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Berger, W.H., Kroenke, L.W., Mayer, L.A., et al., 1993 Proceedings of the Ocean Drilling Program, Scientific Results, Vol. 130 19. PLIOCENE PLEISTOCENE CARBON ISOTOPE RECORD, SITE 586, ONTONG JAVA PLATEAU 1 Jill M. Whitman 2,3 and Wolf H. Berger 2 ABSTRACT Oceanographic changes in the western equatorial Pacific during the past 6 m.y. are inferred from carbon isotopic analyses of planktonic and benthic foraminifers from Ontong Java Plateau (DSDP Site 586). Sample spacing is 1.5 m (ca. 35,000 75,000 yr). An overall trend of δ 13 C toward lighter values is evident for the last 5 m.y. in all four foraminiferal taxa analyzed (G. sacculifer, Pulleniatina, P. wuellerstorfi, and O. umbonatus). This trend is interpreted as an enrichment of the global ocean with 12 C, because of the addition of carbon from organic carbon reservoirs (or lack of removal of carbon to such reservoirs), as a consequence of an overall drop in sea level. Differences between shallow and deep water δ 13 C decrease slightly during this time interval, suggesting a moderate drop in productivity. This drop is not sufficient to explain the drop in sedimentation rate, however, much of which apparently must be ascribed to winnowing effects. A marked convergence in the δ 13 C values of planktonic taxa exists within the last 2 m.y. We propose that this convergence indicates nutrient depletion in thermocline waters, caused by the vigorous removal of phosphate in marginal upwelling regions, or by the stripping of intermediate waters in their source regions. No large shifts are seen in the carbon isotope record of the last 6 m.y., in contrast to the oxygen isotope record. Some indication of cyclicity is present, with a period between 0.5 and 1.0 m.y. (especially in the earlier portion of the record). INTRODUCTION The last 6 m.y. of ocean history are characterized by marked changes in both deep and shallow water circulation, in response to climatic and geochemical changes resulting from mountain building and regression, from the closing of the Middle American Seaway, and from ice buildup in the southern and northern polar regions. The response of the carbon cycle of the ocean to these global changes is of considerable interest. This cycle is an integral part of climatic change, and much of the record in the deep sea is directly tied to the carbon cycle. One place that is ideally suited for studying the record of the ocean's carbon cycle is the Ontong Java Plateau in the western equatorial Pacific. In this region, far from the influence of continental margins and other special conditions, the signals recorded in biogenic sediments have a strong global component. Fluctuations in equatorial upwelling, a phenomenon of global significance, are represented within the sediments on the northernmost portion of the plateau, from which our samples originate. We focus on the stable isotope record of planktonic and benthic foraminifers from this region, using cores recovered at Site 586 of Deep Sea Drilling Project (DSDP) Leg 89 (Shipboard Scientific Party, 1986). Site 586 neighbors Site 289, which was drilled during DSDP Leg 30 (Andrews, Packham, et al., 1975). It is but a short distance from Site 806, drilled during Ocean Drilling Program (ODP) Leg 130 (Kroenke, Berger, Janecek, et al., 1991), so that results can be readily compared among these sites. The results of the oxygen isotope analyses have been published (Whitman and Berger, 1992). Here we present the carbon isotope data and discuss their implications. The taxa we have analyzed are Glo bigerinoides sacculifer, Pulleniatina spp., Planulina wuellerstorfi, and Oridorsalis umbonatus. G. sacculifer is a shallow water species monitoring the mixed layer, Pulleniatina lives in the uppermost 1 Berger, W.H., Kroenke, L.W., Mayer, L.A., et al., 1993. Proc. ODP, Sci. Results, 130: College Station, TX (Ocean Drilling Program). 2 Scripps Institution of Oceanography, University of California, San Diego, La Jolla, Ca 92093 0212, U.S.A. 3 Present address: Department of Earth Sciences, Pacific Lutheran University, Ta coma, WA 98447, U.S.A. thermocline, P. wuellerstorfi and O. umbonatus live on the seafloor and record changes in the chemistry of the water there. SETTING AND STRATIGRAPHY The Ontong Java Plateau is a Texas sized, elevated region east of New Guinea, bearing a layer cake cover of calcareous sediments more than 1 km thick in the shallower portions (Berger and Johnson, 1976). The plateau has long been a favored region for paleoceano graphic investigation, yielding important information on Pleistocene oxygen isotope cycles (Shackleton and Opdyke, 1973, 1976). It has been studied for processes of carbonate sedimentation (Johnson et al., 1977; Berger and Killingley, 1982) and for late Pleistocene pale oceanography (Berger et al., 1987; Hebbeln et al., 1990; Wu et al., 1990; Herguera and Berger, 1991). Also, this area has attracted four different drilling expeditions (DSDP Leg 7, Winterer, Riedel, et al., 1971; DSDP Leg 30, Andrews, Packham, et al., 1975; DSDP Leg 89, Shipboard Scientific Party, 1986; and ODP Leg 130, Kroenke, Berger, Janecek, et al., 1991). Site 586 is close to the equator, at a depth of 2218 m (Fig. 1). It is located on the northeastern upper slope of Ontong Java Plateau (00°29.84'S, 158°29.89'E), about 1 nmi northwest of DSDP Site 289. Our samples derive from the upper 195 m of sediments, recovered by hydraulic piston coring in Holes 586 and 586A. They constitute a continuous and undisturbed record of the last 6 m.y., from the latest Miocene to the present. Sedimentation rates are quite high, ranging from about 40 m/m.y. in the lower portion of the section studied to 20 m/m.y. in the upper part. At the depth (2200 m) and location (equator) of Site 586, the Ontong Java Plateau is bathed in General Pacific Deep Water, about 500 700 m below the Inter mediate Waters of the Pacific (Reid, 1965; Dietrich et al., 1980). The site is located well above the present depth of the foraminifer lysocline in this region (3300 3400 m; Berger et al., 1982; Wu and Berger, 1989). Thus, the isotopic record should be largely unaffected by differential dissolution, and should contain an excellent pale oceanographic record for the late Neogene. The sequence studied here is part of a single lithologic unit (Shipboard Scientific Party, 1986), described as pale green to white foraminifer nannofossil ooze and foraminifer bearing nannofossil ooze (Fig. 2). The sediments of the Pleistocene (0 38 m) contain a larger percentage of foraminifers (up to 60%) compared with those 333

19. PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD, SITE … · 2006. 12. 8. · 19. PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD, SITE 586, ONTONG JAVA PLATEAU1 Jill M. Whitman2,3 and

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  • Berger, W.H., Kroenke, L.W., Mayer, L.A., et al., 1993Proceedings of the Ocean Drilling Program, Scientific Results, Vol. 130

    19. PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD, SITE 586, ONTONG JAVA PLATEAU1

    Jill M. Whitman2 , 3 and Wolf H. Berger2

    ABSTRACT

    Oceanographic changes in the western equatorial Pacific during the past 6 m.y. are inferred from carbon isotopic analyses ofplanktonic and benthic foraminifers from Ontong Java Plateau (DSDP Site 586). Sample spacing is 1.5 m (ca. 35,000-75,000 yr).An overall trend of δ 1 3C toward lighter values is evident for the last 5 m.y. in all four foraminiferal taxa analyzed (G. sacculifer,Pulleniatina, P. wuellerstorfi, and O. umbonatus). This trend is interpreted as an enrichment of the global ocean with 1 2C, becauseof the addition of carbon from organic carbon reservoirs (or lack of removal of carbon to such reservoirs), as a consequence ofan overall drop in sea level. Differences between shallow- and deep-water δ 1 3C decrease slightly during this time interval,suggesting a moderate drop in productivity. This drop is not sufficient to explain the drop in sedimentation rate, however, muchof which apparently must be ascribed to winnowing effects.

    A marked convergence in the δ 1 3C values of planktonic taxa exists within the last 2 m.y. We propose that this convergenceindicates nutrient depletion in thermocline waters, caused by the vigorous removal of phosphate in marginal upwelling regions,or by the stripping of intermediate waters in their source regions. No large shifts are seen in the carbon isotope record of the last6 m.y., in contrast to the oxygen isotope record. Some indication of cyclicity is present, with a period between 0.5 and 1.0 m.y.(especially in the earlier portion of the record).

    INTRODUCTION

    The last 6 m.y. of ocean history are characterized by markedchanges in both deep- and shallow-water circulation, in response toclimatic and geochemical changes resulting from mountain buildingand regression, from the closing of the Middle American Seaway, andfrom ice buildup in the southern and northern polar regions. Theresponse of the carbon cycle of the ocean to these global changes isof considerable interest. This cycle is an integral part of climaticchange, and much of the record in the deep sea is directly tied to thecarbon cycle.

    One place that is ideally suited for studying the record of theocean's carbon cycle is the Ontong Java Plateau in the westernequatorial Pacific. In this region, far from the influence of continentalmargins and other special conditions, the signals recorded in biogenicsediments have a strong global component. Fluctuations in equatorialupwelling, a phenomenon of global significance, are representedwithin the sediments on the northernmost portion of the plateau, fromwhich our samples originate.

    We focus on the stable isotope record of planktonic and benthicforaminifers from this region, using cores recovered at Site 586 ofDeep Sea Drilling Project (DSDP) Leg 89 (Shipboard Scientific Party,1986). Site 586 neighbors Site 289, which was drilled during DSDPLeg 30 (Andrews, Packham, et al., 1975). It is but a short distancefrom Site 806, drilled during Ocean Drilling Program (ODP) Leg 130(Kroenke, Berger, Janecek, et al., 1991), so that results can be readilycompared among these sites.

    The results of the oxygen isotope analyses have been published(Whitman and Berger, 1992). Here we present the carbon isotope dataand discuss their implications. The taxa we have analyzed are Glo-bigerinoides sacculifer, Pulleniatina spp., Planulina wuellerstorfi,and Oridorsalis umbonatus. G. sacculifer is a shallow-water speciesmonitoring the mixed layer, Pulleniatina lives in the uppermost

    1 Berger, W.H., Kroenke, L.W., Mayer, L.A., et al., 1993. Proc. ODP, Sci. Results,130: College Station, TX (Ocean Drilling Program).

    2 Scripps Institution of Oceanography, University of California, San Diego, La Jolla,Ca 92093-0212, U.S.A.

    3 Present address: Department of Earth Sciences, Pacific Lutheran University, Ta-coma, WA 98447, U.S.A.

    thermocline, P. wuellerstorfi and O. umbonatus live on the seafloorand record changes in the chemistry of the water there.

    SETTING AND STRATIGRAPHY

    The Ontong Java Plateau is a Texas-sized, elevated region east ofNew Guinea, bearing a layer-cake cover of calcareous sedimentsmore than 1 km thick in the shallower portions (Berger and Johnson,1976). The plateau has long been a favored region for paleoceano-graphic investigation, yielding important information on Pleistoceneoxygen isotope cycles (Shackleton and Opdyke, 1973, 1976). It hasbeen studied for processes of carbonate sedimentation (Johnson et al.,1977; Berger and Killingley, 1982) and for late Pleistocene pale-oceanography (Berger et al., 1987; Hebbeln et al., 1990; Wu et al.,1990; Herguera and Berger, 1991). Also, this area has attracted fourdifferent drilling expeditions (DSDP Leg 7, Winterer, Riedel, et al.,1971; DSDP Leg 30, Andrews, Packham, et al., 1975; DSDP Leg 89,Shipboard Scientific Party, 1986; and ODP Leg 130, Kroenke, Berger,Janecek, et al., 1991).

    Site 586 is close to the equator, at a depth of 2218 m (Fig. 1). Itis located on the northeastern upper slope of Ontong Java Plateau(00°29.84'S, 158°29.89'E), about 1 nmi northwest of DSDPSite 289. Our samples derive from the upper 195 m of sediments,recovered by hydraulic piston coring in Holes 586 and 586A. Theyconstitute a continuous and undisturbed record of the last 6 m.y.,from the latest Miocene to the present. Sedimentation rates are quitehigh, ranging from about 40 m/m.y. in the lower portion of thesection studied to 20 m/m.y. in the upper part. At the depth (2200 m)and location (equator) of Site 586, the Ontong Java Plateau is bathedin General Pacific Deep Water, about 500-700 m below the Inter-mediate Waters of the Pacific (Reid, 1965; Dietrich et al., 1980). Thesite is located well above the present depth of the foraminiferlysocline in this region (3300-3400 m; Berger et al., 1982; Wu andBerger, 1989). Thus, the isotopic record should be largely unaffectedby differential dissolution, and should contain an excellent pale-oceanographic record for the late Neogene.

    The sequence studied here is part of a single lithologic unit(Shipboard Scientific Party, 1986), described as pale green to whiteforaminifer-nannofossil ooze and foraminifer-bearing nannofossilooze (Fig. 2). The sediments of the Pleistocene (0-38 m) contain alarger percentage of foraminifers (up to 60%) compared with those

    333

  • J.M. WHITMAN, W.H. BERGER

    2°S

  • PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD

    Pidleniatina

    Coiling

    Cor. Ag Uthology % R i « h t

    Table 1. Age control points/events, Site 586.

    " 36.0 m/my

    200

    42. β m/my

    lFor minil f-n noo250 µm was picked, and the total as well as the percentage of certainspecies and genera within the total benthic population were counted:P. wuellerstorfi, O. umbonatus, and all the specimens in the genusUvigerina. These data are listed in Appendix B.

    Stable isotopes were measured on four taxa of foraminifers, asmentioned in the "Introduction" section (this chapter). The planktonicspecimens were picked from the size fraction 355-425 µm and thebenthic species from the fraction >250 µm. The narrow size fractionfor the planktonic foraminifers was used to minimize interferencefrom changing depth ranges and changing vital effects during the lifecycle of the species (Berger et al., 1978). For benthic foraminifers,we did not feel it was necessary to take this precaution, as no evidenceexists that size influences composition (Dunbar and Wefer, 1984). Thesample size for isotopic measurement was 250-1000 µg, dependingupon the availability of specimens; planktonic samples usually con-tained 15-25 individuals, and benthic samples 5-10 individuals. Insome samples, insufficient numbers of one or both of the benthic

    Deptfi

    (mbsf)Age*(Ma) Event From interpolation

    0.0

    *15.4

    36.655.0

    ••64.8 m85.8

    **96.3

    160.2**181.5

    0.00.73

    1.662.472.93.43.8

    5.35.8

    Top of coreBrunhes/Matuyama boundary(Barton and Bloemendahl, 1986)

    LAD G. altispira

    S/D coiling change Pulleniatina

    FAD Pulleniatina

    Pliocene/Pleistocene boundaryMatuyama/Gauss boundary

    Gauss/Gilbert boundary

    Miocene/Pliocene boundary

    •Dates based on Berggren et al. (1985).••Indicates age control points identified in Site 586 sediments. FAD = first appearance datum, LAD

    = last appearance datum, and S/D = sinistral/dextral coiling shift.

    species were present for analysis. In a few instances, duplicatemeasurements were made on one of the planktonic taxa. A VGMicromass 602C mass spectrometer was used to make the measure-ments in the conventional fashion (for details, see Whitman andBerger, 1992). The isotopic data are presented in Appendix C.

    Interpretation of Carbon Isotopes

    We take the δ13C values measured as reflecting the isotopiccomposition of seawater, in the main. This interpretation rests onprevious work on surface sediments of Ontong Java Plateau (Bergeret al., 1978; Vincent et al., 1981). It appears, from these studies, thatmoderately large G. sacculifer and Pulleniatina will yield δ13Cvalues that reflect the composition of the surrounding seawater (not-withstanding the vital effects that undoubtedly play a role in G.sacculifer; see Spero and Williams, 1988). Regarding δ13C in benthicforaminifers, P. wuellerstorfi is thought to be an especially goodrecorder of carbon isotopes in deep water (Woodruff et al., 1980;Graham et al., 1981; Vincent et al., 1981; Woodruff and Savin, 1985;summary in Wefer and Berger, 1991). This species has an epibenthichabitat, which avoids interference from interstitial waters (Grossman,1987; Lutze and Thiel, 1989; McCorkle et al., 1990).

    Changes in the δ13C values of planktonic and benthic foraminiferslargely reflect the shifting of organic matter from one reservoir toanother. Organic carbon is depleted in 1 3C; building up a reservoir,therefore, enriches the ocean in the heavy isotope, and vice versa(Tappan, 1968; Fischer and Arthur, 1977; Shackleton, 1977; Vincentand Berger, 1985; Woodruff et al., 1985; Raymo et al., 1989). Theglacial-to-interglacial contrast for the last cycle, for example, isbetween 0.4‰ and 0.5‰ (Berger and Keir, 1984; Herguera et al.,1992), which corresponds to a transfer of carbon of approximatelyone atmospheric carbon mass (ACM) in and out of organic pools.

    The surface waters are enriched in 13C, as 12C is removed prefer-entially by the sinking of organic matter out of the photic zone. Thecontrast between shallow and deep waters is typically between l‰and 2‰ (Fig. 3A). The difference in δ 1 3 C between planktonic andbenthic foraminifers is a measure of the efficiency of this biologicalpump (Broecker, 1973,1982): a greater difference inδ1 3C points to agreater nutrient concentration, and hence implies greater productivity,other factors being equal (Shackleton et al., 1983a; Boyle, 1988a,1988b). The distribution of oxygen utilization and of nutrients in theoceans (phosphate and nitrate in particular) parallels the trends of12C/13C in the oceans, as oxygen, phosphate, and nitrate all areinvolved in the internal carbon cycle (Kroopnick, 1985). Conversioncoefficients between the different parameters, based on the Redfieldratios or on regressions, are given by Kroopnick (1985) and in Bergerand Spitzy (1988).

    As deep waters move away from their sites of formation (andsource of oxygen), the oxygen concentration continues to decrease

    335

  • J.M. WHITMAN, W.H. BERGER

    (µMOLE/KG)

    2200

    O2(µMOLE/KG)

    200 300

    2 -

    -

    2

    4

    60

    613c

    SAL. l°'ool34

    10 20TEMP l°CI

    I 1

    36 JT

    a

    a

    i

    \

    a\

    \

    é

    / Δ

    /

    1/Δ / NORTH

    i|1 I

    \ 2 4 /

    PACIFIC \ I1

    1

    4 -

    PO4(µMOLE/KG)

    PACIFIC ATLANTIC

    "old "waterO2- poor/Cθ2 - rich

    low 81 3C

    "young" waterC 2-rich /Cθ2-poor

    high S 1 3C

    Figure 3. Internal fractionation of carbon isotopes in the ocean. A. Evidence for preferential removal of 1 2C fromsurface waters by biological pumping (from Kroopnick, 1985). B. Basic pattern of water exchange between thePacific and Atlantic oceans, and resulting δ1 3C distribution in deep waters (from Berger and Vincent, 1986).

    because of the continued combustion of organic matter. Thus, theolder bottom waters of the Pacific have high values of apparentoxygen utilization (AOU; i.e., the difference between observed oxy-gen content and saturation values), and low values of δ1 3C of thedissolved CO2. The overall difference in the δ

    I 3C values of deepPacific and deep Atlantic waters reflects the present asymmetry inbottom-water production, and this asymmetry chan ges through time(Vincent et al., 1980; Shackleton et al., 1983b; Miller and Fairbanks,1985; Keir, 1988; Raymo et al., 1990). The resulting fractionation ofthe isotopes of carbon between the ocean basins is an example ofinterbasin biological fractionation, well-known for phosphate, nitrate,and silica (Sverdrup et al., 1942; Redfield et al., 1963) (Fig. 3B). Thepresent contrast of deep-water values between Atlantic and Pacific istypically near l‰ (Fig. 3B).

    From this brief outline of controls on δ13C patterns in the ocean,it is apparent that a number of different factors must be considered

    simultaneously when interpreting carbon isotope values in foramini-fers. As a rule of thumb, when the benthic and planktonic taxa covary,it reflects a global change in the composition of the oceanic carbonreservoir (from external exchange); however, when the two signalsrecord opposite trends, it is likely a record of changing nutrientcontent in deep waters (driving internal fractionation activity).

    RESULTS

    Overview: Stable Isotopes vs. Depth-in-Hole

    The stable isotope data for Site 586 are plotted against depth inFigure 4. Oxygen isotope data are from Whitman and Berger (1992);carbon isotopes are listed in Appendix C. The oxygen isotopes(Fig. 4A) show clearly different values for the four taxa; averages are-1.31‰, -0.99%o, 2.73%o, and 3.38‰ for G. sacculifer, Pulleniatina,P. wuellerstorfi> a n d O. umbonatus, respectively. A difference of 4‰

    336

  • PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD

    A 5.5

    0 20 40 60 80 100 120 140 160 180 200 220Depth (mbsf)

    B 3 Quaternary] late Pliocene early Pliocene late Miocene

    0 20 40 60 80 100 120 140 160 180 200 220Depth (mbsf)

    Figure 4. Stable isotope stratigraphy for Site 586, plotted within ODP depthframe. A. Oxygen isotopes. B. Carbon isotopes. Lines labeled A, B, and C areevents seen in the oxygen isotope record (see text). SAC = G. sacculifer, PUL= Pulleniatina, WUEL = P. wuellerstorß, and UMBO = O. umbonatus. Theaverage value for each species is marked by a horizontal line labeled with thevalue at right.

    between G. sacculifer and P. wuellerstorfi corresponds to a tempera-ture difference of roughly 20°C. The exact value depends on evapo-ration-precipitation effects and is not readily determined. Clearly, thedifference is increased within the Quaternary (4.3‰; Whitman andBerger, 1992). This is not enough, however, to account for the presenttemperature difference of 27°C, for which about 5.4‰ are expected.The difference of 5.05‰ between G. sacculifer and O. umbonatuscomes close to this expected value (a change of 1°C correspondingto change in δ 1 8 θ of ca. 0.2%o in the present ocean). Thus, O. um-bonatus is closer to equilibrium with seawater than is P. wuellerstorfi>as concerns the oxygen isotope ratio.

    Substantial shifts are seen in the isotopic records; these are labeledas Events "A," "B," and "C." Event A (at 63.5 mbsf) marks a rapidand drastic change in benthic oxygen isotope values, signaling icebuildup or cooling at depth or both. It is the dominant feature in ourrecord. The age of this event is 2.87 Ma in our age model, which agreeswell with the age of a major cooling trend recorded at other sites inthe global ocean (Keigwin, 1986; Loubere and Moss, 1986; Sarntheinand Fenner, 1988; Sarnthein and Tiedemann, 1988; Curry and Miller,1990; Tiedemann, 1991; Whitman and Berger, 1992). The change isless pronounced in the planktonic record than in the benthic one,though it is still present. Thus, a combination of ice buildup and

    deep-water cooling is indicated. (Jansen et al. [1990] show the onsetof ice-rafting at 2.6 Ma in the Norwegian Sea.) Cooling is on the orderof 2°-3°C, judging from the trends seen in these data.

    Event B (at 81 mbsf) marks a change in planktonic δ 1 8 θ values,without a corresponding change in the benthic values. The age of thisevent is 3.36 Ma in our age model. Substantial changes in surface tem-perature in this region of very high temperatures presumably are pre-cluded by strong negative feedback stemming from cloud shading(Ramanathan et al., 1989). Thus, a substantial portion of the changeshould be caused by ice growth. If so, the failure of the benthic foramin-ifers to record such a change is puzzling. Either not much ice growthhas taken place (weakening the argument about strong negative cloudfeedback) or the ice growth was accompanied by a warming of bottomwaters at this depth (2200 m) during this transition period (perhapsfrom an expansion of deep intermediate waters). Event C, denoting adistinct change of planktonic δ 1 8 θ toward lighter values, likewise isnot recorded in the benthic foraminifers. Again, a change in ice mass,combined with a compensating change in deep-water temperature, maybe involved. The event is at 111 mbsf (4.19 Ma in our age model).

    The time between Events C and B (4.2-3.4 Ma) denotes a climaticoptimum (a Pliocene altithermal) in the interval studied. Events B andA are steps leading into a colder climate with increased amounts of ice.After Event A the scatter in the data increases greatly, reflecting theincreasing amplitude of glacial-interglacial sea-level variation thatresults from a buildup of climate-sensitive ice caps (most likely in theNorthern Hemisphere, as seen in glacial-derived sediments; Berggren,1972; Backman, 1979; Shackleton et al., 1984).

    The carbon isotope values of the four taxa (from top to bottomin Fig. 4B: G. sacculifer, Pulleniatina, P. wuellerstorfi, and O. um-bonatus) are distinct, with G. sacculifer having the highest values,as expected for this shallow-dwelling species. P. wuellerstorfi showsvalues close to equilibrium, whereas O. umbonatus is notably de-pleted in 1 3C. None of the taxa show striking events or strong trends.Both G. sacculifer and P. wuellerstorfi show a tendency from morepositive to more negative values over the interval studied, presum-ably indicating the input of organic carbon to the dissolved inorganiccarbon (DIC) pool of the ocean. The mid-Pliocene altithermal hasrelatively low δ1 3C values in G. sacculifer, suggesting lowerednutrient concentrations in the deep waters. At the same time, theO. umbonatus values are relatively high, supporting a hypothesis oflowered productivity. (The assumption is that increased productivity,by supplying organic matter to the seafloor, would tend to lower theδ1 3C values in this species.) Some indication of cyclicity is presentin the record of Pulleniatina.

    Carbon Isotope Patterns: Differences Between Taxa

    The age-plot of stable isotopes (Fig. 5) clearly shows the contrastin sampling density between the upper and lower portions of thesection that is a result of the changing sedimentation rate (Fig. 2).Thus, documentation is better for pre-Event A time than for post-Event A time. In addition, the scatter is distinctly less in pre-Event Atime, which means that each point here is more representative for abroader interval. The oxygen isotopes (Fig. 5A) are given for orien-tation; they are discussed in Whitman and Berger (1992).

    The difference between the values of G. sacculifer and P. wuel-lerstorfi in the youngest portion of the section agrees well with thatseen in the modern oceanographic data, that is, the present composi-tion of dissolved inorganic carbon (Fig. 3). At nearby GEOSECSStation 246 (0°00'S, 178°59'E), the value for δ13C of I C O 2 at thesurface is near+1.5‰ and that at 2200 m is between -0. l‰ to -0.2‰(Kroopnick, 1985), for a difference of 1.6‰ to 1.7‰. Allowing forthe anthropogenic effect of lowered δ13C in present surface waters,the expected difference is near 2‰. The relevant values in theforaminifers (Appendix C) closely reflect this difference. The agree-ment supports our assumption of near-equilibrium precipitation for

    337

  • J.M. WHITMAN, W.H. BERGER

    Table 2. Carbon difference data. A 5.5

    For entire core:G. sacculifer-C. wuellerstorfiC. wuellerstorfi-O. umbonatusG. sacculifer-Pulleniatina

    For subdivided intervals:C. wuellerstorfi-O• umbonatus

    4.4Ma)

    G. sacculifer-Pulleniatina1.8Ma)38-85 m (1.8-3.4 Ma)85-122 m (3.4-^.4 Ma)122-160 m (4.4-5.3 Ma)>160 m (>5.3 Ma)

    G. sacculifer-C. wuellerstorfi

  • PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD

    2.6π

    2.4

    2.2

    2-

    1.8-

    1.6

    1.4-

    1.2-

    1-

    0.8-n K-

    SAC vs WUEL

    • α π _π

    m

    m y = x + 1.9

    π

    o

    +

    Legend• post A+ AtoCπ preC

    -0.5-0.4-0.3-0.2-0.1 0 0.1 0.2 0.3 0.4 0.5δ13C (‰, P. wuellerstorfi)

    Quaternary | late Pliocene | early Pliocene I late Miocene

    2 3 4Age (Ma)

    B -0.4

    -1.8

    UMBOvsWUEL

    y = x - 1 . 1 7

    -0.5-0.4-0.3-0.2-0.1 0 0.1 0.2 0.3 0.4 0.5δ13C (‰, P. wuellerstorfi)

    2.8

    2.6H

    J 2.4

    I 2.2

    I 2̂é 1 8o% 1-6-

    1.41.2

    SAC vs PUL

    c P •

    V ^

    y = x + 1.07

    10.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 1.1 1.21.31.41.5

    δ1 3C (‰, Pulleniatina)

    Figure 6. Scatter plots of δ1 3C values of one taxon vs. another. A. G. sαcculifer

    vs. P. wuellerstorfi• B. O. umbonαtus vs. P. wuellerstorfi• C. G. sαcculifer vs.

    Pulleniatina. The line denotes constant offset, given by the equation in each

    panel. Symbols refer to different time spans as defined by Events A, B, and C

    (as shown in legend).

    1.8

    1.6

    1.4

    1.2

    0.8:

    0.6

    0.4-

    0.2

    0

    Quaternary late Pliocene early Pliocene late Miocene

    -ç -tt-

    π π

    G. sacculifer-Pulleniatina

    carbon

    1.06

    0.31oxygen

    0 2 3 4Age (Ma)

    Figure 7. Differences in stable isotope values between G. sαcculifer and two

    other taxa. A. G. sαcculifer and P. wuellerstorfi B G• sαcculifer and

    Pulleniatina. Open squares denote δ 1 8 θ and crosses denote δ 1 3C values.

    Average differences shown at right and marked by horizontal lines.

    planktonic species (Fig. 6C) shows that correlations change signifi-cantly through time: whereas correspondence is good in pre-Event Ctime, it is very poor in post-Event A time. Again, G. sacculifer showsa large range, the comparison species a smaller one.

    Carbon Isotope Patterns: Difference Trends

    Differences in stable isotope records are further explored in Fig-ure 7, where we compare δ 1 8 θ and δ13C contrasts between taxa in thetime domain. Post-Event A time is characterized by increasing andgreater-than-average differences in δ 1 8 θ values between G. saccu-lifer and P. wuellerstorfi, and an overall decrease in δ13C difference(Fig. 7 A). Atendency for negative correlation between the differencespersists throughout the section, indicating that whenever the thermalgradient increases (increased difference in δ 1 8 θ), there is an increasedlikelihood for a lowered nutrient content in deep waters (decreasedcontrast in δ13C). This same relationship is reflected in the differencestratigraphy of G. sacculifer and Pulleniatina (Fig. 7B): increasedtemperature contrast (i.e., shallowing of the thermocline) parallelsdecreased δ13C contrast (i.e., nutrient content of the thermocline). Thetrend in the Quaternary is especially remarkable in this respect; δ 1 8 θdifferences increase greatly over the general average of 0.31‰,whereas δ13C differences drop way below their average of 1.06%c.

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  • J.M. WHITMAN, W.H. BERGER

    On the whole, the patterns suggest that as the deep ocean becomescolder and the thermocline rises, vertical fractionation becomes lesseffective, that is, nutrient concentrations decrease.

    Carbon Isotope Patterns: Variability Through Time

    The variability of stable isotopes changes through time andreaches a maximum in post-Event A time in almost all categories(excepting δ13C of P. wuellerstorfi) (Fig. 8). Variability is taken as thestandard deviation about the mean of five consecutive values downthe hole, that is, an interval of 8 m (ca. 200-400 k.y., depending onthe sedimentation rate). Among the δ 1 8 θ records (Fig. 8A), the oneof G. sacculifer is the least variable, supporting the notion that surfacetemperatures tend to stay constant in this region because of the strongnegative cloud feedback. (Seasonal changes are on the order of 1°C.)Also, a general cooling (which would be reflected in ice buildup anda lowering of deep-water temperatures) could conceivably increasethe delivery of warm surface waters to the western equatorial Pacificby strengthened trade winds. Thus, the effects of ice buildup wouldbe compensated in part. Variability in the δ 1 8 θ record of Pulleniatinais somewhat greater, especially in the Quaternary, and the same is truefor the benthic species. The range of these fluctuations is readilyaccommodated by an ice effect on the order of l‰. The lesservariability in pre-Event A time, presumably, indicates that the effectfrom the waxing and waning of continental ice on δ 1 8 θ values of theocean was less important before 3 Ma than afterward. Variations indepth-to-thermocline, and in proportion of North Atlantic Deep Water(NADW) component in the waters bathing the site of deposition, alsohave to be considered when discussing these records.

    The corresponding δ13C records show much less change in vari-ability through time, compared with the δ 1 8 θ records (Fig. 8B). In thecarbon isotope record G. sacculifer is the most variable, and P. wuel-lerstorfi the least, as mentioned, reflecting the inertia of deep-watercarbon isotopic composition. There is some indication that variabilityis at a minimum in the extended Pliocene altithermal (Events A to C)in all taxa. If so, variability in the carbon isotope record is a functionof the extent of cooling, that is, ice mass extant. However, no strongincrease is present in overall variability (or in the variation of vari-ability) toward the late Quaternary, as seen in the oxygen isotopes.

    Productivity-related Indices

    The δ1 3C records of the different taxa are closely tied to produc-tivity changes in the ocean. To avoid circular reasoning, it is impor-tant to attempt to reconstruct such changes in ways other thanthrough the carbon isotope record itself. One well-established proxyfor productivity is the abundance of benthic foraminifers in thesediment, which records the supply of organic matter to the seafloor(Altenbach and Sarnthein, 1989; Herguera and Berger, 1991). Over-all trends in this index show a general increase of both abundanceand variability in abundance through the period studied (Fig. 9A).Numbers of benthic foraminifers >250 µm per gram typically showvalues between 10 and 15 in the early Pliocene, increasing tobetween 15 and 30 in the Quaternary. This trend actually representsa decrease (or at best no change) in the accumulation rate of benthicforaminifers, as the sedimentation rate differs by a factor of 2 forthese time periods. Thus, a slight reduction in overall productivityseems indicated. There is some hint of cyclicity on a 1-m.y. scale(Fig. 9A, running average).

    Somewhat surprisingly, the relative abundance of Uvigerina, onthe whole, increases over the period studied, except for some unusu-ally high values in the latest Miocene (Fig. 9B). A high abundanceof Uvigerina has been associated with low oxygen content of bottomwaters and with a high supply of organic matter (Streeter andShackleton, 1979; Douglas and Woodruff, 1981; Corliss, 1982;Corliss et al., 1986; Lutze et al, 1986; Zahn et al., 1986). In the caseat hand, lowered oxygen content would seem to be the preferred

    A 1.8

    2 3 4Age (Ma)

    late Pliocene early Pliocene late Miocene

    Figure 8. Variability of stable isotopes as a function of age. Variability is givenas the standard deviation from the mean of 5 consecutive values downhole (ca.8-m interval). A. Oxygen isotopes. B. Carbon isotopes. SAC = G. sacculifer,PUL = Pulleniatina, WUEL = P. wuellerstorfi, and UMBO = O. umbonatus.In each panel, the horizontal line is zero (offset in steps of 0.4‰ added to rawstandard deviation [SD] values, for separation).

    explanation, considering the lack of evidence for increased produc-tivity. However, temporarily increased output from pulsed productiv-ity, even if occurring on the background of reduced production, mayhave to be considered in the context as an important factor in theUvigerina pattern.

    A general decrease in oxygen in the latest Neogene at this site—orpulsed reduction of oxygen—is not supported by the carbon isotopedata, which do not show increased differences between G. sacculiferand P. wuellerstorfi parallel to the increase in relative abundance ofUvigerina (cf. Figs. 7A and 9B). However, a decrease of oxygen inPacific deep waters is expected from the turning up of deep-waterproduction in the North Atlantic in the late Pliocene, a point taken upin the discussion section.

    It seems unlikely that upwelling, on the whole, should havedecreased in the late Pliocene and since. The increase in the depthgradient in shallow-water temperature (as seen in the oxygen isotopedata of the two planktonic species; Fig. 7B) suggests that the thermo-cline rose (rather than being depressed by warm-water pileup). Also,general considerations involving the strength of trade winds (Ar-rhenius, 1952; Leinen and Heath, 1981) would seem to favor in-creased rather than decreased upwelling. A trend of decreasingproductivity combined with a trend of increased upwelling wouldimply the depletion of nutrients from thermocline waters (Berger and

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  • PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD

    A 50

    2 3 4Age (Ma)

    3 4 5Age (Ma)

    Figure 9. Productivity-related sediment properties, Site 586. A. Benthic fora-minifers per gram of sediment in the fraction >250 µm. B. Log of percentabundance of Uvigerina spp. C. Sand content. The solid lines through the datapoints are 5-point sliding averages. Events labeled A, B, and C as in Figure 5.

    Wefer, 1991; Berger and Herguera, 1992). It remains to be demon-strated, however, that physical upwelling, in fact, proceeded whilenutrients declined. The problem is that the reconstruction of up-welling almost invariably involves productivity-related proxies.

    Sand content reflects, in essence, the ratio between foraminifersand coccoliths, which may be taken as an index of productivity whenwinnowing and dissolution are unimportant (Berger, 1976). The sandfraction shows an overall increase after 3 Ma in the section studied(Fig. 9C). The scatter also greatly increases after 3 Ma. However, theaccumulation of sand does not increase, as the change in sedimenta-tion rate more than compensates for the change in sand content.Overall, then, no increase occurs in the supply of foraminifers. It islikely that the variations in sand content are largely produced bywinnowing, as has been proposed for the late Quaternary in this region(Wu and Berger, 1991). A more detailed age model than the one hereused will be necessary to show this for fluctuations on the scale of1 m.y. and less. An overall increase in winnowing could be the result,

    for example, of increased tidal action at depth, as shelves are no longeravailable for tidal energy destruction after large-scale regression.Other possibilities also must be considered (e.g., trends in tsunamiactivity and frequency of benthic storms).

    DISCUSSION AND SUMMARY

    Major Trends

    The main feature of our oxygen isotope record is Event A, thecentral cooling step that occurs between 2.8 and 2.9 m.y. ago in ourage model. Event A is close enough to the timing of the establishmentof the Panama Isthmus (Keigwin, 1978, 1982) to suggest some sortof causal connection by means of changes in the energy budget(Maier-Reimer et al., 1990). What is surprising is that the carbonisotope record does not clearly reflect this change, although muchevidence exists that a profound change takes place in the way theclimatic system works at that time (expressed as a change in cyclicitiesand covariance patterns of δ 1 8 θ ; see Tiedemann, 1991; Prell, 1984).

    Compared with the oxygen isotope record, the carbon isotoperecord is both more complicated and less eventful (Fig. 10). No majorsteps occur that would point to drastic changes in the carbon compo-sition of the global ocean, either permanently or temporarily. Instead,we see quasi-cyclic variations, presumably reflecting the up and downof sea level, and the accompanying exchange of the ocean's carbonreservoir with shallow marine and terrestrial reservoirs. However,distinct trends are present that reflect the overall climatic change froma world with only a southern ice cap to one with permanent ice massesat both poles and with greatly fluctuating volumes in the north. Thesetrends include an overall decrease in oceanic δ13C, an apparentdecrease in productivity, and an apparent change in deep-water prop-erties, toward greater asymmetry between the Atlantic and Pacificoceans. In the following material, we discuss these trends and offeradditional evidence bearing on their existence and magnitude.

    The trend of a gradual decrease in δ1 3C values since the earlyPliocene, in both benthic and planktonic records, indicates a changein the composition of the global carbon reservoir. We suggest thatglobal regression is responsible, increasing the delivery of lightcarbon from shallow marine or terrestrial organic matter to the oceans.The general trend of a falling sea level over this time period supportsthis hypothesis. The curve presented by Haq et al. (1987) shows adrastic and permanent fall of average sea level shortly after 3 Ma.Increased erosion of soil carbon is compatible with the change in theratio of 87Sr/86Sr in seawater, as preserved in carbonate sediments(Koepnick et al., 1988). Since the early Pliocene, the heavier isotopehas increased, in response to increased delivery from terrestrialsources because the exposure of a greater area of continental rocks toweathering results in an increased flux of radiogenic 87Sr to the oceans(Palmer and Elderfield, 1985; Capo and DePaolo, 1990).

    One could argue, from the decreasing difference in 613C values ofG. sacculifer and P. wuellerstorß within the last 3 m.y. (Fig. 7A) thatan overall decrease in productivity in this time interval occurred atthis site. A lowering of accumulation rates is in accord with thisobservation (Fig. 2). However, a substantial portion of the reductionin the sedimentation rate through the period considered presumablyis caused by increased winnowing rather than by a drop in productiv-ity. The sand fraction changes from a background value near 20% toabout 35% in the latest Quaternary (Fig. 9C). If this difference wereentirely a result of winnowing, the loss of sediment (made up of siltand clay) is given by

    L=l-SJS, (1)

    where So is the initial sand content, and S the final one after winnowing.Thus, about 43% of the sediment is missing under this assumption, fora reduction of the sedimentation rate to 57% of the original value.

    The actual reduction of the sedimentation rate is close to a factor of2, that is, 50% of the original material is lost by winnowing, under the

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  • J.M. WHITMAN, W.H. BERGER

    3 4Age (Ma)

    1 2 3 4 5 6 7Age (Ma)

    Figure 10. Long-term trends in stable isotope compositions of four foraminif-eral taxa: G. sacculifer (SAC), Pulleniatina (PUL), P. wuellerstorfi (WUEL),and O. umbonatus (UMBO). A. Oxygen isotopes. B. Carbon isotopes. Datasmoothed using 5-point sliding averages.

    assumption that winnowing produces the changes in rate. If we use thesand fraction >147 µm in the data of Gardner et al. (1986) to make thesame loss calculation again, we do obtain an apparent reductionapproaching 50%. Thus, a substantial portion of the decrease in accu-mulation rate in the late Neogene, at this site, could easily be a resultof the downslope removal of fine material, including a portion of thefine sand. If this process of downslope removal was widespread duringthe Pleistocene, it would have constantly delivered carbonate to thedeepest waters, increasing the alkalinity there, and lowering the car-bonate compensation depth (CCD). The effect would be a reduction inpCO2 in the atmosphere. A drop of the CCD on the order of 500 m(Peterson and Backman, 1990) would have produced a reduction inpCO2 of about 30 ppm, other factors being equal (Berger and Spitzy,1988). The grain-size patterns, then, caution us not to interpret thechanges in accumulation rate as being mainly caused by changes inproductivity, and the hypothesis of alkalinity increase from winnowinggives us a mechanism to decrease atmospheric CO2 independently ofthe efficacy of the biological pump.

    Thermocline Depletion

    Thermocline depletion is indicated especially for the Pleistoceneby the convergence in the carbon values of the two planktonic species(Fig. 10B). The trend implies that the increased supply of upwellingcold water (as indicated by divergence in δ 1 8 θ values; Fig. 10A) does

    not bring an increased supply of nutrients to fuel the productivity ofsurface waters. Conceivably, increased productivity in the sourceareas of the thermocline water (subantarctic, subarctic, or both)removes nutrients there, resulting in decreased supply to intermediatewaters, and hence to tropical upwelling regions. This mechanism maybe important on a glacial-interglacial scale as well (Keir, 1988).Alternatively, or in addition, nutrients are extracted in the margins asthe planet cools, leaving less for the open ocean. Such an explanationwould be in accord with decreased silica accumulation during glacialperiods, in this region (Lange and Berger, this volume).

    One possible explanation for nutrient depletion of thermoclinewaters in their areas of origin is an increase in the supply of iron duringthe last 6 m.y. (Berger and Wefer, 1991). Iron may be limiting toproductivity in high latitudes, and increased glacial productivity therein consequence of a greater dust supply has been postulated (Martin,1990). Boyle and Keigwin (1987) assumed decreased intermediate-water nutrient content in the North Atlantic during the last glacial andexplained the effect with reduced NADW production and increasedNorth Atlantic intermediate water production. A more general model,postulating transfer of nutrients to deep waters, is given by Boyle(1988a, 1988b). The glacial depletion of nutrients apparently oc-curred in North Pacific intermediate waters down to a depth of2000 m, judging from radiocarbon evidence and stable isotope data(Berger, 1987; Duplessy et al., 1988; Herguera et al., 1991, 1992).Similar data are reported from the Indian Ocean (Kallel et al., 1988).Thus, thermocline depletion during glacial phases is a global phe-nomenon. We assume that the Pleistocene as a whole representsglacial conditions, compared with the Pliocene.

    Comparison with Atlantic Record

    The crude analogy setting the Pleistocene equal to glacial condi-tions and the Pliocene equal to interglacial ones is useful, but it has tobe applied with discretion. For example, it is now generally acceptedthat the production of NADW was reduced during glacials, comparedwith interglacials, so that the asymmetry in δ13C values between theAtlantic and Pacific oceans was reduced during cold periods (e.g.,Shackleton et al., 1983b). Yet, when comparing the deep-ocean recordfrom Site 586 with that from Site 552 in the North Atlantic (Fig. 11),we note that the asymmetry in the δ13C is strong throughout the period.If anything, it has increased slightly during the last 3 m.y.

    In this case, then, the apparent change in the effects of NADW pro-duction is not exactly analogous to the NADW fluctuations within thePleistocene, when NADW is turned down during the glacials (Duplessyet al., 1980; Boyle and Keigwin, 1982; Curry et al., 1988). This situ-ation indicates that contrasts within the Pleistocene cannot be readilyextrapolated to warm/cold contrasts comprising larger time scales.

    ACKNOWLEDGMENTS

    We are indebted to Edith Vincent, Robin Keir, and Frances Parkerfor advice during the initial stages of the work here reported on.Christina Ravelo read an earlier draft of the manuscript and made manyuseful suggestions for improvement. Supported by the U.S. NationalScience Foundation (Grants OCE88-16167 and OCE90-17717).

    REFERENCES*

    Altenbach, A.V., and Sarnthein, M., 1989. Productivity record in benthicforaminifera. In Berger, W.H., Smetacek, V.S., and Wefer, G. (Eds.),Productivity of the Ocean: Present and Past: Chichester (Wiley), 255-269.

    * Abbreviations for names of organizations and publication titles in ODPreference lists follow the style given in Chemical Abstracts Service SourceIndex (published by American Chemical Society).

    342

  • PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD

    Andrews, J.E., Packham, G., et al., 1975. Init. Repts. DSDP, 30: Washington(U.S. Govt. Printing Office).

    Arrhenius, G., 1952. Sediment cores from the East Pacific. Swed. Deep-SeaExped., 1947-1948, 5 (Pts. 1-3).

    Backman, J., 1979. Pliocene biostratigraphy of DSDP Sites 111 and 116 fromthe North Atlantic Ocean and the age of Northern hemisphere glaciation.J. Stockh. Contrib. Geol, 32:115-137.

    Barton, C.E., and Bloemendal, J., 1986. Paleomagnetism of sediments col-lected during Leg 90, Southwest Pacific. In Kennett, J.P., von der Borch,C.C., et al., Init. Repts. DSDP, 90, Pt. 2: Washington (U.S. Govt. PrintingOffice), 1273-1316.

    Belanger, P.E., Curry, W.B., and Matthews, R.K., 1981. Core-top evaluationof benthic foraminiferal isotopic ratios for paleo-oceanographic interpre-tations. Palaeogeogr., Palaeoclimatol., Palaeoecol, 33:205-220.

    Berger, W.H., 1976. Biogenous deep-sea sediments: production, preservationand interpretation. In Riley, J.P., and Chester, R. (Eds.), Treatise onChemical Oceanography (Vol. 5): London (Academic Press), 265-388.

    , 1987. Ocean ventilation during the last 12,000 years: hypothesis ofcounterpoint deep water production. Mar. Geol, 78:1-10.

    Berger, W.H., Bonneau, M.-C, and Parker, F.L., 1982. Foraminifera on thedeep-sea floor: lysocline and dissolution rate. Oceanol. Acta, 5:249-258.

    Berger, W.H., and Herguera, J.C., 1992. Reading the sedimentary record of theocean's productivity. In Falkowski, P.G., and Woodhead, A.D. (Eds.),Primary Productivity and Biogeochemical Cycles in the Sea: New York(Plenum), 455-486.

    Berger, W.H., and Johnson, T.C., 1976. Deep-sea carbonates: dissolution andmass wasting on Ontong-Java Plateau. Science, 192:785-787.

    Berger, W.H., and Keir, R.S., 1984. Glacial-Holocene changes in atmosphericCO2 and the deep-sea record. In Hansen, J.E., and Takahashi, T. (Eds.),Climate Processes and Climate Sensitivity (Vol. 5). Am. Geophys. Union,Maurice Ewing Sen, 29:337-351.

    Berger, W.H., and Killingley, J.S., 1982. Box cores from the Equatorial Pacific:C sedimentation rates and benthic mixing. Mar. Geol, 45:93-125.

    Berger, W.H., Killingley, J.S., and Vincent, E., 1978. Stable isotopes indeep-sea carbonates: Box Core ERDC-92, west-equatorial Pacific.Oceanol. Acta, 1:203-216.

    Berger, W.H., Killingley, J.S., and Vincent, E., 1987. Time scale of Wiscon-sin/Holocene transition: oxygen isotope record in the western equatorialPacific. Quat. Res., 28:295-306.

    Berger, W.H., and Spitzy, A., 1988. History of atmospheric CO2: constraintsfrom the deep-sea record. Paleoceanography, 3:401-411.

    Berger, W.H., and Vincent, E., 1986. Deep-sea carbonates: reading the carbonisotope signal. Geol. Rundsch., 75:249-269.

    Berger, W.H., and Wefer, G., 1991. Productivity of the glacial ocean: discus-sion of the iron hypothesis. Limnol. Oceanogr., 36:1899-1918.

    Berggren, W.A., 1972. Late Pliocene-Pleistocene glaciation. In Laughton,A.S., Berggren, W.A., et al., Init. Repts. DSDP, 12: Washington (U.S. Govt.Printing Office), 953-963.

    Berggren, W.A., Kent, D.V., and Van Couvering, J.A., 1985. The Neogene:Part 2. Neogene geochronology and chronostratigraphy. In Snelling, N.J.(Ed.), The Chronology of the Geological Record. Geol. Soc. LondonMem., 10:211-260.

    Boyle, E.A., 1988a. Vertical oceanic nutrient fractionation and glacial/inter-glacial CO2 cycles. Nature, 331:55-56.

    , 1988b. The role of vertical chemical fractionation in controllinglate Quaternary atmospheric carbon dioxide. /. Geophys. Res., 93:15701-15714.

    Boyle, E.A., and Keigwin, L.D., 1982. Deep circulation of the North Atlanticover the last 200,000 years: geochemical evidence. Science, 218:784-787.

    , 1987. North Atlantic thermohaline circulation during the past 20,000years linked to high-latitude surface temperature. Nature, 330:35-40.

    Broecker, W.S., 1973. Factors controlling CO2 content in the oceans andatmosphere. In Woodwell, G.M., and Pecan, E.V. (Eds.), Carbon and theBiosphere. AEC Symp., 30:32-50.

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

    Capo, R.C., and DePaolo, D.J., 1990. Seawater strontium isotopic variationsfrom 2.5 million years ago to the present. Science, 249:51-55.

    Corliss, B.H., 1982. Linkage of North Atlantic and Southern Ocean deep-watercirculation during glacial intervals. Nature, 298:458-^60.

    , 1985. Microhabitats of benthic foraminifera within deep-sea sedi-ments. Nature, 314:435^38.

    Corliss, B.H., Martinson, D.G., and Keefer, T., 1986. Late Quaternary deep-ocean circulation. Geol. Soc. Am. Bull, 97:1106-1121.

    Curry, W.B., Duplessy, J.-C, Labeyrie, L.D., and Shackleton, N.J., 1988.Changes in the distribution of δ C of deep water XCO2 between the lastglaciation and the Holocene. Paleoceanography, 3:317-341.

    Curry, W.B., and Miller, K.G., 1990. Oxygen and carbon isotopic variation inPliocene benthic foraminifers of the equatorial Atlantic. In Ruddiman, W,Sarnthein, M., et al., Proc. ODP, Sci. Results, 108: College Station, TX(Ocean Drilling Program), 157-166.

    Dietrich, G., Kalle, K., Krauss, W., and Siedler, G., 1980. General Oceanog-raphy: An Introduction (2nd ed.): New York (Wiley-Interscience).

    Douglas, R.G., and Woodruff, F., 1981. Deep sea benthic foraminifera. InEmiliani, C. (Ed.), The Sea (Vol. 7): The Oceanic Lithosphere: New York(Wiley-Interscience).

    Dunbar, R.B., and Wefer, G., 1984. Stable isotopic fractionation in benthicforaminifera from the Peruvian continental margin. Mar. Geol, 59:215-225.

    Duplessy, J.-C, Moyes, J., and Pujol, C, 1980. Deep water formation in theNorth Atlantic Ocean during the last ice age. Nature, 286:479-482.

    Duplessy, J.C., Shackleton, N.J., Fairbanks, R.G., Labeyrie, L., Oppo, D.,Kallel, N., 1988. Deep-water source variations during the last climaticcycle and their impact on the global deep-water circulation. Paleoceano-graphy, 3:343-360.

    Fischer, A.G., and Arthur, M.A., 1977. Secular variations in the pelagic realm.Spec. Publ.—Soc. Econ. Paleontol Mineral, 25:19-50.

    Gardner, J.V., Dean, WE., Bisagno, L., and Hemphill, E., 1986. Late Neogeneand Quaternary coarse-fraction and carbonate stratigraphies for Site 586 onOntong-Java Plateau and Site 591 on Lord Howe Rise. In Kennett, J.P., vonder Borch, C.C., et al., Init. Repts. DSDP, 90: Washington (U.S. Govt.Printing Office), 1201-1224.

    Graham, D.W., Corliss, B.H., Bender, M.L. and Keigwin, L.D., 1981. Carbonand oxygen isotopic disequilibria of recent deep-sea benthic foraminifera.Mar. Micropaleontol, 6:483^497.

    Grossman, E.L., 1987. Stable isotopes in modern benthic foraminifera: a studyof vital effect. J. Foraminiferal Res., 17:48-61.

    Haq, B.U., Hardenbol, J., and Vail, P.R., 1987. Chronology of fluctuating sealevels since the Triassic. Science, 235:1156-1167.

    Hebbeln, D., Wefer, G., and Berger, W.H., 1990. Pleistocene dissolutionfluctuations from apparent depth of deposition in Core ERDC-127P, west-equatorial Pacific. Mar. Geol, 92:165-176.

    Herguera, J.C., and Berger, W.H., 1991. Paleoproductivity from benthic fo-raminifera abundance: glacial to postglacial change in the west-equatorialPacific. Geology, 19:1173-1176.

    Herguera, J.C., Jansen, E., and Berger, W.H., 1992. Evidence for a bathyalfront at 2000-m depth in the glacial Pacific, based on a depth transect onOntong Java Plateau. Paleoceanography, 7:273-288.

    Herguera, J.C., Stott, L.D., and Berger, W.H., 1991. Glacial deep-waterproperties in the west-equatorial Pacific: bathyal thermocline near a depthof 2000 m. Mar. Geol, 100:201-206.

    Jansen, E., Sj0holm, J., Bleil, U., and Erichsen, J.A., 1990. Neogene andPleistocene glaciations in the Northern Hemisphere and late Miocene-Plio-cene global ice volume fluctuations: evidence from the Norwegian Sea. InBleil, U., and Thiede, J. (Eds.), Geological History of the Polar Oceans:Arctic Versus Antarctic: Dordrecht (Kluwer Academic), 677-705.

    Johnson, T.C., Hamilton, E.L., and Berger, W.H., 1977. Physical properties ofcalcareous ooze: control by dissolution at depth. Mar. Geol, 24:259-277.

    Kallel, N., Labeyrie, L.D., Juillet-Leclerc, A., and Duplessy, J.-C, 1988. Adeep hydrological front between intermediate and deep-water masses inthe glacial Indian Ocean. Nature, 333:651-655.

    Keigwin, L.D., 1978. Pliocene closing of the Isthmus of Panama, based onbiostratigraphic evidence from nearby Pacific Ocean and Caribbean Seacores. Geology, 6:630-634.

    , 1982. Isotopic paleoceanography of the Caribbean and east Pacific:role of Panama uplift in the late Neogene time. Science, 217:350-353.

    -, 1984. Stable isotopic results on upper Miocene and lower Plioceneforaminifers from Hole 552A. In Roberts, D.G., Schnitker, D., et al., Init.Repts. DSDP, 81: Washington (U.S. Govt. Printing Office), 595-597.

    -, 1986. Pliocene stable-isotope record of Deep Sea Drilling ProjectSite 606: sequential events of 1 8O enrichment beginning at 3.1 Ma. InRuddiman, W.F., Kidd, R.B., Thomas, E., et al., Init. Repts. DSDP, 94, Pt.2: Washington (U.S. Govt. Printing Office), 911-920.

    Keigwin, L.D., Aubry, M.-P., and Kent, D.V., 1987. North Atlantic lateMiocene stable-isotope stratigraphy, biostratigraphy and magnetostratig-raphy. In Ruddiman, WE, Kidd, R.B., Thomas, E., et al., Init. Repts. DSDP,94, Pt. 2: Washington (U.S. Govt. Printing Office), 935-963.

    Keir, R.S., 1988. On the late Pleistocene ocean geochemistry and circulation.Paleoceanography, 3:413-445.

    343

  • J.M. WHITMAN, W.H. BERGER

    Koepnick, R.B., Denison, R.E., and Dahl, D.A., 1988. The Cenozoic 87Sr/86Srcurve: data review and implications for correlation of marine strata.Paleoceanography, 3:743-756.

    Kroenke, L.W., Berger, W.H., Janecek, T.R., et al., 1991. Proc. ODP, Init.Repts., 130: College Station, Texas (Ocean Drilling Program).

    Kroopnick, P.M., 1985. The distribution of 13C of CO2 in the world oceans,Deep-Sea Res., Pt. A, 32:57-84.

    Leinen, M., and Heath, G.R., 1981. Sedimentary indicators of atmosphericactivity in the northern hemisphere during the Cenozoic. Palaeogeogn,Palaeoclimatol., PalaeoecoL, 36:1-21.

    Loubere, P., and Moss, K., 1986. Late Pliocene climatic change and the onsetof Northern Hemisphere glaciation as recorded in the northeast AtlanticOcean. Geol. Soc. Am. Bull., 97:818-828.

    Lutze, G.F., Pflaumann, U., and Weinholz, P., 1986. Jungquartàre Fluktu-ationen der benthischen Foraminiferenfaunen in Tiefseesedimenten vorNW Afrika—Eine Reaktion auf Produktivitàtsànderungen im Ober-flachenwasser. "Meteor" Forschungsergeb., 40:163-180.

    Lutze, G.F., and Thiel, H., 1989. Epibenthic foraminifera from elevatedmicrohabitats: Cibicidoides wuellerstorfi and Planulina ariminensis.J. Foraminiferal Res., 19:153-158.

    Maier-Reimer, E., Mikolajewicz, U., and Crowley, T., 1990. Ocean generalcirculation model sensitivity experiment with an open Central AmericanIsthmus. Paleoceanography, 5:349-366.

    Mammerickx, J., and Smith, S.M., 1985. Bathymetry of the North CentralPacific. Geol. Soc. Am., Map and Chart Sen, MC-52.

    Martin, J.H., 1990. Glacial-interglacial CO2 change: the iron hypothesis.Paleoceanography, 5:1-13.

    McCorkle, D.C., Keigwin, L.D., Corliss, B.H., and Emerson, S.R., 1990. Theinfluence of microhabitats on the carbon isotopic composition of deep-seabenthic foraminifera. Paleoceanography, 5:161-185.

    Miller, K.G. and Fairbanks, R.G., 1985. Oligocene to Miocene carbon isotopecycles and abyssal circulation changes. In Sundquist, E.T., and Broecker,W.S. (Eds.), The Carbon Cycle and Atmospheric CO2: Natural VariationsArchean to Present. Am. Geophys. Union, Geophys. Monogr. Sen,32:469-^86.

    Palmer, M.R., and Elderfield, H., 1985. Sr isotope composition of sea waterover the past 75 Myr. Nature, 314:526-528.

    Peterson, L.C., and Backman, J., 1990. Late Cenozoic carbonate accumulationand the history of the carbonate compensation depth in the westernequatorial Indian Ocean. In Duncan, R.A., Backman, J., Peterson, L.C.,Peterson, L.C., etal.,Proc. ODP, Sci. Results, 115: Washington (U.S. Govt.Printing Office), 467-507.

    Pisias, N.G., Shackleton, N.J., and Hall, M.A., 1985. Stable isotope andcalcium carbonate records from hydraulic piston cored Hole 574A: high-resolution records from the middle Miocene. In Mayer, L., Theyer,E, Thomas, E., et al., Init. Repts. DSDP, 85: Washington (U.S. Govt.Printing Office), 735-748.

    Prell, W.L., 1984. Covariance patterns of foraminiferal δ 1 8 θ : an evaluationof Pliocene ice volume changes near 3.2 million years ago. Science,226:692-693.

    Ramanathan, V., Barkstrom, B.R., and Harrison, E.F., 1989. Climate and theearth's radiation budget. Phys. Today, 42:22-32.

    Raymo, M.E., Ruddiman, W.F., Backman, J., Clement, B.M., and Martinson,D.G., 1989. Late Pliocene variation in Northern Hemisphere ice sheets andNorth Atlantic deep water circulation. Paleoceanography, 4:413^46.

    Raymo, M.E., Ruddiman, WE, Shackleton, N.J., and Oppo, D.W., 1990.Evolution of Atlantic-Pacific δ1 3C gradients over the last 2.5 m.y. EarthPlanet. Sci. Lett., 97:353-368.

    Redfield, A.C., Ketchum, B.H., and Richards, F.A., 1963. The influence oforganisms on the composition of sea-water. In Hill, M.N. (Ed.), The Sea(Vol. 2): New York (Wiley-Interscience), 26-77.

    Reid, J.L., 1965. Intermediate Waters of the Pacific Ocean. Johns HopkinsOceanogr. Stud., No. 2.

    Sarnthein, M., and Fenner, J., 1988. Global wind-induced change of deep-seasediment budgets, new ocean production and CO2 reservoirs ca. 3.3-2.35Ma BP. Philos. Trans. R. Soc. London Ser. B, 318:487-504.

    Sarnthein, M., and Tiedemann, R., 1988. Toward a high-resolution stableisotope stratigraphy of the last 3.4 million years: Sites 658 and 659 offnorthwest Africa. In Ruddiman, W., Sarnthein, M., et al., Proc. ODP, Sci.Results, 108: College Station, TX (Ocean Drilling Program), 167-185.

    Shackleton, N.J., 1977. Carbon-13 in Uvigerina: tropical rainforest history andthe equatorial Pacific carbonate dissolution cycles. In Andersen, N.R., andMalahoff, A. (Eds.), The Fate of Fossil Fuel CO2 in the Oceans: New York(Plenum), 401^27.

    Shackleton, N.J., Backman, J., Zimmerman, H., Kent, D.V., Hall, M.A.,Roberts, D.G., Schnitker, D., Baldauf, J.G., Desprairies, A., Hom-righausen, R., Huddlestun, P., Keene, J.B., Kaltenback, A.J., Krumsiek,K.A.O., Morton, A.C., Murray, J.W., and Westburg-Smith, J., 1984. Oxy-gen isotope calibration of the onset of ice-rafting and history of glaciationin the North Atlantic region. Nature, 307:620-623.

    Shackleton, N.J., Berger, A., and Peltier, W.R., 1990. An alternative astronomi-cal calibration of the lower Pleistocene time scale based on ODP Site 677.Trans. R. Soc. Edinburgh: Earth Sci., 81:251-261.

    Shackleton, N.J., and Hall, M.A., 1984. Oxygen and carbon isotope stratigra-phy of Deep Sea Drilling Project Hole 552A: Plio-Pleistocene glacialhistory. In Roberts, D.G., Schnitker, D., et al., Init. Repts. DSDP, 81:Washington (U.S. Govt. Printing Office), 599-609.

    Shackleton, N.J., Hall, M.A., Line, J., and Shuxi, C, 1983. Carbon isotopedata in Core VI9-30 confirm reduced carbon dioxide concentration in theice age atmosphere. Nature, 306:319-322.

    Shackleton, N.J., Imbrie, J., and Hall, M., 1983. Oxygen and carbon isotoperecord of east Pacific Core VI9-30: implications for the formation of deepwater in the late Pleistocene North Atlantic. Earth Planet. Sci. Lett.,65:233-266.

    Shackleton, N.J., and Opdyke, N.D., 1973. Oxygen isotope and paleomagneticstratigraphy of equatorial Pacific Core V28-238: oxygen isotope tempera-tures and ice volumes on a 105 and 106 year scale. Quat. Res., 3:39-55.

    , 1976. Oxygen-isotope and paleomagnetic stratigraphy of PacificCore V28-239 late Pliocene to latest Pleistocene. In Cline, R.M., and Hays,J.D. (Eds.), Investigation of Late Quaternary Paleoceanography andPaleoclimatology. Mem.—Geol. Soc. Am., 145:449^-64.

    Shackleton, N.J., and Vincent, E., 1978. Oxygen and carbon isotope studies inRecent foraminifera from the southwest Indian Ocean. Mar. Micropaleon-tol, 3:1-13.

    Shipboard Scientific Party, 1986. Site 586. In Moberly, R., Schlanger, S.O., et al.,Init. Repts. DSDP, 89: Washington (U.S. Govt. Printing Office), 213-235.

    Spero, H.J., and Williams, D.F., 1988. Extracting environmental informationfrom planktonic foraminiferal δ 1 3C data. Nature, 335:717-719.

    Streeter, S.S., and Shackleton, N.J., 1979. Paleocirculation of the deep NorthAtlantic: 15,000-year record of benthic foraminifera and oxygen-18. Sci-ence, 203:168-171.

    Sverdrup, H.U., Johnson, M.W., and Fleming, R.H., 1942. The Oceans:Their Physics, Chemistry and General Biology: Englewood Cliffs, NJ(Prentice-Hall).

    Tappan, H., 1968. Primary production, isotopes, extinctions and the atmos-phere. Palaeogeogr., Palaeoclimatol., PalaeoecoL, 4:187-210.

    Tiedemann, R., 1991. Acht Millionen Jahre Klimageschichte von NordwestAfrika und Palào-Ozeanographie des angrenzenden Atlantiks: Hochau-flösende Zeitreihen von ODP-Sites 658-661. Ber. Repts., Geol.-Palaont.Inst. Univ. Kiel, 46:1-190.

    Vincent, E., and Berger, W.H., 1985. Carbon dioxide and polar cooling in theMiocene: the Monterey hypothesis. In Sundquist, E.T., and Broecker,W.S. (Eds.), The Carbon Cycle and Atmospheric CO2: Natural Vari-ations Archean to Present. Am. Geophys. Union, Geophys. Monogr. Ser.,32:455-468.

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

    , 1981. Stable isotope composition of benthic foraminifera from theequatorial Pacific. Nature, 289:639-643.

    Wefer, G., and Berger, W.H., 1991. Isotope paleontology: growth and compo-sition of extant calcareous species. Mar. Geol., 100:207-248.

    Whitman, J.M., 1989. Stable isotope record of foraminifera from Ontong JavaPlateau for the last 6 million years, DSDP Site 586 [Ph.D. dissert.]. Univ.of California, San Diego.

    Whitman, J.M., and Berger, W.H., 1992. Pliocene-Pleistocene oxygen isotoperecord Site 586, Ontong Java Plateau. Mar. Micropaleontol., 18:171-198.

    Williams, D.F., Sommer, M.A., and Bender, M.L., 1977. Carbon isotopiccompositions of recent planktonic foraminifera of the Indian Ocean. EarthPlanet. Sci. Lett., 36:391^03.

    Winterer, E.L., Riedel, W.R., et al., Init. Repts. DSDP, 7: Washington (U.S.Govt. Printing Office).

    Woodruff, F, and Savin, S.M., 1985. δ 1 3C values of Miocene Pacific benthicforaminifera: correlations with sea level and biological productivity. Ge-ology, 13:119-122.

    Woodruff, J., Savin, S.M., and Douglas, R.G., 1980. Biological fractionationof oxygen and carbon isotopes by recent benthic foraminifera. Mar.Micropaleontol., 5:3-11.

    344

  • PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD

    Wu, G., and Berger, W.H., 1989. Planktonic foraminifera: differential dissolu-tion and the Quaternary stable isotope record in the west-equatorial Pacific.Paleoceanography, 4:181-198.

    , 1991. Pleistocene δ 1 8 θ records from Ontong-Java Plateau: effectsof winnowing and dissolution. Mar. Geol, 96:193-209.

    Wu, G., Herguera, J.C., and Berger, W.H., 1990. Differential dissolution:modification of late Pleistocene oxygen isotope records in the westernequatorial Pacific. Paleoceanography, 5:581-594.

    Zahn, R., Winn, K., and Sarnthein, ML, 1986. Benthic foraminiferal δ1 3C andaccumulation rates of organic carbon: Uvigerina peregrina group andCibicidoides wuellerstorfi• Paleoceanography, 1:27-42.

    Date of initial receipt: 9 March 1992Date of acceptance: 31 August 1992Ms 130B-030

    early Pliocene late Miocene

    1 2 3 4 5 6 7Age (Ma)

    Figure 11. Comparison of North Atlantic isotope record (Site 552) with theOntong Java record (Site 586). Data for Hole 552A from Shackleton and Hall(1984), Shackleton et al. (1984), Keigwin (1984), and Keigwin et al. (1987),The sampling interval at Hole 552A was 10 cm in the upper and lower sectionsof the hole; between 3.5 and 5.0 Ma, the sampling interval was generally 1.5m, although sometimes it was larger. Data given in Shackleton and Hall (1984)for various benthic species were converted to P. wuellerstorfi equivalents usingaverage offset values. Correlation of the records is based on the time scale ofBerggren et al. (1985) and on available paleomagnetic data (Shackleton andHall, 1984; Keigwin et al., 1987). Constant sedimentation rates were assumedbetween age control points, and data were smoothed by calculating slidingaverages. Note increased asymmetry after the major cooling event (Event A).

    345

  • J.M. WHITMAN, W.H. BERGER

    APPENDIX A

    Grain-size Data, Sand Fractions

    Core, section,interval (cm)

    89-586-1-1,49-522-1, 50-522-2, 55-572-3, 55-572-4, 55-572-5, 55-572-6, 55-573-1, 50-523-2, 49-503-3, 50-523-4, 50-523-5, 50-523-6, 51-534-1,50-524-2,51-534-3, 50-524-4, 50-524-5, 50-524-6,51-535-1, 60-625-2, 60-625-3, 58-605-4, 60-625-5, 60-625-6, 60-62

    89-586A-1-1, 50-521-2, 48-501-3, 50-521-4, 50-521-5, 50-521-6,48-502-1, 50-522-2,48-502-3, 50-522-4, 50-522-5, 50-522-6, 48-503-1, 50-523-2, 48-503-3, 50-523-4, 50-523-5, 50-523-6, 48-504-1, 50-524-2,48-504-3, 50-524-4, 50-524-5, 50-524-6, 48-505-1,50-525-2, 48-505-3, 50-525-4, 50-525-5, 50-525-6,48-506-1,50-526-2, 48-506-3, 50-526-4, 50-526-5, 50-526-6, 50-52

    Depth(m)

    0.501.813.364.866.367.869.36

    11.3212.8114.3115.8117.3118.8220.8122.3223.8125.3126.8128.3230.4131.9133.3934.9136.4137.91

    39.8141.2942.8144.3145.8147.2949.4150.8952.4153.9155.4156.8959.0160.4962.0163.5165.0166.4968.6170.0971.6173.1174.6476.0978.2179.6981.2182.7184.2185.6987.8189.2990.8192.3193.8195.31

    Age(m.y.)

    0.020.090.160.230.300.370.440.540.610.680.750.810.880.971.031.101.171.231.301.391.461.521.591.651.72

    1.801.871.932.002.072.132.222.292.362.422.492.552.652.712.782.842.912.953.013.053.093.143.183.223.283.333.373.413.453.503.563.603.643.693.733.77

    Weight(g)

    4.95825.34622.43184.45595.50783.99963.70276.13485.90645.62634.06815.64864.23196.90077.28066.40166.32817.10056.84037.74267.30196.97586.48036.55776.3180

    6.31865.28065.02814.68826.20075.49055.61574.20934.79544.45445.73057.14204.93875.70805.37983.75783.43505.63364.29466.46695.71237.14254.98277.52696.13015.86565.00955.29236.16395.02765.16385.71774.99694.55103.07555.4780

    >63 µm

    32.735.433.627.534.321.840.827.421.929.0

    8.541.631.131.231.527.929.224.324.721.823.627.629.639.022.3

    16.329.113.423.318.718.936.627.3

    6.318.812.015.321.213.111.122.414.824.321.522.226.728.122.920.223.421.523.812.628.026.128.733.018.327.326.117.0

    Weight percent>149µm

    23.225.724.021.425.216.029.118.814.617.16.3

    31.523.022.523.621.422.617.416.315.915.519.222.629.916.4

    11.423.59.4

    17.212.812.826.118.83.8

    12.37.49.9

    13.07.67.0

    15.410.216.914.015.218.319.815.512.915.813.615.67.0

    18.316.919.222.311.818.017.110.7

    >250 µm

    19.019.217.217.420.112.921.914.510.812.05.1

    25.118.117.017.816.316.712.410.612.110.812.317.423.312.3

    8.118.66.4

    12.59.59.5

    19.512.52.48.75.16.98.55.25.1

    10.97.7

    12.610.211.213.315.011.59.2

    12.210.011.35.0

    12.712.914.316.88.9

    13.513.47.7

    Core, section,interval (cm)

    7-1,50-527-2, 48-507-3, 50-527-4, 50-527-5, 50-527-6, 48-508-1,50-528-2, 48-508-3, 50-528-4, 50-528-5, 50-52

    89-586-8-6, 48-509-1, 50-529-2, 48-509-3, 50-529-4, 50-529-5, 50-529-6, 48-50

    10-1,50-5210-2, 48-5010-3, 50-5210-4, 50-5210-5, 50-5210-6, 48-5011-1,50-5211-2,48-5011-3,50-5211-4,50-5211-5,50-5211-6,48-5012-1, 50-5212-2, 48-5012-3, 50-5212-4, 50-5212-5, 50-5212-6, 48-5013-1, 50-5213-2, 48-5013-3, 50-5213-4, 50-5213-5, 50-5214-1,50-5214-2, 48-5014-3, 50-5215-1, 50-5215-2, 48-5015-3, 50-5215-4, 50-5215-5, 50-5215-6, 50-5216-1,50-5216-2, 50-5216-3, 50-5216-4, 50-5216-5, 50-5216-6, 50-5217-1,50-5217-2, 50-5217-3, 50-5217-4, 50-5217-5, 50-5217-6, 50-52

    Depth(m)

    97.4198.89

    100.41101.91103.41104.89107.01108.49110.01111.51113.01

    114.49116.61118.09119.61121.11122.61124.09126.21127.69129.21130.71132.21133.69135.81137.29138.81140.31141.81143.31145.41146.89148.41149.91151.41152.89155.01156.49158.01159.51161.01162.80163.79165.31167.21168.69170.21171.71173.21174.71175.70177.71179.21180.71182.21183.71185.81187.31188.81190.31191.81193.31

    Age(m.y.)

    3.833.863.903.933.974.004.054.094.124.164.19

    4.234.284.314.354.384.424.454.504.544.574.614.64 •4.684.734.764.804.834.874.904.954.995.025.065.095.135.185.215.255.285.325.365.385.425.465.505.535.575.615.645.665.715.755.785.825.855.905.945.976.016.046.08

    Weight(g)

    6.29217.51316.22167.07666.18066.48708.51369.00325.49698.00019.2394

    6.95504.98664.96084.83773.91285.23633.81046.10916.72125.24926.45906.98676.27245.62025.50765.72704.83327.55426.50455.00235.81494.57864.17513.61594.03986.63834.88215.41003.16056.61196.40876.72746.29475.73036.21325.59135.62395.67346.12482.93804.63054.01923.30113.39383.68905.33674.12886.34193.84033.92665.4382

    >63 µm

    23.730.521.621.622.514.725.223.516.913.121.4

    28.713.818.023.219.618.617.919.822.723.816.127.327.127.138.727.627.921.021.920.214.714.910.715.919.221.720.016.011.517.117.417.718.816.418.325.018.818.322.121.523.325.520.618.020.521.618.425.523.724.415.2

    Weight percent>149 µm

    13.920.414.014.114.08.7

    16.814.110.58.3

    12.9

    19.78.2

    11.614.512.011.811.413.112.813.89.3

    16.316.715.426.514.716.711.710.812.28.39.15.88.18.7

    13.311.19.47.1

    11.011.610.09.97.5

    10.416.09.6

    11.013.413.815.318.113.910.513.714.513.917.016.616.310.3

    >250 µm

    8.814.510.210.410.15.6

    12.19.86.55.58.4

    13.34.96.88.37.57.77.79.57.78.55.7

    10.011.09.6

    19.88.4

    11.27.46.78.75.26.03.64.64.47.96.16.04.77.68.36.76.54.86.8

    10.95.87.09.19.0

    10.412.99.86.38.79.17.3

    11.311.611.07.2

    346

  • PLIOCENE-PLEISTOCENE CARBON ISOTOPE RECORD

    APPENDIX B

    Benthic Foraminifer Counts

    Core, section,interval (cm)

    89-586-1-1,49-522-1, 50-522-2, 55-572-3, 55-572-4, 55-572-5, 55-572-6, 55-573-1,50-523-2, 49-503-3, 50-523-4, 50-523-5, 50-523-6,51-534-1, 50-524-2,51-534-3, 50-524-4, 50-524-5, 50-524-6,51-535-1,60-625-2, 60-625-3, 58-605-4, 60-625-5, 60-625-6, 60-62

    89-586A-1-1, 50-521-2, 48-501-3, 50-521-4, 50-521-5, 50-521-6,48-502-1, 50-522-2,48-502-3, 50-522-4, 50-522-5, 50-522-6, 48-503-1,50-523-2,48-503-3, 50-523-4, 50-523-5, 50-523-6,48-504-1, 50-524-2,48-504-3, 50-524-4, 50-524-5, 50-524-6, 48-505-1,50-525-2, 48-505-3, 50-525-4, 50-525-5, 50-525-6,48-506-1, 50-526-2,48-506-3, 50-52

    Depth(m)

    0.501.813.364.866.367.869.36

    11.3212.8114.3115.8117.3118.8220.8122.3223.8125.3126.8128.3230.4131.9133.3934.9136.4137.91

    39.8141.2942.8144.3145.8147.2949.4150.8952.4153.9155.4156.8959.0160.4962.0163.5165.0166.4968.6170.0971.6173.1174.6476.0978.2179.6981.2182.7184.2185.6987.8189.2990.81

    Age(m.y.)

    0.020.090.160.230.300.370.440.540.610.680.750.810.880.971.031.101.171.231.301.391.461.521.591.651.72

    1.801.871.932.002.072.132.222.292.362.422.492.552.652.712.782.842.912.953.013.053.093.143.183.223.283.333.373.413.453.503.563.603.64

    Weight(g)

    4.95825.34622.43184.45595.50783.99963.70276.13485.90645.62634.06815.64864.23196.90077.28066.40166.32817.10056.84037.74267.30196.97586.48036.55776.3180

    6.31865.28065.02814.68826.20075.49055.61574.20934.79544.45445.73057.14204.93875.70805.37983.75783.43505.63364.29466.46695.71237.14254.98277.52696.13015.86565.00955.29236.16395.02765.16385.71774.9969

    Benthic foraminifers (N)

    Total

    94702694

    105666688

    14010117413299

    13619324911721512023092

    13887

    109122

    12210717497

    13510417265

    12266

    10514285

    1031017681

    10886

    18510511410810171765499977751

    106151

    Pw

    10117

    13201310141817132421161726

    82917251417145

    12

    7858

    1879282

    1074562385386

    117527

    119723

    31

    Ou

    9832878583888

    151914122711188

    169

    118

    153

    115

    2010167

    108

    14139

    10122684

    176

    174

    10852574323

    Uv

    930

    361777

    12331661182

    124651

    3399

    19626

    1015

    210

    39262

    129

    516

    1716127

    136

    14146

    27121

    13412117

    10111222

    Core, section,interval (cm)

    6-4, 50-526-5, 50-526-6, 50-527-1,50-527-2, 48-507-3, 50-527-4, 50-527-5, 50-527_6, 48-508-1,50-528-2,48-508-3, 50-528-4, 50-52

    89-5 86 A-8-5, 50-528-6, 48-509-1, 50-529-2, 48-509-3, 50-529-4, 50-529-5, 50-529-6, 48-50

    10-1,50-5210-2, 48-5010-3, 50-5210-4, 50-5210-5, 50-5210-6, 48-5011-1,50-5211-2,48-5011-3,50-5211-4,50-5211-5,50-5211-6,48-5012-1, 50-5212-2,48-5012-3, 50-5212-4, 50-5212-5, 50-5212-6, 48-5013-1, 50-5213-2, 48-5013-3, 50-5213-4, 50-5213-5, 50-5214-1, 50-5214-2, 48-5014-3, 50-5215-1, 50-5215-2, 48-5015-3, 50-5215-4, 50-5215-5, 50-5215-6, 50-5216-1, 50-5216-2, 50-5216-3, 50-5216-4, 50-5216-5, 50-5216-6, 50-52

    Depth(m)

    92.3193.8195.3197.4198.89

    100.41101.91103.41104.89107.01108.49110.01111.51

    113.01114.49116.61118.09119.61121.11122.61124.09126.21127.69129.21130.71132.21133.69135.81137.29138.81140.31141.81143.31145.41146.89148.41149.91151.41152.89155.01156.49158.01159.51161.01162.80163.79165.31167.21168.69170.21171.71173.21174.71175.70177.71179.21180.71182.21183.71

    Age(m.y.)

    3.693.733.773.833.863.903.933.974.004.054.094.124.16

    4.194.234.284.314.354.384.424.454.504.544.574.614.644.684.734.764.804.834.874.904.954.995.025.065.095.135.185.215.255.285.325.365.385.425.465.505.535.575.615.645.665.715.755.785.825.85

    Weight(g)

    4.55103.07555.47806.29217.51316.22167.07666.18066.48708.51369.00325.49698.0001

    9.23946.95504.98664.96084.83773.91285.23633.81046.10916.72125.24926.45906.98676.27245.62025.50765.72704.83327.55426.50455.00235.81494.57864.17513.61594.03986.63834.88215.41003.16056.61196.40876.72746.29475.73036.21325.59135.62395.67346.12482.93804.63054.01923.30113.39383.6890

    Benthic foraminifers (TV)

    Total

    6264

    12078

    10061

    143967663

    1467895

    1057468607254654664674789695339646049717873

    1016345454384575969

    11118885715889542749

    109424497393856

    Pw

    81282

    104773

    1117

    17

    124584953435

    134345

    139

    152016168688

    2311121012187035634

    27513832

    Ou

    46

    13787

    18959

    121013

    108322585556385025395

    105443234357

    151520i

    098027315

    Uv

    01

    122228211243

    3283540010030101023436

    13210002

    242348

    85

    1524611

    122

    1218102

    Notes: Pw = Planulina wuellerstorfi, Ou = Oridorsalis umbonatus, and Uv = Uvigerina spp. N = number of samples.

    347

  • J.M. WHITMAN, W.H. BERGER

    APPENDIX C

    Isotope Data (‰), Four Foraminifer Taxa

    Depth(m)

    0.501.813.364.866.367.869.36

    11.3211.3212.8112.8114.3114.3115.8115.8117.3117.3118.8220.8120.8122.3223.8125.3125.3126.8128.3230.4131.9133.3934.9136.4137.9139.8139.8141.2942.8144.3145.8147.2949.4150.8950.8952.4153.9155.4156.8956.8959.0160.4962.0163.5163.5165.0166.4968.6170.0971.6173.1174.6476.0978.2178.2179.6981.2182.7184.2185.6985.6987.8187.8189.29

    Age(m.y.)

    0.020.090.160.230.300.370.440.540.540.610.610.680.680.750.750.810.810.880.970.971.031.101.171.171.231.301.391.461.521.591.651.721.801.801.871.932.002.072.132.222.292.292.362.422.492.552.552.652.712.782.842.842.912.953.013.053.093.143.183.223.283.283.333.373.413.453.503.503.563.563.60

    G. sacculifer

    δ 8 θ

    -1.48-0.83-1.49-0.83-1.07-1.23-1.05-1.29

    -1.44

    -1.13

    -0.91

    -1.36

    -1.16-1.75

    -1.06-1.28-1.28-1.64-1.23-1.33-1.36-1.15-1.17-1.30-0.92-1.27-0.83-1.24-1.31-1.00-1.16-1.13-1.37-1.10-1.20-1.54-1.00-1.16-1.01-1.33-1.26-1.27-1.26-1.49-1.43-1.67-1.53-1.53-1.32-1.44-1.35-1.46-1.41-1.47-1.36-1.34-1.29-1.58-1.77-1.48-1.47-1.76-1.73-1.51-1.89

    δ 1 3 c

    2.001.591.681.821.701.721.361.80

    1.29

    0.78

    1.53

    1.68

    1.611.50

    1.901.641.611.611.611.791.881.781.541.731.621.481.671.561.941.611.762.171.961.532.041.671.901.711.761.821.382.022.081.942.081.971.721.792.122.051.771.881.841.671.841.881.661.831.961.961.881.571.711.741.63

    Pulleniatim

    δ'8O

    -0.77-0.13-0.75-0.24-0.68-0.90-0.73-0.84-0.80-0.96-1.10-0.64-0.58-0.18-0.28-0.87-0.91-0.79-0.80-0.88-1.02-0.81-1.16

    -0.47-1.03-0.43-0.88-0.68-0.90-0.64-0.98-0.87

    -1.18-0.85-1.04-1.05-1.23-1.03-1.10

    -0.89-0.97-0.98-1.06

    -1.08-1.12-1.24-1.18

    -1.26-1.21-1.27-1.20-1.22-1.28-1.25-1.30-1.24

    -1.26-1.29-1.25-1.43-1.40

    -1.48

    -1.51

    spp.

    δ"C

    1.241.100.971.351.090.880.620.981.010.770.610.640.521.291.150.591.000.670.610.561.130.540.99

    1.050.830.860.830.480.890.760.770.80

    0.840.390.510.640.870.430.78

    0.700.820.740.60

    0.850.910.660.84

    0.540.630.750.720.650.670.810.420.74

    0.750.860.870.730.63

    0.73

    0.45

    P. wuellerstorfi

    δ 1 8 θ

    2.903.613.043.693.303.293.203.08

    3.29

    3.54

    3.69

    2.83

    3.102.66

    2.853.232.94

    3.372.983.393.012.922.923.183.243.06

    2.75

    2.883.092.983.10

    3.15

    3.21

    2.893.10

    2.80

    2.442.35

    2.652.402.232.552.62

    2.742.682.402.19

    2.42

    δ 1 3 c

    0.08-0.49-0.15-0.34-0.03-0.15-0.17-0.21

    -0.32

    -0.41

    -0.21

    -0.09

    -0.14-0.18

    -0.08-0.34

    0.07

    -0.13-0.10-0.14-0.17-0.36-0.16-0.35-0.31-0.14

    -0.13

    -0.26-0.27

    0.14-0.31

    -0.16

    -0.19

    0.07-0.19-0.18

    -0.040.17

    0.05-0.01-0.03-0.13

    0.13

    0.33-0.04

    0.15-0.01

    O. umbonatus

    δ 1 8 o

    3.734.54

    3.933.923.833.93

    3.86

    4.09

    3.69

    3.673.22

    3.674.13

    3.46

    3.893.654.12

    3.783.673.733.933.773.58

    3.474.10

    3.443.963.603.693.43

    3.903.403.683.63

    3.323.713.47

    3.183.17

    3.103.393.133.063.053.21

    3.24

    3.303.082.86

    δ 1 3 c

    -1.15-1.34

    -1.58-1.15-1.55-1.41

    -1.13

    -1.58

    -1.22

    -1.18-1.58

    -1.20-1.02-0.73

    -1.22-1.14-1.32-1.26-1.15-1.07-1.08-0.98-1.22

    -0.90-1.50-1.25-1.19-0.87-1.69-1.17

    -1.35-1.30-1.62-1.19

    -1.25-1.22-1.29

    -1.27-1.03

    -1.14-1.06-1.19-1.39-1.11-1.03

    -1.07

    -1.00-0.87-1.17

    Depth(m)

    90.8192.3193.8193.8195.3195.3197.4197.4198.89

    100.41101.91103.41104.89107.01107.01108.49110.01111.51113.01114.49116.61118.09119.61121.11122.61124.09126.21127.69129.21130.71132.21133.69135.81137.29138.81140.31141.81143.31145.41146.89148.41149.91151.41152.89155.01156.49158.01159.51161.01162.80163.79165.31167.21168.69170.21171.71173.21174.71175.70177.71179.21180.71182.21183.71185.81187.31188.81190.31191.81193.31

    Age(m.y.)

    3.643.693.733.733.773.773.833.833.863.903.933.974.004.054.054.094.124.164.194.234.284.314.354.384.424.454.504.544.574.614.644.684.734.764.804.834.874.904.954.995.025.065.095.135.185.215.255.285.325.365.385.425.465.505.535.575.615.645.665.715.755.785.825.855.905.945.976.016.046.08

    G. sacculifer

    δ 1 8 θ

    -1.73-1.80-1.73

    -1.65

    -1.42-1.55-1.65-1.67-1.50-1.70-1.62-1.30-1.55-1.39-1.65-1.49-1.54-1.28-1.18-1.24-1.29-1.31-1.17-1.11-1.17-1.22-1.24-1.24-1.08-1.10-1.24-1.34-1.31-1.34-1.22-1.38-1.19-1.33-1.17-1.24-1.26-1.14-0.97-1.20-1.06-1.40-1.17-1.44-1.21-1.20-1.20-1.18-1.19-1.21-1.20-1.16-1.22-1.32-1.23-1.24-1.24-1.20-1.19-1.18-1.10-1.13.-1.05-1.06

    δ 1 3 c

    1.781.611.85

    1.50

    2.032.071.711.431.501.541.931.662.072.011.941.791.732.021.812.201.952.051.831.751.842.151.911.582.171.902.122.252.092.102.18

    2.222.512.372.16

    2.092.292.311.802.062.072.012.111.541.922.062.352.231.952.092.431.812.252.472.052.101.972.301.752.012.001.991.741.97

    PuUeniatina spp.

    δ 8 θ

    -1.46-1.41-1.23-1.15-1.35-1.35-1.20

    -1.29-1.43-1.33-1.43-1.40-0.99

    -1.04-1.00-0.95-1.10-1.00-0.85-0.94-0.87-0.99-0.93-0.91-1.09-0.99-1.01-1.18-1.09-1.11-1.00-0.95-1.10-0.94-0.95-1.14-0.96-1.21-0.91-0.90-1.00-0.95-0.79-1.01-0.95-1.26-0.71-0.92-0.84-1.07-0.88-0.83-0.87-0.74-0.75-0.60-0.86-0.67-0.80-0.80

    δ 1 3 c

    0.760.620.810.600.550.530.60

    0.640.510.850.580.741.04

    0.930.830.860.860.850.881.040.690.710.530.620.640.620.690.380.810.710.941.140.980.960.950.971.280.740.390.660.860.930.550.510.780.610.870.881.071.031.251.251.100.961.290.741.081.231.081.17

    P. wuellerstorfi

    δ 1 8 θ

    2.552.622.46

    2.60

    2.372.282.632.322.442.54

    2.44

    2.392.362.442.712.412.382.422.58

    2.472.402.27

    2.532.572.412.302.452.372.432.462.812.802.552.402.802.602.382.502.412.482.42

    2.782.69

    2.612.81

    2.65

    2.442.322.49

    2.14

    δ 1 3 c

    -0.06-0.15-0.06

    0.18

    0.02-0.20-0.13-0.15-0.01

    0.23

    0.17

    0.060.04

    -0.020.000.300.290.00

    -0.08

    0.02-0.22

    0.19

    0.340.370.310.400.340.120.400.24

    -0.10-0.02

    0.100.19

    -0.150.030.10

    -0.130.01

    -0.310.10

    0.200.28

    -0.050.22

    0.24

    -0.13-O.03-0.02

    -0.15

    0. umbonatus

    δ 1 8 θ

    3.083.033.07

    3.23

    3.02

    3.122.923.172.832.933.06

    2.963.153.213.083.10

    3.093.273.05

    2.922.98

    2.803.05

    2.81

    2.78

    3.272.952.99

    3.04

    3.10

    2.99

    δ 1 3 c

    -0.79-1.10-1.10

    -1.13

    -1.09

    -0.76-1.31-1.24-1.26-1.01-0.81

    -1.04-0.77-0.98-1.16-1.15

    -0.82-1.42-1.40

    -1.22-1.30

    -1.04-1.22

    -1.26

    -0.98

    -1.59-1.27-1.77

    -1.27

    -1.02

    -1.33

    348


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