

(interval in cm)

Sub-bottomdepth

(m)

Age

(Ma)δ

Benthic(°/oo PDB)

1 8 Q δ 1 3 C


δ 1 8 θ δ 1 3 C

38-3.38-4,38-5,38-6.38-7,39-1,39-2.39-3.39-4.39-5.39-6.41-1.41-2,41-3,41-4.41-5.41-6,42-1.42-2.42-3.42-4.42-5.42-6.43-1,43-2.43-3.43-4.43-5.44-1.44-2.44-3.44-4.44-5.44-6.

45-1,45-2,45-4,45-5,46-1,46-2,46-3,46-4,46-5,46-6,47-1,47-2,47-3,47-4,47-5,47-6,48-1,48-2,48-3,48-4,48-5,48-6,48-7,

*49-l,49-1,49-2,

*49-2,49-3,

*49-3,49-4,49-4,

*49-5,49-5,

*50-l,*50-3,50-3,

•51-1,•51-2,•51-3,*51-4,•51-4,•51-5,•51-6,

35-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-37

35-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-3735-37

348.9350.4351.9353.4354.9355.5357.0358.5360.0361.5363.0374.7376.2377.7379.2380.7382.2384.3385.8387.3388.8390.3391.8393.8395.3396.8398.3399.8403.5405.0406.5408.0409.5411.0413.0414.5417.5419.0422.7424.5425.7427.2428.7430.2423.3433.8435.3436.8438.3439.8441.8443.3444.8446.3447.8449.7450.8451.5451.5453.0453.0454.5454.5456.0456.0457.5457.5461.0464.0464.0470.7472.2474.0475.2475.2476.7478.2

10.5810.6610.7310.8110.8910.9211.0011.0811.1611.2311.3111.9212.0012.0312.0612.0912.1312.1712.2012.2412.2812.3212.3612.4112.4512.4912.5312.5712.6712.7112.7512.7912.8312.8712.9212.9613.1313.2713.6013.7413.8714.0014.1414.2714.4614.6014.7314.8715.0015.1415.3215.4515.5915.7315.8616.0016.1416.2016.2016.3416.3416.5016.5018.0018.0018.0618.0618.1918.3118.3118.5718.6318.6918.7418.7418.8018.86

1.941.431.801.891.311.831.931.591.721.451.941.591.591.511.651.521.141.861.761.832.171.741.961.341.061.220.961.381.291.321.291.421.481.501.501.331.661.611.281.181.301.401.421.310.931.231.261.411.341.330.980.931.121.000.970.340.900.95

0.820.960.861.001.09

1.27

0.840.43

1.111.150.90

1.211.34

0.780.520.690.920.760.840.761.190.860.960.851.201.101.201.171.061.131.171.201.401.451.291.241.070.860.880.901.070.840.901.011.021.171.131.241.091.181.261.201.311.441.441.441.511.361.751.742.112.112.191.651.571.751.791.741.461.831.51

1.661.201.751.361.41

1.52

1.241.09

0.890.790.94

1.191.04

-0.93-0.80-0.44-0.39

-0.79-0.34-0.71-0.40-0.29-0.80-0.69-0.53-0.83-0.56-0.65-0.57-0.67-0.53-0.55-0.86-0.47-0.82-1.01-0.65

0.21-0.24-0.42-0.76-0.92-0.54-0.51-0.74-0.47•0.51-1.07-0.89-0.70-0.84-0.73-0.69-0.94-1.03-1.09-0.77-1.30-1.17-0.90-1.24-1.00-1.23-1.03-0.94-1.05-1.18-1.04-0.79-0.54

0.85-0.75

-0.83-0.68-0.750.93

-0.51-1.040.78

-0.56-0.55-0.87-0.711.10

-0.45-0.00

1.791.881.881.82

1.931.791.721.571.881.501.511.601.621.431.601.671.791.652.032.021.891.772.081.881.631.741.981.882.181.962.282.332.522.152.002.212.072.302.162.472.462.372.422.092.471.702.693.082.882.212.202.682.352.582.442.392.331.772.42

-0.68 2.35

2.201.722.371.642.152.141.351.721.871.951.941.081.801.74

Hole 590B (Cont.)*52-l, 35-37 480.3 18.94*52-2, 35-37 481.8 19.00*52-3, 35-37 483.3 19.16*52-4, 35-37 484.8 19.32*52-5, 35-37 486.3 19.48*52-6, 35-37 487.8 19.64*53-l, 35-37 489.8 19.86*53-2, 35-37 491.3 20.02*53-3, 35-37 492.8 20.18*53-4, 35-37 494.3 20.34*53-5, 35-37 495.8 20.50*53-6, 35-37 497.4 20.67

1.071.331.071.101.101.091.281.131.301.280.881.20

1.070.980.771.171.321.171.341.401.351.010.981.20

Hole 591a

20-1, 35-3720-1, 128-13020-2. 35-3720-2, 128-13020-3, 35-3720-3, 128-13020-4, 35-3720-4, 128-13020-5, 35-3720-5, 128-13020-6, 35-3720-6, 128-13020,CC21-1,35-3721-1, 128-13021-2, 35-3721-2, 128-13021-3, 35-3721-3, 128-13021-4, 35-3721-4, 128-13021-5,35-3721-5, 128-13021-6, 35-3721-6, 128-13022-1,35-3722-1, 128-13022-2, 128-13022-3, 35-3722-3, 128-13022-4, 35-3722-4, 128-13022-5, 35-3722,CC23-1,35-3723-1, 128-13023-2, 35-3723-2, 128-13023-3,35-3723-3, 128-13023-4, 35-3723-4, 128-13023-5, 35-3723-5, 128-13023-6, 35-3723-6, 128-13023,CC24-1,35-3724-1, 128-13024-2, 35-3724-2, 128-130

174.6175.6176.1177.1177.6178.6179.1180.1180.6181.6182.1183.1183.8184.1185.1185.6186.6187.1188.1188.6189.6190.1191.1191.6192.6193.6194.6196.1196.6197.6198.1199.1199.6202.7203.0204.0204.5205.5206.0207.0207.5208.5209.0210.0210.5211.5212.1212.5213.4214.0214.9

3.373.453.493.573.613.693.733.813.853.933.974.044.104.124.204.244.324.364.444.484.564.604.664.694.754.814.864.954.985.045.075.085.085.115.125.135.135.145.155.165.165.175.185.195.195.205.225.235.255.265.29

2.272.112.222.162.132.162.232.302.182.362.042.012.272.172.262.132.292.352.132.792.252.382.262.132.172.242.352.402.202.112.432.472.312.092.302.342.072.172.282.382.101.352.382.332.262.532.242.402.402.172.27

-0.97-0.87

-0.41-0.08-1.21-0.92-1.16-0.48-0.60-0.52-0.39

1.681.74

1.941.852.152.342.331.951.731.971.84

Core-Section

(interval in cm)

SITE 591Sub-bottom

depth Age

(m) (Ma)

δ 1 8 θ(°/oo PDB)

δ 1 8 θ

δ 1 3 c

(°/oo PDB)

δ 1 3 C

0.640.800.830.680.730.700.710.920.820.580.580.500.720.760.860.810.730.580.680.280.650.320.670.610.720.790.630.630.880.630.350.630.800.480.430.520.320.260.330.150.580.560.490.630.730.370.490.560.690.690.37

Cibicidoides kullenbergi used except in Samples 591-23-1, 35-37 cm,and 591-25-1, 128-130 cm, where C. wuellerstorfi was used.

1409

J. P. KENNETT

Core-Section

(interval in cm)

Hole 591 (Com24-3, 35-3724-3, 128-13024-4, 35-3724-4, 128-13024-5, 25-2724-5, 128-13024-6, 35-3724,CC25-1, 35-3725-1, 128-13025-2, 35-3725-2, 128-13025-3, 35-3725-3, 128-13025-4, 35-3725-4, 128-13025-5, 35-3725-5, 128-13025-6, 35-3725-6, 128-13025,CC26-1, 35-3726-1, 128-13026-2, 35-3726-2, 128-13026-3, 35-3726-3, 128-13026-4, 35-3726-4, 128-13026-5, 35-3726,CC27-1, 35-3727-1, 128-13027-2, 35-3727-2, 128-13027-3, 35-3727-3, 128-13027-4, 35-3727-4, 128-13027-5, 35-3727-5, 128-13027-6, 35-3727-6, 128-13027,CC28-1, 35-3728-2, 35-3728-3, 35-3728-4, 35-3728-5, 35-3728-6, 35-3729-1, 35-3729-2, 35-3729-3, 35-3729-4, 35-3729-5, 35-3729-6, 35-3729,CC30-1, 35-3730-2, 35-3730-3, 35-3730,CC31-1,35-3731-2, 35-3731-3,35-3731-4, 35-3731-5,35-3731-6, 35-3731,CC

Sub-bottomdepth

(m)

t.)215.5216.4217.0217.9218.5219.4220.0221.5221.9222.8223.4224.3224.9225.8226.4227.3227.9228.8229.4230.3231.1231.5232.4233.0233.9234.5235.4236.0236.9237.5240.7241.0242.0242.5243.5244.0245.0245.5246.5247.0248.0248.5249.5250.3250.6252.5253.6255.1256.7258.7260.2261.7263.2264.7266.2267.7269.3269.7271.2272.7274.3274.7276.2277.7279.2280.7282.2283.1

Age(Ma)

5.305.325.335.365.375.405.415.455.465.485.505.525.535.565.575.595.615.635.655.675.695.705.725.745.765.775.805.815.835.855.935.945.965.976.006.016.036.056.076.086.116.126.156.176.176.226.256.336.436.526.636.726.816.906.987.077.177.197.287.377.467.487.577.667.757.837.927.97

δ 1 8 θ(°/oo PDB)

δ 1 8 θ

2.192.232.352.612.212.092.272.302.232.362.212.272.092.112.442.192.392.312.332.452.302.182.262.242.362.042.222.242.222.312.332.022.072.262.462.332.292.262.272.182.332.242.222.202.322.402.332.282.222.262.222.212.102.102.122.222.312.252.152.012.332.092.122.091.952.042.082.31

δ 1 3 C(°/oo PDB)

δ 1 3 C

0.480.750.830.330.710.590.800.620.520.380.570.800.670.340.330.440.580.480.680.690.780.600.450.570.370.470.510.370.600.650.640.590.490.390.590.490.400.350.260.190.390.510.320.530.280.670.490.470.620.490.370.760.590.600.600.800.930.660.800.890.800.830.750.850.950.860.800.85


Hole 591B1-1, 35-371-2, 35-371-3, 35-371,CC2-1, 35-372-2, 35-372-3, 35-372,CC3-1, 35-373-2, 35-373-3, 35-373-5, 35-373,CC4-1, 35-374-2, 35-374-3, 35-374-4, 35-374-5, 35-374,CC5-1, 35-375-2, 35-375-3, 35-375-4, 35-376-1, 35-376-2, 35-376-3, 35-376-4, 35-376-5, 35-376,CC7-1,35-377-1,35-377-2, 35-377-3, 35-377-4, 35-377-5,35-377-6, 35-377,CC8-2, 35-378,CC9-2, 35-379,CC10-1, 35-3710-2, 35-3710-3, 35-3711-1, 35-3711, CC12-1, 35-3712-2, 35-3712,CC13-1,35-3713-2,35-3713,CC14-1, 35-3714-2, 35-3714-6, 35-3714,CC15-1, 35-3715-3, 35-3717-1, 35-3717-2, 35-3717-3, 35-3717-4, 35-3718-1, 35-3718-2, 35-3718-3, 35-3718-4, 35-3718-5, 35-3718-6, 35-3719-1,35-3719-2, 35-3719-4, 35-3719-5, 35-3720-1, 35-3720-2, 35-3720-3, 35-3720-5, 35-3721-1, 35-37

Sub-bottomdepth

(m)

271.0272.5274.0280.2280.5282.0283.5289.8290.2292.0293.2296.2299.4299.8301.3302.8304.3305.8308.8309.2310.7312.2313.7318.5320.0321.5323.0324.5327.6328.0328.0329.5331.0332.5334.0335.5337.2339.0346.8348.7356.4356.8358.3359.8366.3375.6376.0377.3385.2385.5387.0394.8395.2396.7402.7404.4404.8407.8424.0425.5427.0428.5433.5435.0436.5438.0439.5441.0443.2444.7447.7449.2452.8454.3455.8458.8462.3

Age(Ma)

7.277.357.447.807.827.918.008.378.398.508.578.748.938.959.049.139.229.309.489.509.569.629.679.869.919.97

10.0310.0810.2010.2210.2210.2710.3310.3910.4410.5010.5710.6310.9311.0011.4411.4611.5511.6312.0012.1912.2012.2612.6312.6412.6113.0713.0913.1613.4413.5113.5313.6714.4214.4914.5614.6314.8614.9315.0015.1115.2215.3315.4915.6015.8315.9416.2016.2316.2616.3216.39

δ 1 8 θ

(°/oo

2.292.222.232.042.152.201.952.032.171.961.052.102.172.382.212.342.192.342.202.342.332.272.382.032.402.292.352.292.212.242.242.152.022.092.232.292.442.092.302.102.121.741.892.142.081.811.421.982.011.842.121.901.951.851.321.772.121.691.821.611.701.641.581.331.691.541.551.871.621.361.411.401.141.111.361.020.71

δ 1 3 C

PDB)

0.820.950.990.781.020.981.010.841.021.110.710.961.251.190.821.011.031.080.971.111.140.931.001.381.101.141.331.191.361.181.180.951.040.991.020.980.990.820.690.790.720.850.971.101.410.961.411.131.220.940.660.971.180.861.300.871.131.031.251.091.541.521.351.241.521.601.801.911.561.561.401.521.391.401.511.631.74

Species(benthic)

C. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. wuellerstorfiC. wuellerstorfiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. wuellerstorfiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. kullenbergiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC, wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiC. wuellerstorfiO. umbonatusO. umbonatus

1410



Sub-bottomdepth(m)

Hole 591B (Cont.)21-2, 35-3721-3, 35-3722-1,35-3722-3, 35-3722-4, 35-3722-5, 35-3722-6, 35-3723-1, 35-3723-2, 35-3723-3, 35-3723-4, 35-3723-6, 35-3724-2, 35-3724-3, 35-3724-4, 35-37

463.8465.3472.0475.0476.5478.0479.5481.5483.0484.5486.0489.0492.7494.2

. 495.7

Age(Ma)

16.4216.4516.5916.6516.6816.7116.7416.7816.8116.8416.8716.9317.0017.0317.06

(°/oo

0.680.800.870.990.931.041.110.920.800.300.420.620.870.690.54

δ 1 3 C

PDB)

1.291.491.741.912.012.281.971.872.031.541.301.821.281.631.05

O.O.

o.o.o.o.o.o.o.0.

o.o.o.o.o.

Species(benthic)

umbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatusumbonatus

1411

deep sea drilling project initial reports volume 90 · 2007. 4. 25. · 42. miocene to early...

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