The last interglacial ocean

QUATERNARY RESEARCH 21, 123-224 (1984)

The Last Interglacial Ocean CLIMAP Project Members

Coordination and Compilation: WILLIAM F. RUDDIMAN Editor: ROSE MARIE L. CLINE

TASKGROUPLEADERS

Antarctic Ocean: JAMES D. HAYS?

Indian Ocean: WARREN L. PRELL* North Atlantic Ocean: WILLIAM F. RUDDIMANt

Pacific Ocean: TED C. MOORERS South Atlantic Ocean: NILVA G. KIPP*

BARBARA E. MOLFINol Ice Sheet: GEORGE H. DENTONS

TERENCE J. HUGHES*

William L. Balsam# Charlotte A. Brunner** Jean-Claude Duplessyti- Ann G. Esmay? James L. Fastook$

John Imbrie* Lloyd D. KeigwinO Thomas B. Kelloggs Andrew McIntyret Robley K. Matthews*

Received August 10, 1982

Alan C. Mixi Joseph J. Morley? Nicholas J. Shackleton@ S. Stephen Streeter§§ Peter R. Thompsonii~

The final effort of the CLIMAP project was a study of the last interglaciation, a time of minimum ice volume some 122,000 yr ago coincident with the Substage 5e oxygen isotopic minimum. Based on detailed oxygen isotope analyses and biotic census counts in 52 cores across the world ocean, last interglacial sea-surface temperatures (SST) were compared with those today. There are small SST departures in the mid-latitude North Atlantic (warmer) and the Gulf of Mexico (cooler). The eastern boundary currents of the South Atlantic and Pacific oceans are marked by large SST anomalies in individual cores, but their interpretations are precluded by no-analog problems and by discordancies among estimates from different biotic groups. In general, the last interglacial ocean was not significantly different from the modern ocean. The relative sequencing of ice decay versus oceanic warming on the Stage 6/5 oxygen isotopic transition and of ice growth versus oceanic cooling on the Stage Se/5d transition was also studied. In most of the Southern Hemi- sphere, the oceanic response marked by the biotic census counts preceded (led) the global ice- volume response marked by the oxygen-isotope signal by several thousand years. The reverse pattern is evident in the North Atlantic Ocean and the Gulf of Mexico, where the oceanic response lagged that of global ice volume by several thousand years. As a result, the very warm temperatures associated with the last interglaciation were regionally diachronous by several thousand years.

* Brown University, Providence, Rhode Island 02912. I’ Lamont-Doherty Geological Observatory of Columbia University, Palisades, New York 10964. $ University of Maine, Orono, Maine 04473. 8 University of Rhode Island, Kingston, Rhode Island 02881. /I Exxon, Stratigraphic Prediction Section, Houston, Texas 77001. # Southampton College, Southampton, New York 11968. ** University of California, Berkeley, California 94720. t? Centre des Faibles RadioactivitCs, Laboratoire mixte CNRS-CEA, 91190 Gif sur Yvette, France. $$ Godwin Laboratory, Free School Lane, Cambridge CB2 3RS, England. 8§ Chevron, Le Habra, California 90631. (/(I Arco Oil & Gas Company, Dallas, Texas 75221.

123 0033-5894/84 $3.00 Copyright 0 1984 by the University of Washingtor All tights of reproduction in any form reserved.

124 CLIMAP PROJECT MEMBERS

These regional lead-lag relationships agree with those observed on other transitions and in Iong- term phase relationships; they cannot be explained simply as artifacts of bioturbational translations of the original signals.

Contenrs. Introduction. Data Base. Choice of cores. Oxygen isotopic ice volume record and chronology. Definition of the last interglaciation. Age of the last interglaciation. Sea-surface temperature estimates. Absolute abundance counts. Reconstruction of the last interglacial ocean. Methods. Choice of oxygen isotopic 5e level. Choice of modem SST value. Results. Lead-lag relutionships: SST and ice. Methods. Offsets in transition midpoints. Cross-correlation analysis. Results. Problems and complications. Validity of the 6t80/ice volume assumption. Temperature and dissolution overprints. Meltwater complications and oceanic mixing time. Summary of 6’80. Validity of the SST estimates. No-analog conditions. Discordant multiple estimates. Bioturbational and bottom-current mixing. Impact of mixing on Stage 5e SST estimates. Impact on lead-lag relationships. Summary of problems and their impacts. Stage 5e SST reconstruction. Lead-lag relationships. Other evidence 0ff3’~0/SST phasing. Evidence from other transitions. Comparison with long-term phase relationships. Implications of diachronous regional responses. Additional evidence of last interglnciul climate. Extent of last interglacial ice sheets. Correlation to pollen records. Other terrestrial sequences. Shallow marine sequences. Other oceanic evidence. Sum- mary. Conclusions. Acknowledgments. References.

INTRODUCTION

In the final three years (1977- 1980) of its decade-long existence, CLIMAP (Climate, Long-Range Investigation, Mapping and Prediction) focused a considerable part of its research effort on the last interglacial ocean. That study, reported here, had two major objectives: (1) to compare the last interglacial ocean with the modern ocean and (2) to define the sequence of regional oceanographic changes that occurred during the shift both into, and out of, the last period of minimal ice volume. The latter objective involved a study of leads and lags in estimated SST in various oceanic regions relative to the oxygen isotopic record of global ice volume.

DATA BASE

Choice of Cores

The last interglacial study is based on 52 cores (Fig. 1, Table 1). This coverage spans all the major oceans, with maximum coverage in the North and South Atlantic.

Three criteria primarily determined the final core coverage. First, because our stratigraphy is based entirely on oxygen isotopic records, we were restricted to regions of calcareous microfossils. Areas lacking foraminifera, like the high-latitude North and South Pacific and high-latitude South-

ern Ocean, are thus blank spaces on the map. Second, to minimize problems in- volving sediment mixing (bioturbation), we only included cores with sedimentation rates in excess of 1.0 cm/1000 yr. This diminished our coverage of mid-ocean waters in the central subtropical gyres, although temperature changes in such areas have generally been small anyway. Third, we sought to focus our major effort on regions of high-amplitude temperature change, both to maximize the signal-to-noise ratio in our temperature estimates and from the conviction that regions of maximum variation are inherently important parts of the climate system.

Finally, the preexisting state of stratigraphic knowledge influenced the final coverage. Regions with a long history of cu- mulative research effort (e.g., the North At- lantic Ocean) naturally emerged with a larger data base than areas where the basic stratigraphic framework had to be generated almost entirely within this project (e.g., the Indian Ocean).

Oxygen Isotopic Ice Volume Record and Chronology

We have relied on oxygen isotopic variations in foraminifera for (1) basic stratigraphic control and (2) a first-order indicator of global ice volume against which to

THE LAST INTERGLACIAL OCEAN

FIG. 1. Locations and names of cores used in the last interglacial project.

TABLE 1. CORE LOCATIONS AND DEFTHS

Core Latitude Longitude Depth

Cm) Core Latitude Longitude Depth

Cm)

Al80-73 Dl17 E49-18 K-11 K708- 1 Ml2392-1 MD73025 RC8-39 RCS-145 RC 10-65 RCll-86 RCll-120 RCll-210 RCl l-230 RCl2-294 RCl2-339 RCl3-205 RCl3-228 RCl3-229 RCl5-61 RC17-69 RCl7-98 TRl26-23 TRl26-29 v12-122 Vl8-68

0”lO’N 42”06’N 46”03 ‘S 71”47’N 50”OO’N 25”lO’N 43”49’S 42”53’S 33”35’N O”4l’N

35”47’S 43”3 1 ‘S

l”49’N 8”48’S

37=16’S 9’08’N 2”17’S

22”2O’S 25”3O’S 40”37’S 31”3O’S 13”13’S 20”29’N 21”26’N 17”OO’N 54”33’S

23”OO’W 3749 52”45’W 3300 90”09’E 3253

l”36’E 2900 23”45’W 4053 16”5l’W 2573 51”19’E 3284 42”2l’E 4330 62”23’W 2743

108”37’W 3588 18”27’E 2829 79”52’E 3135

140”03’W 4420 1 lO”48’W 3259 lo”O6’W 3308 90“02’E 3010 S”ll’E 3731

1 l”l2’E 3204 1 l”l8’E 4191 77”12’W 3771 32”36’E 3380 65”37’E 3409 95”37’W 2410 93”57’W 2700 74”24’W 2800 77”5l’W 3982

Vl9-29 v19-53 V21-146 V22-38 V22- 108 V22-174 V22-182 V22-196 V23-82 V25-59 V27-20 V27-86 V28-14 V28-56 V28-127 V28-238 V28-304 V28-345 V29-29 V29-179 v30-97 V32-126 V32-128 v34-88 Y71-6-12 Y72-11-l

3”35’S 17”Ol’S 37”4l’N 9”33’S

43”ll’S lo”O4’S O”33’S

13”50’N 52”35’N

l”22’N 54”OO’N 66”36’N 64”47’N 68”02’N 1 l”39’N 1”Ol’N

28”32’N 17”4O’S 5’07’N

44”OO’N 4l”OO’N 35”19’N 36”28’N 16”3l’N 16”26’S 43”lS’N

83”56’W 3157 113”3l’W 3058 163”02’E 3968 34”15’W 3797 3”15’W 4171

12”49’W 2630 17”16’W 3937 18”58’W 3728 21”56’W 3974 33”29’W 3824 46”12’W 3510

l”07’E 2900 29”34’W 1855 6”07’W 2941

80”08’W 3227 160”29’E 3120 134”08’E 2942 117”57’E 1904 77”35’E 2673 24”32’W 3331 32”56’W 3371

117”55’E 3870 117”lO’E 3623 59”32’E 2100 7753’W 2734

126”23’W 3000


phase local changes in SST. Both of these uses of the iY80 record derive from the fact that periodic preferential storage of at60 over #*O in ice sheets is the major control over al80 records in foraminifera (Shack- leton and Opdyke, 1973). As a result, the ice-volume signal dominant in all a’*0 curves from the deep sea is generally considered globally synchronous within the mixing time of the deep ocean, estimated to be <lOOO yr for the modern ocean.

Regional temperature and precipitation- evaporation balances also affect the S”O record (Emiliani, 1955). Benthonic foraminifera are generally preferred to planktonic foraminifera for 6t80 analysis, because the local climatic signals that can be strong at the sea surface are attenuated in the deep sea. In some low-latitude areas, however, surface temperature changes appear to have been sufficiently small to leave the ice-volume signal relatively intact (Shack- Ieton and Opdyke, 1973). Conversely, several recent studies have suggested that a significant part of the algO signal in benthonic foraminifera may reflect changes in bottom-water temperature (Duplessy, 1978; Duplessy et al., 1980; Broecker, 1981).

With these factors in mind, our oxygen isotopic strategy in the last interglacial project was to analyze benthonic foraminifera wherever possible; to supplement these records with planktonic foraminiferal curves in some areas; and to measure only planktonic foraminifera where benthonic foraminifera were unavailable. In some low-latitude regions, because of time and efficiency constraints, only planktonic foraminifera were analyzed.

Our final choices for 6t80 analysis in- clude planktonic foraminifera in 36 cores and benthonic foraminifera in 34 cores. Of the total 52 cores analyzed, 18 have both kinds of records. All al80 analyses for the 52 cores are listed in Appendix 1,’ with

’ Readers may request the appendixes on microfiche by written request from William E Ruddiman, Lamont-Doherty Geological Observatory of Co- lumbia University, Palisades, N.Y. 10964.

each value reported relative to PDB and the laboratory responsible for analysis of each core indicated.

Most species of benthonic foraminifera have been shown to calcify their tests out of oxygen isotopic equilibrium with am- bient sea water. Adjustments of these species to equilibrium have been proposed in a number of publications (Shackleton and Opdyke, 1973; Shackleton, 1974, 1977; Du- plessy et al., 1975; Streeter and Shack- leton, 1979; Duplessy, 1978; Belanger et al., 1981; Vincent et al., 1981; Graham et al., 1981). For several species, there is considerable range in the estimates of the necessary adjustments; Vincent et al. (1981) even suggest that the degree of disequilibrium for some species is variable in space and time. We have selected from among the several published adjustments the values shown in Table 2 for use in this study.

Adjusted Z1*O values are listed in Ap- pendix 1 to the right of the actual measured values. For the plotted records of each core, we used the adjusted values for the benthonic foraminifera noted above, and the unadjusted values for all planktonic foraminifera.

All 6t3C analyses from this project are listed in Appendix 1. We made no attempt to adjust these values toward an assumed equilibrium; all are reported relative to PDB. We have not attempted to interpret the 613C analyses (but see Duplessy et al., 1984).

Definition of the lust interglaciation. Two “typical” records of oxygen isotopic change in benthonic foraminifera extending from the present back through the last interglaciation are shown in Figure 2 (data from Shackleton, 1977). One is an Atlantic core from the west coast of Saharan Africa; the other is from the Panama Basin sector of the eastern equatorial Pacific. Both show similar isotopic changes. Characteristically, oxygen isotope Substage 5e emerges as the last time that isotopic values were as light as they are today at the core tops. This is widely interpreted as indicating that Sub-

THE LAST INTERGLACIAL OCEAN 127

TABLE 2. 6i*O ADJUSTMENTS USED FOR BENTHONIC FORAMINIFERA

Species Correction

(%d

Uvigerina peregrina Uvigerina senticosa Cibicides wuellerstorfi (-Planulina wuellerstorfi) Cibicides kullenbergi Melonis barleanum Melonis pompilioides (=Nonion spp.) Oridorsalis tener Oridorsalis umbonatus mixed Oridorsalis spp. Pyrgo murrhina Pyrgo depressa Pyrgo oblonga Pyrgo rotaliana Hoeglundina elegans Favocassidulina favus Globocassidulina subglobosa Gyroidina spp.

1 0.00 : j +0.64 Shackleton and Opdyke (1973)

+ 0.73 Graham et al. (1981)

+ 0.40 Streeter and Shackleton (1979)

+ 0.36’ +0.15a +0.136

0.00

-0.51”

-0.37” -0.34

0.00 -0.10

0.00 I

Source

Shackleton (1974)

Shackleton (unpublished) Belanger et al. (1981) Graham et al. (1981) Shackleton (1974)

Duplessy et a/. (1975)

Belanger et a/. (1981) Graham et al. (1981)

Shackleton (unpublished)

u Derived entirely through offset relative to C. wuellerstorfi. b Derived in part through offset relative to C. wuellerstorfi.

stage 5e was the last time that there was as small a volume of ice on earth as there is today.

Oxygen isotopic records published earlier indicated relatively little structure within “interglacial” isotopic Stage 5 (Em- iliani, 1955, 1966), but later work showed major positive isotopic excursions during Substages 5d and 5b with Substages 5c and 5a never returning to the extremely light 6180 values characteristic of Substage 5e (Ninkovich and Shackleton, 1975; Shack- leton, 1977). As a result, the concept of the last interglaciation shrank from the 52,000- yr length of Stage 5 to the 11,000 yr of Sub- stage 5e (Shackleton, 1969). Oxygen isotopic Substage 5e is thus our definition of the peak of the “last interglaciation,” the last time that there was as little ice as today.

Through the sequence of changes into and out of the last interglaciation, the two records in Figure 2 show an oxygen isotopic signature that is typical of that found in benthonic foraminifera from most of the world’s oceans. The values are characteristically around 5..0%0 in Stage 6. rise to

3.0%0 in Substage 5e, and then fall to 4.0%. in Substage 5d. Many other records show changes with smaller or larger amplitudes but a similar pattern. In some other cores, the total amplitude may be maintained but with the absolute values systematically shifted toward lighter or heavier values. Oxygen isotopic signals in surface-dwelling planktonic foraminifera often show somewhat smaller amplitudes of change. In cores with both benthonic and planktonic records, the two oxygen isotopic signals are usually in phase. In those cores where the two signals are out of phase, we relied on the benthonic foraminiferal signal rather than the planktonic. These, then, are the basic signals we have used to define the last interglacial level in these cores.

Age of the last interglaciation. The basic chronologic framework of the oxygen isotopic record of the last 130,000 yr was defined by the work of Broecker et al. (1968). They placed the Stage 5 boundaries at 127,000 and 75,000 yr B.I? , creating a time scale some 25% longer than that previously used by Emiliani (1%6). This assessment


lSOTOPlC STAGES

+5

v19-29

8’*0 vs. P.D.B. +4

Ml2392 -1

8”O vs. P. D.B. +4

I i

3 DEPTH I IN CORE (ml

4

MORE LESS MORE LESS GLOBAL- GLOBAL GLOBAL- GLOBAL

ICE ICE ICE ICE

FIG. 2. Oxygen isotopic records of the last 150,000 yr from Pacific core Vl9-29 and Atlantic core Ml2392-1 (data from Shackleton (1977)). Last interglacial level (Substage 5e) marks the last time that isotopic values were as light as they are today, suggesting global ice volumes at least as small as those today.

depended both on 231Pa/23c’Th dating within deep-sea cores and on correlation of isotopic Substages Sa, 5c, and 5e with Bar- bados high sea-level terraces I, II, and III. This chronology has passed the test of a decade’s vigorous application and is widely accepted. It has been corroborated by isotopic work on Barbados mollusks (Shack- leton and Matthews, 1977). It has also provided plausible links between orbitally controlled insolation changes and climatic responses on earth (Hays et al., 1976; Moore et al., 1977; Imbrie and Imbrie, 1980; Ruddiman and McIntyre, 1981a).

The major challenge to the Broecker/Ku chronology is the frequency of dates on high sea-level terraces in the range 140,000-130,000 yr B.P. Several papers argue that this concentration of U-series dates on apparently unaltered corals is real (Chappell, 1974; Moore, 1982). Moore

argues that a high stand of sea level 2 m above the modern levels occurred at -135,000 yr B.P.

The oxygen isotopic record, however, shows no large-amplitude negative excursions just prior to the Stage 60 boundary at -127,000 yr B.P. (Fig. 2); it certainly shows no full-scale excursion to full-interglacial values. Such a sea-level excursion could only exist within the Broecker/Ku time scale if oxygen isotopes were concluded to be totally insensitive as monitors of ice volume. This is not a reasonable conclusion.

Alternatively, these high sea levels might argue for a different time scale, with the lower boundary of Stage 5 at -140,000 instead of -127,000 yr B.P. This, however, would raise many other problems. It would mean that the large body of work that convincingly links orbital forcing and the cli-

matic response on earth at the Broecker/Ku three biotic components of the sediments: time scale over the last 130,000 yr is invalid foraminifera, radiolaria, or coccoliths, (2) and must be some totally fortuitous acci- selection or development of the appropriate dent. It would also require that a major de- regional transfer function to apply to the glaciation occurred at a time (140,000 yr biotic counts, and (3) estimation of SST for BP) when orbital forcing provides no un- the winter and summer seasons. usually large impetus; this contrasts poorly All biotic census counts for the 52 cores with the strong deglacial impetus of very in this project are listed in Appendix 2. high summer insolation in the Northern There are 39 cores with counts based on Hemisphere at 127,000 yr B.P. (Milanko- planktonic foraminifera, 18 cores with vitch, 1941). counts on radiolaria, and 6 cores with

We conclude that the Broecker/Ku time counts on coccoliths. Nine cores have scale is essentially correct and that the ter- counts from two biotic groups, and one has race dates at 140,000-130,000 yr BP. must three. be in error, probably due to contamination. As with previous CLIMAP publications, The possibility of contamination was not the emphasis on the three biotic groups rejected by Moore (1982). varies from region to region. The North At-

We thus use the estimate of Broecker and lantic and Indian Ocean exclusively rely on Ku that the midpoint age of Termination II foraminifera. The Antarctic and Pacific (the isotopic 6/5 boundary) is 127,000 yr tend to emphasize radiolaria, and the South B.P. This estimate recurs (to within 1000 yr) Atlantic utilizes all three biotic groups. in numerous recent minor revisions of the These choices reflect a variety of factors, late Quaternary time scale (Hays et al., including the preservation state of the 1976; Kominz et al., 1979; Morley and CaCO, and/or SiO, fraction, the regional di- Hays, 1981). For the isotopic 5e/5d versity trends of the various biotic groups boundary, we use the estimate of 116,000 in each of the modern oceans, and the de- yr B.P. from Shackleton (1969). This con- sirability of obtaining multiple biotic rec- strains isotopic Substage 5e to a length of ords in areas where one biotic group is sus- 11,000 yr, which is one-half of a preces- pect. sional cycle. And it places the midpoint of Typically, the biotic census counts are Substage 5e at about 122,000 yr B.P., under based on 2300 individuals; this range rep- the assumption of rough geometrical sym- resents the optimal compromise between metry. This value lies in the middle of the the conflicting goals of precision and effi- range of commonly cited dates between ciency. The transfer-function equations 120,000 and 125,000 yr B.P. (Broecker et used in this study are listed in Table 3, with al., 1968; Mesolella et al., 1969) and is pre- the publication in which the equation was cisely congruent with multiple dates ob- first described. Several of the equations tained on Hawaii (Ku et al., 1974) and Yu- used have not been published by the re- catan (Szabo et al., 1978). It also matches spective CLIMAP members and are the age derived from a model that estimates marked as “in preparation” (in prep). the response time of ice sheets to orbital The standard errors of estimate of the insolation forcing (Imbrie and Imbrie, seasonal SST are listed in Table 3. Most of 1980). the transfer functions have standard errors

Sea-Surface Temperature Estimates of estimate in the range l.o”-2.O”C, with an average of 1.5”C. This represents a rela-

The methodology involved in estimating tively small portion of the temperature SST for the last interglacial interval is es- signal found in high-latitude oceans, where sentially that published in various articles the total glacial/interglacial difference may in Cline and Hays (1976). It consists of exceed 10°C. This error of estimate may, three steps: (1) census counts of one of the however, equal the entire glacial/intergla-



TABLE 3. TRANSFER FUNCTIONS USED FOR EACH CORE AND SOURCE REFERENCES

Core

Transfer function

name Source

Seasonal standard error of estimate

TW

A180-73 D117 E49-18

K-11 K708- 1 M12392-1 MD73025

RC8-39

RC8-145 RClO-65 RCll-86

RCll-120

RCll-210 RCl l-230 RC12-294

RC12-339 RC13-205

RC13-228

RC13-229

RClS-61 RC 17-69 RC17-98 TR126-23 TR126-29 v12-122 V18-68

V19-29

v19-53 V21-146 V22-38

V22-108

v22- 174

V22-182

V22-196 V23-82

FA13 RAN3

FA13’ FA13’

RAN3

F12 RAN3

FA13 RP7 CA8 FA2OU7 F12 RAN3

RP7 RP8 CA8 FA20 FI2 FAZOUS RSAl CA8 FA2OU7 RSAl FA2OU7 RSAl RP8 F12 FI2 FG6 FG6 FA3 RAN3

RP7 RP8 FPl2D RP8 CA8 FA20 RAN3

CA8 FA20 CA8 FA20 FA13 FA3

Gardner (1973) 1.0 0.9 Kipp ( 1976) 1.2 1.4 J. Hays, J. Morley, L. Burckle,

D. Clarke, D. Cooke (in prep) Ruddiman and Glover (1975) Ruddiman and Glover ( 1975) Molina-Cruz and Thiede (1978) J. Hays, J. Morley, L. Burckle,

1.1 1.2 1.2 1.3

1.4 1.4 1.4 1.1

D. Clarke, D. Cooke (in prep) Hutson and Prell (1980) J. Hays, J. Morley, L. Burckle,

1.1 1.4 1.3(Aug) l.l(Feb)

D. Clarke, D. Cooke (in prep) Kipp (1976) Moore et al. (1980) B. Molfino and A. McIntyre (in prep) N. Kipp (in prep) Hutson and Prell (1980) J. Hays, J. Morley, L. Burckle,

1.1 1.4 1.2 1.4 1.8 1.5 2.0 1.8 1.4 1.8 1.3(Aug) 1. l(Feb)

D. Clarke, D. Cooke (in prep) Moore et a/. (1980) Moore ef al. (1980) B. Molfino and A. McIntyre (in prep) N. Kipp (in prep) Hutson and Prell (1980) N. Kipp (in prep) Morley (1979) B. Mollino and A. McIntyre (in prep) N. Kipp (in prep) Morley ( 1979) N. Kipp (in prep) Morley (1979) Moore et a[. (1980) Hutson and Prell (1980) Hutson and Prell (1980) Brunner (1979) Brunner (1979) Imbrie, Van Donk, and Kipp (1973) J. Hays, J. Morley, L. Burckle,

1.1 1.4 1.8 1.5 2.4 1.8 2.0 1.8 1.3 1.2 1.3(Aug) I .l(Feb) 1.5 1.3 1.4 1.4 2.0 1.8 1.4 1.8 1.4 1.4 1.4 1.8 1.4 1.4 2.4 1.8 1.3(Aug) 1 .l(Feb) 1.3(Aug) I.l(Feb) 1.0 0.9 1.0 0.9 1.5 1.7

D. Clarke, D. Cooke (in prep) Moore et al. (1980) (downcore) Moore et a/. (1980) (core top) Moore er al. (1980); Thompson (1977) Moore et al. (1980) B. Moltino and A. McIntyre (in prep) N. Kipp (in prep) J. Hays, J. Morley, L. Burckle,

1.1 1.4 1.8 1.5 2.4 1.8 3.0 1.5 2.4 1.8 2.0 1.8 1.3 1.2

D. Clarke, D. Cooke (in prep) B. Molfino and A. McIntyre (in prep) N. Kipp (in prep) B. Molfino and A. McIntyre (in prep) N. Kipp (in prep) Kipp (1976) Imbrie, Van Donk, and Kipp (1973)

1.1 1.4 2.0 1.8 1.3 1.2 2.0 1.8 1.3 1.2 1.2 1.4 1.5 1.7

T,


TABLE 3-Continued

Core

Transfer function

name Source

Seasonal standard error of estimate

TW T,

V25-59 v21-20 v21-86 V28-14 V28-56 V28-127 V28-238 V28-304 V28-345 V29-29 V29-179 v30-97 V32-126 V32-128 V34-88 Y71-6-12 Y72-11-l

FA13 FAl3’ FAl3’ FAl3’ FAl3’ FA13 FPl2E RP8 F12 F12 FA13’ FAl3’ FPl2E FPl2E F12 RP8 RP8

Kipp (1976) 1.2 1.4 Ruddiman and Glover (1975) 1.2 1.4 Ruddiman and Glover (1975) 1.2 1.4 Ruddiman and Glover (1975) 1.2 1.4 Ruddiman and Glover (1975) 1.2 1.4 Kipp (1976) 1.2 1.4 Thompson (1981) 2.5 1.5 Moore et al. (1980) 2.4 1.8 Hutson and Prell (1980) 1.3(Aug) l.l(Feb) Hutson and Prell (1980) 1.3(Aug) l.l(Feb) Ruddiman and Glover (1975) 1.2 1.4 Ruddiman and Glover (1975) 1.2 1.4 Thompson (1981) 2.5 1.5 Thompson (1981) 2.5 1.5 Hutson and Prell (1980) 1 .l(Feb) 1.3(Aug) Moore et a/. (1980) 2.4 1.8 Moore et al. (1980) 2.4 1.8

cial signal in some low-latitude oceanic re- dian Ocean now follows the Southern gions. Hemisphere caloric seasons.

Most of the transfer functions in Table 3 are “caloric” equations. This means that modern atlas SST values at all calibration sites are entered with the convention that the cold and warm seasons are called “winter” and “summer,” respectively, rather than being given monthly designations. This basically reflects the reversal of caloric seasons near the equator. For cores lying a few degrees north of the equator but following the Southern Hemisphere caloric seasons, the modern atlas SST values are again entered with the cold season as “winter” and conversely (McIntyre et al., 1976).

Indian Ocean transfer function F12 is the only “calendar” equation used (Hutson and Prell, 1980). This means that all modern atlas SST values are entered with the designation of a specific month (February or August). This is done because it is difficult to infer past locations of the thermal equator, which trends erratically even in the modern Indian Ocean due to summer upwelling. Much of the northwestern In-

All estimates of SST that appear in Fig- ures 3 through 54 are listed in Appendix 3. In most cases, seasonal SST designations follow the normal hemispheric conventions (e.g., “summer” means August in the Northern Hemisphere and February in the Southern Hemisphere). For several At- lantic and Pacific cores located north of the geographic equator but south of the caloric equator, the normal hemispheric conventions are reversed. In these cores “winter” means August and “summer” means Feb- ruary. Cores where this occurs are noted in Table 4. In several figures and tables perti- nent to the Indian Ocean, we have desig- nated the colder SST value as winter and the warmer as summer, with the August and February designations noted in the figure captions and table footnotes.

Absolute Abundance Counts

In the initial stages of research on the last interglacial project, bioturbational mixing of deep-sea sediments was viewed merely as a low-pass filter that smoothed the orig-

TABLE 4. STAGE 5e LEVEL (cm) AND SST VALUES IN “C FOR ISOTOPIC STAGE Se,

MODERN ATLAS AND CORE TOPS

Core 5e

level T,

Se

T,

Atlas Core top

T, T, TW TS

Al80-73 D117 E49-18 K-11 K708-1 M12392-1 MD73025 RC8-39

340 255 470 282

24.0 12.2

6.2 3.1

825 13.1 824 16.1

26.6 20.2

8.0 8.66

1351 826

4.4

RC8-145 193 RC lo-65 272 RCll-86 286

9.2(Aug) 5.0

17.7

RCll-120 420

RCll-210 185 RCll-230 232 RC 12-294 276

23.4 19.5 14.7 8.0(Aug) 6.9

25.6 26.3 15.3

19.9 21.4

6.1 lO.l(Feb) 6.9b

24.9 26.1 23.8 18.9 lZ.S(Feb)

9.3 27.7 26.4 20.3

13.6 19.1 RCl2-339 322 26.2(Aug) 26.7(Feb) RCl3-205 222 23.7 27.1

RCl3-228 620

RC13-229 290

26.4 22.4’ 16.4 23.56 21.3 23.9

RC15-61 580 RCl7-69 230 RCl7-98 302 TRl26-23 600 TRl26-29 820 v12-122 280 V18-68 450 V19-29 830 v19-53 219 V21-146 370 V22-38 228

22.2 16.6 11.2 18.1 15.8 18.5

3.6 22.1(Aug) 24.7(Aug) 21.0 21.7 23.6

6.4 23.6 27.0 17.3

V22-108 505 V22-174 280

V22-182 402

24.1 25.1

7.9 22.2 23.5 23.7

V22-196 V23-82 V25-59 V27-20 V27-86 V28-14 V28-56 V28-127 V28-238 V28-304 V28-345 V29-29 V29-179 v30-97 V32-126 V32-128 V34-88 Y71-6-12 Y72-11-l

567 765 338 494

21.5 22. I

9.6 23.6

6.3 321 6.0 515 8.4

0.1 24.1 27.0 23.1

235 528 203 430 532 415 515 574 204 291 500 156

1205

25.1(Aug) 25.8(Aug) 14.9 14.6 18.0 17.2 24.3(Feb) 15.6 Il.8

11.26 25.5(Feb) 27.7(Feb) 28.4 27.7 26.6

8.5 27.5 29.3b 24.gb 25.7 27.3 10.1

25.5 25.9 26.4 23.2 27.6 15.6 27.4 10.6 10.3b 12.8

6.1 27.8 29.8 26.1b 27.2(Feb) 26.4(Feb) 20.6 21.5 27.0b 26.7 26.O(Aug) 20.1 16.6b

24.9 27.5” 10.7 21.8

24.6 26.7

7.0 3.6

10.5 18.1 6.6 6.7(Aug)

8.1 7.9

16.3 21.7

9.5 10.6(Feb)

18.6 26.9 23.6 26.0” 15.2 20.0

9.0(Aug) lI.g(Feb)

2.2 11.0

7.5’ 17.0

PC PC

3.2 4.8 PC

5.8 7.6 R TW 16.5 23.9 PC

16.8 21.3 C TW 16.5 22.1 F TW

6.3(Aug) 9.1’(Feb) F PC 9.6 13.0 R TW

27.2 12.6 12.5 27.2(Aug) 21.7 21.2 14.7 15.3 15.3 16.7 14.6

27.9 PC 16.9 C TW 17.0 F TW 28.2’(Feb) TW 23.7 F TW 25.6’ R PC 20.3’ C TW 19.7’ F TW 21.4 R PC 21.6 F TW 19.2’ R TW

22.8 22.9 25.9

6.5 19.5 23.9 13.5 22.1 24.0

29.2’ 29.0’ 27.4

8.8’

8.2 23.9 24.3 21.9 23.9

25.8’,” 28.3 22.0 25.1 26.8 10.8

25.8 26.6 25.8 26.2

TW TW PC PC PC TW PC

C TW F TW

TW C PC F PC C PC F PC

9.3 26.0

4.6 7.3 8.5 1.0

14.0’ 27.4’ 10.8 11.5 12.4 6.8

25.4 27.4’ 30.6 29.9

TW TW PC TW TW TW PC TW

26.O(Aug) 28.O(Feb) PC 25.6(Aug) 25.3’CFeb) PC 12.1 18.0 TW 14.5 21.6’ PC 16.8 25.2’ PC 17.7 25.7 TW 24.4(Feb) 25.4(Aug) PC 20.5 25.4’ PC

25.7 26.5” 24.2 26.2 12.8 17.9

F R

C F F R

28.1(Aug) 27.9(Feb) 23.7 27.6

15.3 20.4

F R

16.1 20.2

10.8 20.2(Aug) 24.5CAug) 22.7 23.3 26.2

5.3 20.7 23.1 14.5 25.6

15.5 24.6(Feb) 28.4(Feb) 28.9 28.9 28.0

7.3 25.5 25.4 24.0 27.2

8.6 23.3

10.0 25.5

23.7 26.9

20.9 10.0 26.4

3.9 6.0

26.6 15.1 26.7” 11.0 12.0 9.5 8.4

28.0

3.2 2.1

26.1 28.9 19.8 24.4(Aug) 27.8CAug) 13.2 14.6 16.0 15.0 24.8CFeb) 17.3 10.3

30.0 28.4 28.8(Feb) 27.9(Feb) 19.8 22.6 24.6 24.1 24.5(Aug) 23.5 15.0

Nofe. Aug = August; Feb = February; C = coccoliths; F = foraminifera: R = radiolaria: PC = piston core; TW = trigger- weight core.

’ Northern Hemisphere cores that follow the seasonal pattern of the Southern Hemisphere. For these cores T, = August and T, = February.

b Stage 5e SST estimates are from the nearest sample to the stage 5e level. If there are temperatures au equal distance above and below the stage Se level, then the oldest temperature is listed. For all other cores the SST estimates listed are from the stage 5e level.

’ The core-top SST estimates are from a level within the mixed layer (2-5 cm). All other core-top SST estimates are from the top (O-l cm).

* V19-29 core-top SST estimates were generated using transfer function RPE while downcore SST estimates were generated using transfer function RP7.

132


inal record. With the later publications of Peng et al. (1977), Duplessy (1978), and Hutson (1980), it became clear that bioturbation is a far more complicated factor than initially realized. These papers demonstrated that major oxygen-isotope transitions can be not only smoothed but also shifted upward or downward from their original level in the cores. This occurs because of variations in the absolute abundance of the isotopic signal carrier, usually a single species of foraminifera. Normal bioturbational mixing sends unequal absolute amounts of the signal carrier upward and downward in the core. This causes the isotopic signal to cascade from regions where the signal carrier is abundant to regions where it is scarce (Hutson, 1980).

Ruddiman et al. (1980a) subsequently demonstrated that biotic curves are subject to similar translational offsets due to bioturbation. Again, the problem was traced to absolute abundance changes: any paleoclimatic signal based on a sedimentary component not referenced to total sediment weight is susceptible to translational offsets and other complications.

The implications for this project are critical because one of our major goals is to define lead-lag relationships between the global 6i*O and local SST signals. These complications mean that the lead-lag relationship now observed in a given core may be an artifact of mixing.

To gain some control over this problem, in the last year of the project we gathered as much absolute abundance data as possible. We attempted to obtain per-gram abundance counts for (1) the species of foraminifera used for isotopic analysis and (2) the biotic groups on which the SST estimates were based. All such data obtained are plotted in Figures 3 through 54 and listed in Appendixes I and 3. Abundance variations in the isotopic signal carrier are plotted on the left next to the isotopic signal; variations in the SST signal carrier are plotted on the right next to the SST curve. CaCO, values for several cores with

coccolith SST estimates are plotted as rough abundance indicators in place of actual per-gram coccolith counts (Appendix 4).

RECONSTRUCTION OF THE LAST INTERGLACIAL OCEAN

This section summarizes our comparison between the modern ocean and the last interglacial ocean at the ice-volume minimum marked by oxygen isotopic Substage 5e.

Methods

Two methodological choices are necessary for this comparison: first, a protocol for defining the exact isotopic 5e level, and, second, a decision for selecting modern SST (atlas values or core-top estimates).

Choice of oxygen isotopic 5e level. Be- cause we have sampled these cores in far greater detail than is customary in deep-sea studies, the sharply developed peak often characteristic of 6180 Substage 5e has in many cores expanded into a broader, plateau-like feature. On such a feature, the choice of the isotopically lightest analysis is not as easy or as meaningful as on <he more peaked curves. In some cases, the isotopically lightest al80 analysis may just reflect random effects from the _tO. 12%00 analytical reproducibility. We thus chose not to base our 5e selections on any single analysis.

Instead, we passed a three-point running average over the raw data and selected the lightest value from among these averages. Use of this kind of filter suppresses the effects of one-point peaks without unduly at- tenuating actual structure in the isotopic curves. Our choices of the isotopic Se level in each core are listed in Table 4 and marked by a star on each downcore plot (Figs. 3-54). Summer and winter SST estimates at the selected levels in each core are listed in Table 4.

In nine cores (K7081, Fig. 7; RC8-39, Fig. 10; RCll-210, Fig. 15; RCll-230, Fig. 16; RC12-339, Fig. 18; V19-29, Fig. 29; V22-182, Fig. 35; V28-304, Fig. 45; and V34-88, Fig. 52), the three-point running


PLANKTONIC S”O

-1.5 -2.0 SST Tw and Ts

320

i

t

280

FIG. 3. Record of the last interglaciation in equatorial Atlantic core Alt?O-73. Data from Gardner and Hays (1976) and Emiliani (1955). Sample at 375 cm is a mixture of individuals from 370 and 380 cm. Filled diamonds contain mostly G. sacculifer but also G. ruber. For this figure, and successive figures through 54, the filled star marks the last interglacial level. A, and A, are atlas picks of winter and summer SST at this site. A is used without subscripts where the seasonal scales are separate. CT, and CT, are estimated winter and summer SST for the core-top sample, with superscripts F, R, and C designating foraminifera, radiolaria, and coccoliths, respectively, in cores with multiple biotic groups. CT is used without subscripts where the seasonal scales are separate.

average still produced a broad and some- times noisy plateau instead of a well-defined isotopic minimum. For these cores, we chose the Se level in approximately the middle of the plateau produced by the smoothing.

Choice of modern SST value. There are two possible ways to select values representative of modern SST at each core site. Because each option has advantages and drawbacks, we have used both as standards for comparison to the estimated temperatures in the last interglacial ocean.

The most direct comparison is with the atlas values of SST used by the CLIMAP Project for calibrating transfer-function equations during the last decade (Imbrie

and Kipp, 1971; Cline and Hays, 1976; CLIMAP, 1981). Atlas values have the ad- vantage of being based on actual measurements taken over a known time interval (the last 50 yr). These values are listed in Table 4 and plotted in Figures 3-54 along the SST scale. (A, is summer atlas temperature; A, is winter atlas temperature; A is used without subscripts where the seasonal scales are separate.)

Two potential problems may occur with the use of atlas values of SST. First, the last 50 yr represent only 1% of the 5000-yr duration of the modern ice-volume minimum. Modern (atlas) SST thus may not be representative of the last 5000 yr; evidence of significant climatic variations during this


BENTHIC 8’“O

4.5 4.0

q i

i.

SST Tw and Ts

35 6 10 Aw ,4 18 A

122 a .

b 200

* 210

1 220

230

240

b 250

. 260

, 270

. 280

FIG. 4. Record of the last interglaciation in North Atlantic core D117. See Figure 3 legend for additional explanation of symbols.

PLANKTONIC 8”0 SST T w and Ts

3.5 3.0 2.5 2.0

FIG. 5. Record of the last interglaciation in Indian/Antarctic core E49-18. Part of data from Hays et al. (1976). See Figure 3 legend for additional explanation of symbols.

L - T 0 5000 10,oofJ

PLANK FORAMlgm SED PLANK 8”0sp/gm SED

FIG. 6. Record of the last interglaciation in Norwegian Sea core K-l 1. Data from Duplessy et al. (1975) and Kellogg et al. (1978). See Figure 3 legend for additional explanation of symbols.

BENTHIC S’*O SST T w SST Ts

(0)

q Mekm,r IO 0 U”!ger,n‘7 I PLANK FORAMlgm SED BENTHIC d% sp/gm SI

FIG. 7. Record of the last interglaciation in North Atlantic core K708-1. Data in part from Ruddiman et al. (1977, 1980b). See Figure 3 legend for additional explanation of symbols.

136

DEPTH BENTHIC 8% SST T w and Ts

CM 5.0 4.5 4.0 3.5 10 10 12 14 16 18 A, 20 hs Z? c

1 l

7.20

740

760

780

800

820

840

860

880

900

920

940


720

740

760

780

800

820

840

860

SSO

900

920

940

FIG. 8. Record of the last interglaciation in North Atlantic core M12392-1. Data from Thiede (1977) and Shackleton (1977). See Figure 3 legend for additional explanation of symbols.

PLANKTONIC s’oo BENTHIC 8% SST Tw SST Ts DEPTH 50 40 30 4

CT J.

6 A + 9.5 CM 1 L . .

40 30 20

FIG. 9. Record of the last interglaciation in Indian/Antarctic core MD73025. Isotopic data from Duplessy (1978); Nonion is synonymous with Melonis. See Figure 3 legend for additional explanation of symbols.


FIG. 10. Record of the last interglaciation in subantarctic core RC8-39. Foraminiferal temperatures were obtained using a calendar equation: T, is synonymous with August and T, is synonymous with February. See Figure 3 legend for additional explanation of symbols.

PLANKTONIC de0 SST T w SST Ts

FIG. 11. Record of the last interglaciation in North Atlantic core RC8-145. See Figure 3 legend for additional explanation of symbols.

interval exists at both decadal and millen- values are left after subtracting the ob- nial time scales (Mitchell, 1961; Lamb, served (atlas) temperatures from the esti- 1972; Hays et al., 1976; Thiede, 1977). mated temperatures derived from transfer-

Second, in some cores large residual function analyses of the core-top fauna or


PLANKTONIC 8”O SST T w SST TS

139

t 300

t

310

RADlgm SED

FIG. 12. Record of the last interglaciation in equatorial Pacific core RC 10-65. Data in part from Romine and Moore (1981). See Figure 3 legend for additional explanation of symbols.

flora. One published example is the map of residual values in Figure 28 of Kipp (1976). The mismatches between estimated and observed atlas temperatures in her plot are geographically scattered but may reach values larger than 2°C in some cores. This discrepancy raises a difficult question: if the transfer-function analysis of the core- top sample misses the atlas value by 2”C, how should we view an equal 2°C offset of the last interglacial estimate from the atlas values? In such a case, the last interglacial estimate would match the core-top estimate and imply no real temperature difference between the last interglaciation and today.

Because of this problem, we decided to use the core-top estimates (where available) as a second line of comparison, with closest attention to cores in which the core- top estimates are significantly offset from atlas temperatures. Available core-top estimates are listed in Table 4 and plotted in Figures 3-54 along the SST scale. (CT, is estimated summer SST at the core top; CT, is estimated winter SST. Where season is indicated on the scale, the subscript s or w

is omitted. For scores with multiple biotic estimates, the superscripts R, F, and C indicate radiolarian, foraminiferal, and coccolith core-top estimates.) Finally, we have used published data on detailed changes in SST during the Holocene. Several areas display significant changes in Holocene SST: for example, the subantarctic (Hays et al., 1976), the eastern South Atlantic (Embley and Morley, 1980), and the Can- aries Current (Thiede, 1977). For these few areas, we have widened the scope of the comparison with the last interglacial ocean, using not just the atlas and core-top values but the trends of the last 5000 yr as well (see section on Other Evidence of 6t80/SST Phasing: Evidence from Other Transitions).

Results

The results discussed in this section are based only on direct observations made from the plotted data (Figs. 3-54); in the section Problems and Complications, we will discuss the impact of complicating factors in altering these direct observations. Estimated February and August SST at the


DEPT

H

PLAN

KTO

NIC

PO

3.5

3.0

2.5

2.0

-

SST

Tw

an

d Ts

1 380

. 40

0

2 +

410

m

r .

420

2

CM

380

’

390

*

400

<

410

q

420

4

430

’

440

*

450

<

460

4 b 460

470

’ ’ 470

480

4 ’

480

00

55 7’0

’ 85

PLAN

K S’

*Osp

/gm

SE

D.

* .

RAD

/gm

SE

D 50

00

15,00

0 25

,000

%

coco

s

PLAN

K.

FORA

M/

gmSE

D

FIG.

14

. R

ecor

d of

the

last

inte

rgla

ciatio

n in

sub

anta

rctic

co

re R

CI

I-120

. Pa

rt of

dat

a pr

evio

usly

in

Hay

s et

aI.

(197

6).

Fora

min

ifera

l te

mpe

ratu

res

were

ob

tain

ed

usin

g a

cale

ndar

eq

uatio

n:

T,

is s

ynon

ymou

s wi

th

Augu

st

and

T, i

s sy

nony

mou

s wi

th

Febr

uary

. Se

e Fi

gure

3

lege

nd

for

addi

tiona

l ex

plan

atio

n of

sym

bols

.


DEPTH CM

PLANKTONIC 8”O SST T w SST Ts

0.0 -0.5 -1.0

22 .

0 SO 160

PLANK 8 D sp /gm SED

FIG. 15. Record of last interglaciation in equatorial Pacific core RCl l-210. Data in part from Romine and Moore (1981). See Figure 3 legend for additional explanation of symbols.

DEPTH CM

PLANKTONIC S”O

FIG. 16. Record of the last interglaciation in Pacific Ocean core RCl I-230. Data in part from Romine and Moore (1981). See Figure 3 legend for additional explanation of symbols.

DEPT

H

PLAN

KTO

NIC

S”O

BE

NTHI

C E’

b SS

T T

w

SST

Ts

CTC

CTF

A 5.

0 4.

5 4.

0 3.

5 3.

0 12

14

16

18

20

22

C

TrC

Tca

A

J %

0

20

40

Uvig

/g

m

SE0

SST

0 25

00

5oco

0

5 ID

C

wve

N/g

m

SED

0 G

,“,,O

,O

0 Cd

ru

s/lcr

sk7r

rh

0 TW

FO

RIM

c,

T

w C

OCCD

PL

ANK

FOR

AMS/

gm

SED

40

60

-To

PLAN

K 8”

0sp/

gm

SED

0 G

,dlg

rn

0 cd

r*

Nm

0 Tr

Fc

JRm

.4

a 7s

co

cco

a “w

garm

o .

*/gm

FO

RAM

%

tac

o,

m L

lwg

/gm

n

% c

oca,

FIG

. 17

. R

ecor

d of

the

las

t in

terg

lacia

tion

in S

outh

At

lant

ic

core

R

C12

-294

. Se

e Fi

gure

3

lege

nd

for

addi

tiona

l ex

plan

atio

n of

sym

bols

.

DEPT

H CM

I‘ PL

K 0 8’

%sp

/gm

90

0 . SE

D 1eo

o - ‘0

4

8 BE

NTHI

C 8”

0sp,

gm

SED

l&C

PL

ANK

FDR

AM/g

m

SED.

0 un

g* /

qm

. o,k

?/p

J FI

G.

18.

Rec

ord

of t

he l

ast

inte

rgla

ciatio

n in

no

rther

n In

dian

O

cean

co

re

RC

l2-3

39.

Tem

pera

ture

s we

re

obta

ined

us

ng

a ca

lend

ar

equa

tion:

T,

is

sy

nony

mou

s wi

th

Augu

st

and

T,

is s

ynon

ymou

s wi

th

Febr

uary

. Se

e Fi

gure

3

lege

nd

for

addi

tiona

l ex

plan

atio

n of

sym

bols

,


I I

I I

Rods

/gm

SED

FIG

. 20.

R

ecor

d of

the

last

int

ergl

acia

tion

in S

outh

At

lant

ic

core

R

C13

-228

. Se

e Fi

gure

3

lege

nd

for

addi

tiona

l ex

plan

atio

n of

sym

bols

.

,b--

2r

,ff

a,

a --

, -=

a 4’

. 40

00

12.0

00

20.0

’00

PLK

FO!?

AM/g

m

SED

RAD

/gm

SE

D

FIG

. 21

. R

ecor

d of

the

las

t in

terg

lacia

tion

in S

outh

At

lant

ic

core

R

C13

-229

. D

ata

from

Em

bley

an

d M

orle

y (1

980)

an

d M

orle

y an

d H

ays

(198

1).

See

Figu

re

3 le

gend

fo

r ad

ditio

nal

expl

anat

ion

of s

ymbo

ls.


DEPTH BENTHIC S”O SST T w SST Ts

5.0 4 0 3.0 8 10 12 14 A-w5 5

SST

0 TW RAD 8 TS Am 0 17,500 35,000 a%l? RAD

RADlgm SED

FIG. 22. Record of the last interglaciation in South Pacific core RC15-61. See Figure 3 legend for additional explanation of symbols.

FlANI(ToNIC s ‘-0 SST T w SST Tr DEPTH ,5

CM L

230 I

FIG. 23. Record of the last interglaciation in southern Indian Ocean core RC 17-69. Temperatures were obtained using a calendar equation: T, is synonymous with August and T, is synonymous with February. See Figure 3 legend for additional explanation of symbols.



DEPTH 0.0 -0.5 -1.0 -!.5 2.0 26 27 28 CM

1 29

/////// TURBIDITE

////// /

PLANK. B’bsp/gm SED. PLANK FORAM/gm SED

FIG. 24. Record of the last interglaciation in southern Indian Ocean core RC 17-98. Temperatures were obtained using a calendar equation: T, is synonymous with August and T, is synonymous with February. See Figure 3 legend for additional explanation of symbols.

FIG. 25. Record of the last interglaciation in Gulf of Mexico core TR126-23. See Figure 3 legend for additional explanation of symbols.


FIG. 26. Record of the last interglaciation in Gulf of Mexico core TR126-29. See Figure 3 legend for additional explanation of symbols.

PLANKTONIC 8”O SST T w SST Ts 0.0 -0.5 - 1.0 -1.5 - 2.0 25 26 A-328 0

DEPTH A . . 22 m-27 4

CM

180

260

320

22 23 24 25 A-926 2 . CT-+25 9

180

220

240

260

320

FIG. 27. Record of the last interglaciation in Caribbean core V12-122. Data from Broecker and Van Donk (1970) and Imbrie et al. (1973). See Figure 3 legend for additional explanation of symbols.

DEPTH CM

380

400

420

440

460

480

500

520

540


35 3.0 2.5 6 I ! 8 CT 10


35 3.0 2.5 6 n 8 CT 10 .

4 n 6 CT 8

0 1$000 30.000

RAD /gin SED

FIG. 28. Record of the last interglaciation in subantarctic core V18-68. See Figure 3 legend for additional explanation of symbols.

BENTHIC I)“0 SST T w SST Ts

to 4.5 4.0 3.5 3.0 20 A DEPTH 22 24 “Q6 28

CM 14 16 18 q20 0 22 24

FIG. 29. Record of the last interglaciation in Pacific core V19-29. Data from Shackleton (1977) and Romine and Moore (1981). See Figure 3 legend for additional explanation of symbols.

151



DEPTH 0.7 0.2 -0.3 25.4tP.29 30 * 28.3CCT l .

CM 23. , t t? 26 27

23.9tCT .

200 l . 200

205 < ’ 205

210 l ’ 210

215 4

235 ’ ’ 235

240 l b 240

it PLANK. 8”Osp /gm SED I;5 3;o E lDOopLAN;:ORI\M,gfiD

FIG. 30. Record of the last interglaciation in South Pacific core V19-53. Data in part from Romine and Moore (1981). See Figure 3 legend for additional explanation of symbols.

DEPTH CM 300 4

310 4

320 4

330.

340 l

350 4

360 *

370 ’

380 4

390 *

400 ’

410 ’

420 l

BENTHIC 8”O

5.5 5.0 4.5 4D 3.5 3.0 SST Ts

29 24J A 300 .

l 320

. 330

b 340

,354

t 360

I 400

410

20,000 ao,&o L 420

20 30 40

RAD/gm SE0 o/o caco3

FIG. 3 1. Record of the last interglaciation in North Pacific core V2 I- 146. Data in part from Thompson (1981). See Figure 3 legend for additional explanation of symbols.

DEPT

H

PLAN

KTON

IC

6’80

BENT

HIC

6’80

SST

Tv

SST

Ts

5.0

4.5

4.0

3.5

“‘L

A 3.0

28

-0.5

-1.0

-1.5

-2.0

23

ZZ.l-C

TC

1

PL

0 40

00

BOO0

PLAN

K FO

RAM/

gm

SED

20

35

5c

%

taco,

FIG

. 32

. R

ecor

d of

the

las

t in

terg

lacia

tion

in S

outh

At

lant

ic

core

V2

2-38

. Se

e Fi

gure

3

lege

nd

for

addi

tiona

l ex

plan

atio

n of

sym

bols

.


FIG. 33. Record of the last interglaciation in subantarctic core V22-108. See Figure 3 legend for

PLANKTONIC 8”O SST T w SST Ts n CT

DEPTH 3.0 2.5 2.0 8 10 12 14 16

CM 6 8CTA 10 12

460 l

additional explanation of symbols.

last interglacial level are shown in Figures 55a and b. Where two or more biotic groups were analyzed, each estimate is annotated by the appropriate code: F for foraminifera, R for radiolaria, and C for coccoliths. The dashed contours of SST shown in Figures 55a and b are modern atlas values.

A comparison of the plotted temperatures against the atlas contours shows that the last interglacial ocean was in general very similar in temperature to today’s ocean. To facilitate the comparison, we plotted the temperature differences between the two time slices in Figures 56a and b and listed all temperature values in Table 4. These data show that almost 60% of the SST estimates in the last interglacial ocean differ from today’s atlas values by amounts less than the typical f. 1 .O”- 1.5”C standard error of estimate.

Many of the apparently large differences visible in Figures 56a and b do not hold if the Stage 5e SST estimates are compared against core-top estimates rather than atlas values. For example, temperature estimates significantly cooler than modern values occur in several cores from the Southern Ocean, particularly the four cores in the Indian Ocean sector (E49-18, Fig. 5; MD73025, Fig. 9; RC8-39, Fig. 10; RCll- 120, Fig. 14). Radiolarian winter SST estimates at the isotopic Stage 5e level average 2°C cooler than atlas values today; summer values average more than 3°C cooler. The subantarctic transfer function based on radiolaria, however, yields core-top estimates for several of these sites that are also considerably cooler than the atlas values (Table 4). Underestimation of core-top temperatures in this area, particularly in summer,

~s[oqml(s

30 uoym

e~dxa

. V

-- ------a

- --

I) . I

q cp \

/ a

b

Hld30

Sl PUD

Ml

ISS O

S& 9

3lHlN38 O

S,9 3lNOlM

NVld

156 CLtMAP PROJECT MEMBERS

IEPTH


C,dwu&Mpomp/gm SED

FIG. 36. Record of the last interglaciation in Atlantic core V22-196. CaCO, data from Hays and Peruzza (1972). See Figure 3 legend for additional explanation of symbols.

157

FIG. 37. Record of the last interglaciation in North Atlantic core V23-82. SST estimates from Sancetta et a/. (1973) and isotopic data from Ruddiman and McIntyre (1979). See Figure 3 legend for additional explanation of symbols.


DEPTH CM

PLANKTONIC S”O EENTHIC S”O SST T w SST Ts 4 5.0 4.5 4.0 3.5 27y 29

0.5 0.0 -0.5 -1.0 -1.5 PO 22 24 I “kf A-+26.4

” _I ”

BENTHIC s’“O sp/gm SED

FIG. 38. Record of the last interglaciation in equatorial Atlantic core V25-59. See Figure 3 legend for additional explanation of symbols.

FIG. 39. Record of the last interglaciation in North Atlantic core V27-20. See Figure 3 legend for additional explanation of symbols.

THE LAST INTERGLACIAL OCEAN 1.59

FIG. 40. Record of the last interglaciation in Norwegian Sea core V27-86. Data from Streeter et al. (1982). See Figure 3 legend for additional explanation of symbols.

FIG. 41. Record of the last interglaciation in Norwegian Sea core V28-14. See Figure 3 legend for additional explanation of symbols.


FIG. 42. Record of the last interglaciation in Norwegian Sea core V28-56. Data in part from Kellogg et al. (1978). See Figure 3 legend for additional explanation of symbols.

! FIG. 43. Record of the last interglaciation in Caribbean core V28-127. See Figure 3 legend for addi-

tional explanation of symbols.

161 THE LAST INTERGLACIAL OCEAN

BENTHIC 8 “0 SST T w PLANKTONIC 8 “0 50 45 40 35 30 SSTTs 28 CT 4

DEPTH vg

CM -05 . -10 -15 -20 23 25 27 * AL9CT-M6

220

230

240

PX).

240.

FIG. 44. Record of the last interglaciation in equatorial Pacific core V28-238. Planktonic isotopic data from Shackleton and Opdyke (1973). See Figure 3 legend for additional explanation of symbols.

l8

%

2, 16

coca,

FIG. 45. Record of the last interglaciation in North Pacific core V28-304. Data in part from Thompson (1981). See Figure 3 legend for additional explanation of symbols.



470. l 470

0 1000 2000 0 m,ooo 20,000

PLANK. Porp/gmSED PLANK. FORAM,gm SE0

FIG. 47. Record of the last interglaciation in northern Indian Ocean core V29-29. Temperatures were obtained using a calendar equation: T, is synonymous with August and T, is synonymous with February. See Figure 3 legend for additional explanation of symbols.

,EPTH CM 460

473 I

b

FIG. 48. Record of the last interglaciation in North Atlantic core V29-179. Isotopic data from Streeter and Shackleton (1979) and SST estimates from Ruddiman and McIntyre (1979). See Figure 3 legend for additional explanation of symbols.

q f

7 PO

SS

T

PLK

8”O

r,Vgm

SE

D ’

6 12

BENT

HlC

81’0

sp/gm

SE

D

FIG

. 49

. R

ecor

d of

the

last

int

ergl

acia

tion

in N

orth

At

lant

ic

core

V3

0-97

. D

ata

in p

art

from

Ru

ddim

an

and

McI

ntyr

e (1

981a

). Ar

rows

on

be

ntho

nic

fora

min

ifera

l ab

unda

nce

curv

es i

ndica

te

min

imum

(3

) va

lues

. Se

e Fi

gure

3

lege

nd

for

addi

tiona

l ex

plan

atio

n of

sym

bols

.

DEPTH CM

0 ii0 300 PLPNK a”0 sp /gm SED


PLANKTONIC 6’*0 SST T w SST Ts

2.0 1.5 23 psp 27

10 12 14 16J*5’ 18 .

I

0 ,000 2000 PLANK. FORAM/gm SE0 % taco,

165

FIG. 50. Record of the last interglaciation in North Pacific core V32-126. Data in part from Thompson (1981). See Figure 3 legend for additional explanation of symbols.

BENTHIC 8”O SST T w SST Ts

“i” t5 4.0 3.5 18 22 A DEPTH J cT26

A CM 6 10 14 78 . *40 ,

250

260

:

PLANK. FORAM /gm SED % coca,

FIG. 5 I. Record of the last interglaciation in North Pacific core V32-128. Data in part from Thompson (1981). See Figure 3 legend for additional explanation of symbols.

PLAN

KTO

NIC

6’*0

BENT

HIC

6’*0

SST

Tw

an

d Ts

DEPT

H

CM

5.0

4.5

1.0

0.5

0.0

-0.5

-1.0

1.5

. .

0

4.0

3.5

3.0

L 5

CL

4 CT

* 2%

23

42

5 27

L

. -

- -

- -

a 2

15.7

0

12.50

0 25

.000

PLAN

K.

FDRA

M/Q

n SE

D.

6o

7-o

%

taco

,

FIG

. 52.

Rec

ord

of th

e la

st in

terg

lacia

tion

in n

orth

ern

Indi

an

Oce

an c

ore

V34-

88.

Tem

pera

ture

s we

re

obta

ined

us

ing

a ca

lend

ar

equa

tion:

T,

is

syn

onym

ous

with

Fe

brua

ry

and

T, i

s sy

nony

mou

s wi

th

Augu

st.

See

Figu

re

3 le

gend

fo

r ad

ditio

nal

expl

anat

ion

of s

ymbo

ls.


BENTHIC L”O SST Tw SST Ts

A 5.0 4.5 4.0 3.5 18 20 22 24 cT 26

SST d TW RAO RAD/gm SED

FIG. 53. Record of the last interglaciation in South Pacific core Y71-6-12. See Figure 3 legend for additional explanation of symbols.

FIG. 54. Record of the last interglaciation in North Pacific core Y72-1 l-l. Data in part from Heusser and Shackleton (1979). See Figure 3 legend for additional explanation of symbols.

CLIMAP PROJECT MEMBERS

FIG. 5.5. (a) Estimated February SST temperature at the last interglacial level. Contours are modem SST values. (b) Estimated August SST temperature at the last interglacial. Contours are modem SST values.


is a general characteristic of transfer-function equation RAN3 (Table 3). If we compare the Stage 5e temperature estimates against the core-top estimates, most of the negative temperature anomalies fall to values of less than 1°C or even become positive anomalies (Table 4). In addition, estimates from foraminifera in cores RCS-39 (Fig. 10) and RCll-120 (Fig. 14) point to temperatures the same as, or even warmer than, those today. These major discrepan- cies in radiolarian and foraminiferal estimates are discussed in the section on Prob- lems and Complications, Discordant Mul- tiple Estimates.

We conclude that the last interglacial subantarctic ocean may have been slightly cooler than it is today, but the measured differences are well within the error of estimate.

For the remaining suite of cores from the last interglaciation, comparison of the Stage 5e SST estimates against the core-top estimates reduces the larger anomalies to <I 5°C in many cores, even reversing the sign of the anomaly in some. On the other hand, the core-top comparison creates new anomalies >lS”C in a significant number of other cores.

Of these remaining anomalies, those that are larger only by comparison to core-top estimates or to atlas values, but not to both, are suspect. Anomalies that survive both comparisons, but that disagree with those from other biotic groups, or those in nearby cores, are also suspect. In the latter case, of course, small-scale local anomalies are one plausible interpretation.

With these qualifications in mind, three areas still show Stage 5e temperature estimates with sufficiently consistent diver- gence from those of today to warrant additional discussion: (1) the mid-latitude North Atlantic, 35” to 65” N; (2) the western equatorial Atlantic, Caribbean, and Gulf of Mexico; and (3) the eastern boundary currents of the South and equatorial Atlantic. We have focused attention on these areas

because temperature anomalies recur in more than single isolated cores.

In the mid- to high-latitude North At- lantic, Stage 5e February SST values were often 1.0” to I.S”C warmer than modern atlas values (Fig. 56a). For a number of cores, the last interglacial temperatures are also warmer relative to the core-top estimates (e.g., K708-1, Fig. 7; winter SST in V27-20, Fig. 39; and V29-179, Fig. 48). Futhermore, SST estimates at levels adjacent to the 5e picks in these cores were generally warmer still and would create even larger anomalies. On the other hand, some cores show no significant temperature differences (V23-82, Fig. 37; V30-97, Fig. 49). We conclude that there is some evidence in this area of a lo-2°C increase in February SST during the last interglaciation relative to today, but that the results are not entirely consistent. We are particularly wary of the Stage 5e picks and hence the SST anomalies in three Norwegian Sea cores (K-11, Fig. 6; V27-86, Fig. 40; V28- 56, Fig. 42) because of inadequate (or no) benthonic foraminiferal 6i80 records and the likelihood of meltwater complications in the planktonic foraminiferal 6*80 record (Kellogg et al., 1978).

The second area of divergent estimates is the western equatorial Atlantic, Caribbean, and Gulf of Mexico. In this region, Stage 5e winter temperature estimates in five cores (TR126-23, Fig. 25; TR126-29, Fig. 26; V12-122, Fig. 27; V25-59, Fig. 38; and V28-127, Fig. 43) were 1.6”-2.8”C colder than that today. In view of the relatively small interglacial/glacial SST ranges in these areas, these 2°C anomalies represent considerable departures. Three out of five of these anomalies hold up when the Stage Se estimates are compared to core-top values (Table 4). Complicating factors will be discussed in the section on Problems and Complications, Bioturbational and Bottom- Current Mixing.

The third area of large anomalies includes the Atlantic eastern boundary cur-


rents, including three cores in the South At- lantic (RCll-86, Fig. 13; RCl3-228, Fig. 20; and RC13-229, Fig. 21) and one in the North Atlantic (V22-196, Fig. 36). In these cores, estimated temperatures at the last interglacial level depart from modern values by as much as 45°C but in both positive and negative directions. The integrity of these estimates is compromised by other factors to be discussed in the section on Problems and Complications, Validity of the SST Estimates.

Scattered cores in other areas show temperature anomalies lying outside the error of estimate. We view these singular occurrences with caution. The complications to be discussed in the section on Problems and Complications could well be responsible for these isolated divergent estimates. In none of the other areas do we have the clustering of cores needed to make a convincing case for systematic SST anomalies in the last interglacial ocean.

LEAD-LAG RELATIONSHIPS: SST AND ICE

This section summarizes the timing of local SST variations relative to changes in global ice volume across two intervals: the glacial-to-interglacial ice-volume decrease across the isotopic Stage 6/5 boundary and the interglacial-to-glacial ice-volume increase across the isotopic Substage Se/Sd boundary.

Methods

The rationale for determining leads and lags is shown in Figure 57. Most areas of the world ocean show SST changes that fluctuate with 6180/ice-volume variations; that is, periods of high 8’*0 values (large ice volume) roughly equate with periods of cold SST and conversely. This is particularly true for the regions with high-amplitude SST variations such as the mid-latitude and high-latitude subpolar oceans, the eastern boundary currents, and the equatorial divergences. In most of these regions, however, there are offsets in timing

between the local SST and global @*O/ice- volume curves. These offsets are the subject of this section. We have used two techniques for determining lead-lag relationships in the region of oxygen isotopic Substage 5e: offsets in transition midpoints and cross-correlation analysis.

Offsets in transition midpoints. The most direct method of determining leads and lags was to pick the amplitude midpoints of a given isotopic and SST transition (Fig. 57), measure the relative offset of the midpoints in cm of core length, determine the mean sedimentation rate across the 5e interval, and convert the cm offset to a time offset in years. This had to be done separately for the glacial-to-interglacial transition between isotopic Stages 6 and 5 and for the interglacial-to-glacial transition between isotopic Substages 5e and 5d. This treatment provided discrete views of the different kinds of leads and lags on ice-decay and ice-growth transitions.

Depth picks for the transition midpoints of the summer and winter SST curves are listed with the isotopic transition midpoints in Table 5 (the Stage 6/5 #*O transition) and Table 6 (the Substage 5e/Sd ai80 transition). The benthonic foraminiferal isotopic curves were used in preference to planktonic foraminiferal curves in cores where both were available.

Not all 6’*0 or SST curves rise and fall so smoothly as those portrayed schematically in Figure 57. For cores in which the midpoint value of 6180 or SST of a transition is attained at more than one depth level (whether due to analytical noise or actual structure on the curve), each individual crossing of the midpoint depth level in the appropriate direction is recorded in Tables 5 and 6, as well as the average of all these values. In three cores (V19-29, V30-97, Y72-1 l-l), we ignored a one-point oscilla- tion in the 6isO curves as mixing artifacts.

Also listed in Tables 5 and 6 are the cm offsets between the SST and 6’*0 midpoints. Positive values indicate that the SST curve changes earlier than the 6180


FIG. 56. (a) Difference in “C of estimated February SST between last interglacial and today (value at oxygen isotopic Substage 5e minus value taken from modem atlas data). (b) Difference in “C of estimated August SST between last interglacial and today (value at oxygen isotopic Substage 5e minus value taken from modem atlas data).


GLOBAL ICE VOLUME

MAX-MIN

‘C SUMMER SST S’*O vs PO E! “C SUMMER SST

5 10 15 5 4 3 5 10 15 L

/

DEPTI IN

CORE

FIG. 57. Schematic representation of the meaning of leads and lags. The St80 record is used as a standard of reference. “Ocean leads” means that the SST signal moves toward a given climatic state before the 6t80/ice-volume signal responds. The interglacial state is defined by convention as minimum ice volume and maximum temperature, and glacial climatic state as maximum ice volume and minimum temperature.

curve, and conversely. For cores with multiple depth picks for either the 6t80 or SST midpoint, average depth values for all individual crossings of the midpoint amplitudes were used to determine the relative offsets. The average values are boldface in Tables 5 and 6.

No picks of SST transition midpoints and/or of 6i80 transition midpoints were made for a number of cores in Tables 5 and 6 for a variety of reasons. Some cores showed no significant SST change in summer or winter, particularly cores in central subtropical gyres (e.g., RC8-145, Fig. 11). Equatorial cores often have substantial winter season changes but minimal summer SST variations (e.g., V25-59, Fig. 38).

We did, however, pick midpoints for some cores in which the amplitudes of the SST curves lie within the 2” to 3°C two-way range of the normal error of estimate (Table 3). These were made if the estimates tracked in a sufficiently smooth way to allow the basic cool-warm-cool trend to stand out clearly despite the low amplitude (e.g., core TR126-23, Fig. 25).

Other cores contain SST signals that do not show the typical cold-warm-cool sequence used as a standard for the analysis

(e.g., core RC17-98, Fig. 24). This may indicate that some low-latitude areas simply do not follow the cold = glacial and warm = interglacial theme of the rest of the world ocean. Large regions of the central subtropical gyres of the Pacific Ocean were slightly warmer at the last glacial maximum, 18,000 yr ago, than today (CLIMAP, 1976; 1981); this could suggest an antiphase relationship between ice volume and SST in some areas. Another possibility could be complications from high-frequency SST changes of various origins (see section on Other Evidence of 6r80/SST Phasing, Com- parison with Long-Term Phase Relation- ships).

All cores with these problems were considered indeterminant and not included in this analysis (blank spaces in Tables 5 and 6). Many were, however, subjected to cross-correlation analysis (next section). Ultimately, the choices were based on sub- jective judgements; Figures 3-54 and Ap- pendixes l-4 provide the necessary raw data for the critical reader to evaluate whether or not the right cores were included or rejected in the lead-lag analysis.

The next step was to determine the sedimentation rates to be used to convert

173 THE LAST INTERGLACIAL OCEAN

TABLE 5. RELATIVE OFFSET BETWEEN MIDPOINTS OF STAGE 6/5 &I80 TRANSITION AND ASSOCIATED DEGLACIAL SST WARMING

Core SST,

Midpoint picks (cm/depth of core)

SST, 6’80

Relative offsets (in cm)b

SSTJ8’80 SST,/6’80

A180-73 357 D117 270 E49-18 497

K-11 K708-1 M12392-1 MD73025

RC8-39C

833 1427

(1414,144O) 84oF

(836,843) 882 R

RC8-145 RClO-65

RCl l-86 RCll-120’

RCll-210

RCl I-230

RC12-294

296 305 (291,301) (302,307)

439 F 440 F 449 R 449R

RC 12-339’

RC13-205

3Ooc 3oOc (294,298, (293,296, 300,308) 301,308)

301 F 301 F 348 351

(342,349,353) (348,353) 230 F 230 F

RC13-228

236 R 236 R (232,236,241) (232.236,241)

667 C 676 C (657,676) (661,677,691)

664R 679 R RC13-229 RCIS-61

RC 17-69’ 242d

RC 17-98’ TR126-23 TR126-29 v12-122 V18-68 V19-29

600 607 830 823

505d 845

v19-53 V21-146

V22-38 243 F

270 497

358 268 511

(489,532)

838 1405

852 F (834,841,880)

882R

844 1377

(1370,1383) 863

206 290

(287,292)

439

191 (187,195)

249 (242,250,255)

290d (286,293)

336

233’ (232,234)

651 (646,655)

236d (232.235,240)

310d 594

(586,601) 243d

(239,246) 315 612 831d

499d 470d 848 858

(839.847.857) (853,863)

385 (378,391)

236

-1 +2

- 14 +2

-14

-11 +50

-23 F

+19 R

+6

-6 +28

-11 F

+19 R

+ 15

OF +lF + 10 R + 10 R

+10 c + 10 c

+ll F +ll F + 12 +15

-3 F -3F

+3 R +3 R

+16C +2s c

+13 R +28 R

-1

-12 -1

+35 -13

-1

-5 -8

+29 -10

+7 F


TABLE 5-Continued

Core

Midpoint picks (cm/depth of core)

Relative offsets

SST, SST, VO w (in Cm)b SST /8i80 SST I@‘0 I

V22-108

V22-174

V22- 182

V22-196

V23-82 V25-59

V27-20 V27-86 V28-14 V28-56 V28-127

V28-238

V28-304 V28-345c V29-29’ V29-179 v30-97 V32-126 V32-128 v34-88f Y71-6-12

Y72-11-l

543d (531,555)

305 Fd

530d

301 Fd (299.303)

422 Cd 414 F

(410,417) 583d 583d

364 (361,367)

507

528

507

527

547 (544,550)

226 (224,228)

551

583 589

522 521 586 586 223 220

562 2OOd

1251 (1244,1257)

562 206d

(201,211) 1258

(1243,1253, 1278)

530d

294d (287,295,301)

414

609 (601,617)

808d 359d

512

537d

554

223 (219,226)

571 433 527 587 223d

544 184d

(181,186) 1238

(1234,1241)

t 13 0

tll F +7F

+8 C OF

-26 -26

+5

-5 -5

-9 -10

-7 -3

+3

+ 12 + 18

-5 -6 -1 -1

0 -3

+18 + 18 + 16 +22

+ 13 +20

Note. C = Coccolith midpoint/offset; F = foraminiferal midpoint/offset; R = radiolarian midpoint/offset. u Cores with multiple midpoint crossings are listed with the average value boldfaced and actual multiple picks

in parentheses. No picks were made in cores with indistinct b’*O/SST trends or insufficient sampling density across transitions.

b Negative values denote SST lags @O. Positive values denote SST leads 8’*0. c Foraminiferal temperatures were obtained using a calendar equation; T, is synonymous with August and

T, is synonymous with February. d Uncertain picks due to truncation of bottom of record (in Stage 6). e Used only Cibicides wuellerstorfi to define Stage 5e 8i80 value. f Foraminiferal temperatures were obtained using a calendar equation; T, is synonymous with February and

T, is synonymous with August.

depth offsets of transition midpoints into here a value of 11,500 yr to represent ex- time offsets. We have previously summa- actly one-half of the 23,000-yr precessional rized the evidence that oxygen isotopic cycle that presumably gives the isotopic Substage 5e lasted - 11,000 yr. We used curve its Substage 5e/5c form. We calcu-


TABLE 6. RELATIVE OFFSET BETWEEN MIDMINTS OF STAGE 5e/5d @so TRANSITION AND ASSOCIATED GLACIAL SST COOLING

Core SST,

Midpoint picks (cm/depth of core)a

SST, ?PO

Relative offsets (in cm)b

SST /6’*0 w SST 16’aO s

A180-73 D117 E49-18

K-11 K708- 1 M12392-1 MD73025 RC8-39C

RC8-145 RC 10-65 RCl l-86 RCll-120C

RCll-210

RCl l-230 RC12-294

RC12-339<

RC13-205

RCl3-228 RC13-229 RC15-61

RC17-69’ RC17-98<

TR126-23 TR126-29 v12-122 V18-68

V19-29

v19-53 V21-146 V22-38 V22-108 V22-174

v22-182 V22-196

249 462

(450,457,479)

768 788

1290 797 F

858 R

405 F 409 F 425 R 425 R

252 c (239,246,250,

254,269) 249 F 317

(306.318,321,324) 214 F 216 R

563

449

815 (806,823)

508 280 F

(278,281)

544 (540,547)

238 469

(459,479)

768 780

1292 796 F

858 R

255 C (239,256,270)

249 F 324

(317,325,330) 214 F 218 R

(216,220)

450d

815 (807,823)

508

525 (511,539)

247 427

(409,444)

787 808

1234 772

(766,777)

187

407 (403,410)

179 (177,181)

218 249d

311

209e

588d 281 575

(573,577)

288 (284,292)

567 778d

421 (410,431)

802

360

481d +27 i27 270d +lO F

533 +11 -8

i-2 -9 +35 i-42

-19 - 19 -20 -28 +56 +58 +25 F +24 F

+86R +86 R

-2F +2 F +18R +18 R

+3 c +6C

OF OF +6 + 13

+5 F +5 F +7R +9R

-4

+28 +29

+ 13 +13


TABLE 6-Continued

Core

Midpoint picks (cm/depth of core)a

Relative offsets

SST, SST, PO

SST~~

IS

$” cdb

SST /8’*0 s

V23-82 695 708 (699,716)

733 (727,738)

328

(463;;4)

505d

-38 -25

+1

-3

+11

0

-2

+9

V25-59 V27-20 470

V27-86 V28-14 V28-56 V28-127

502

514 512 (499,515,529) (495,529)

503

V28-238 V28-304 V28-345c V29-29’

526 384

(380,387) 487

(484,489) 562

V29-179 480 478 -7

- 15

-9

- 17 v30-97 547 545 (542,552) (540,550)

V32-126 V32-128 V34-88f

(48:0;124) (462:;2, + 16

+39

-21

+45 Y71-6-12 Y72-11-l 1208

(1205,1211) 1214 1169

(1159,1178)

Note. C = Coccolith midpoint/offset; F = foraminiferal midpoint/offset; R = radiolarian midpoint/offset. (2 Cores with multiple midpoint crossings are listed with the average value boldfaced and actual multiple picks

in parentheses. No picks were made in cores with indistinct #*O/SST trends or insufficient sampling density across transitions.

b Negative value denotes SST lags 6’*0. Positive value denotes SST leads 6’*0. c Foraminiferal temperatures were obtained using a calendar equation; T, is synonymous with August and

T, is synonymous with February. d Uncertain picks due to truncation of top of record (in Stage 5d). e Used only Cibicides wuellerstorfi to define Stage 5e 8’*0 value. f Foraminiferal temperatures were obtained using a calendar equation; T, is synonymous with February and

T, is synonymous with August.

lated mean Substage 5e sedimentation rates in each core by dividing the thickness of core lying between the isotopic Stage 6/5 boundary (listed in Table 5) and the Sub- stage 5e/5d boundary (listed in Table 6) by the 11,500-yr interval of time. The resulting mean sedimentation rates that were calculated are listed in Table 7. For several cores with inadequate isotopic definition of the 5e interval, we derived the mean sedimenta-

tion rate from Stage 5e to the core top (Table 7).

There is no assurance that the mean sedimentation rates calculated even for the Substage 5e interval apply specifically at the 615 and 5e/5d transitions; however, no finer chronostratigraphic distinctions can be made with reasonable confidence, and we thus consider these rates as optimal estimates for the lead-lag analysis.


TABLE 7. MEAN SEDIMENTATION RATES IN cm/1000 ~~FORTHE LAST INTERGLACIAL INTERVAL

Core Sedimentation

rate0 core Sedimentation

rate”

Also-73 D117 E49-18 K-11 K708- 1 M12392-1 MD73025 RCS-39 RC8-145 RClO-65 RCll-86 RCll-120 RCll-210 RC 11.230 RC12-294 RC 12-339 RC13-205 RC13-228 RC13-229 RClS-61 RCl7-69 RCl7-98 TRl26-23 TRl26-29 v12-122 VlS-68

2.Sb 1.8 7.3 2.3b 6.8b 3.1

12.4 7.9 1.7 2.2b 2.3b 2.8 1.0 2.7 3.6’ 2.2 2.ld 5.5c 2.5 1.7 1.96 2.3 3.9 4.6’ 2.3h 3.7b

V19-29 v19-53 V21-146 V22-38 V22-108 V22-174 V22-182 V22-196 V23-82 V25-59 V27-20 V27-86 V28-14 V28-56 V28-127 V28-238 V28-304 V28-345 V29-29 V29-179 v30-97 V32-126 V32-128 V34-88 Y71-6-12 Y72-11-l

4.9 1 .Sb 2.2 1.9” 4.3c 2.1’ 3.3* 6.6 6.5’ 2.7’ 3.7 2.6b 2.8’ 1.9* 4.4 1 .I6 3.56 3.9 4.3 3.5 2.2 1.7b 2.4’ 4.9 1.3b 6.0’

y Values are based on mean rate during 11,500-yr interval from Stage 6/5 to 5elSd transition midpoints except as in b.

b Sedimentation rates are calculated on the entire interval from isotopic Stage 5e peak to the top of the core (122,000-yr interval).

c Value is uncertain because of slow sedimentation rates, noisy isotopic trends, and/or possible truncation of record before reaching full Stage 6 or Stage 5d isotopic maxima.

d Used Cibicides wuellerstorfi only for determination of Stage 5e.

The resulting lead-lag offsets expressed in thousands of years are shown in map form in Figures 58a and b and 59a and b for the 615 and 5el5d isotopic transitions, respectively.

Cross-correlation analysis. Although di- rected toward the goal of determining iead- lag relationships, cross-correlation analysis achieves this result in a different (and com- plementary) manner by integrating the relative offset of entire data sets rather than focusing just on the rapid transitions. In es- sence, the technique involves moving one data series progressively across the other to determine which relative offset gives the highest correlation coefficient. This offset determines the lead or lag.

In this project, the cross-correlation analysis was done in the depth, rather than in the alternative time, domain. To compute the cross-correlation values, we first made linear interpolations at l-cm intervals between actual data points in both the 6180 and SST records. The two records were then translated across one another at l-cm increments. The cm offsets which maxi- mized the correlation coefficient for each comparison of 6180 with SST and the correlation coefficients at the level of maximum correlation are listed in Table 8, with positive cm offset values again indicating earlier response of the local SST curve than the 6180 curve, and conversely.

Finally, we converted cm offsets to time offsets using the calculated sedimentation rates (Table 7). These lead-lag values in thousands of years are mapped in Figures 60a, b. Cores in which the correlation coefficient never exceeded 0.50 were omitted, as were those cores left out of the previous lead-lag analysis for the reasons stated. These omissions are the blank spaces in Table 8 and the “indeterminant” cores in Figure 60.

Results

Calculated lead-lag offsets at the Stage 6/5 isotopic transition are mapped in Figure 58, those at the 5e/5d transition in Figure 59, and those from the cross-correlation analysis in Figure 60.

Several basic patterns are common to all these maps. In general, changes in Southern Hemisphere SST curves occurred earlier than al80 (-ice-volume) changes (Hays et al., 1976; Hays, 1978). The observed SST lead is commonly some 2000- 4000 yr, with the most consistent pattern in the highest latitudes (Tables 5, 6,8). In general, the radiolarian estimates suggest a substantial SST lead in the region at 40% whereas foraminiferal estimates suggest a much smaller lead or even no significant lead or lag (Figs. 10, 14). Most east-central South Atlantic cores show SST leads, extending this pattern northward to the


FIG. 58. (a) Observed lead or lag (in lo3 yr) of estimated local February SST relative to 6’sO (-ice volume) across the isotopic Stage 6/S boundary; derived by measuring offsets of transition midpoints. (b) Observed lead or lag (in 10’ yr) of estimated local August SST relative to St*0 (-ice volume) across the isotopic Stage 6/5 boundary; derived by measuring offsets of transition midpoints.


FIG. 59. (a) Observed lead or lag (in lo3 yr) of estimated local February SST relative to 6’*0 (-ice volume) across the isotopic Stage 5e/5d boundary; derived by measuring offsets of transition midpoints. (b) Observed lead or lag (in lo3 yr) of estimated local August SST relative to 6’*0 (-ice volume) across the Substage 5e/5d boundary; derived by measuring offsets of transition midpoints.


6

m I.

RCl2-

294

+ 8

(0.8

81)

F +9

(0

.849

) F

V28-

14

- 5

(0.9

19)

-5

(0.8

96)

+ 11

(0

.652

) C

+ 17

(0

.593

) c

V28-

56

RC12

-339

b +

10

(0.6

98)

+ 12

(0

.690

) V2

8-12

7 -

2 (0

.763

) RC

13

-205

’ 0

(0.8

07)

F 0

(0.8

55)

F V2

8-23

8 +

I (0

.834

), +

10

(0.7

49)

+ 1

(0.7

12),

-8

(0.6

29).

+ 10

(0.

612)

+ 8

(0.4

68)

R +

8 (0

.468

) R

V28-

304

-9

(0.8

19).

+7

(0.8

06)

+ 8

(0.8

10).

- 8

(0.3

02)

RCl3-

228

- 14

(0.

226)

F

- 14

(0.

180)

+2

0 (0

.163

) F

G.

cong

. G.

to

ng.

+21

(0.8

02)

C +

2 I

(0.7

39)

c -

13

(0.7

47).

+4

(0.4

98).

0 (0

.603

) G.

du

r. +

28

(0.8

76)

R +

27

(0.8

95)

R -

2 (0

.485

) G.

dt

rr.

+ 10

.5

(0.9

37)

RCl3-

229

-20

(0.5

57)

F -2

0 (0

.467

) F

V28-

345b

+

14 (

0.86

8)

+ 10

.5

(0.9

37)

RCl5-

61

+ 10

.5

(0.4

43).

-6

(0.3

35)

- 7

(0.7

32)

V29-

29b

+ 15

(0.

069)

+

14 (

0.11

1)

RC

17-6

9b

+ 6

(0.7

05)

+ 9

(0.5

28)

V29-

179

-4

(0.9

11)

- 5

(0.9

18)

RC17

-98b

-

0.5

(0.3

38)

v30-

97

- 2

(0.4

85)

- 2

(0.4

75)

TRl26

-23

- 8

(0.8

78)

- 7

(0.6

65)

V32-

126

- 4

(0.9

26)

-3

(0.9

42)

TRl26

-29

8.5

(0.6

92)

- 19

(0.

657)

V3

2-12

8 -

10 (

0.83

6)

- 11

(0

.817

) v1

2-12

2 0

(0.7

87)

G.

sacc

ctlife

r -2

(0

.841

) G.

sa

ccul

ifer

V34-

88d

+47

(0.3

41),

+21

(0.2

02)

+ 11

(0

.741

) +0

.5

(0.8

60)

G.

rube

r -0

.5

(0.8

92)

G.

rube

f Y7

1-6-

12

+ 14

(0.

504)

+

13

(0.5

46)

Y72-

11-l

+ 25

(0

.770

) +

25.5

(0

.623

)

Nofe

. C

= Co

ccol

ith

tem

pera

ture

es

timat

es:

F =

fora

min

ifera

l te

mpe

ratu

re

estim

ates

; R

= ra

diol

aria

n te

mpe

ratu

re

estim

ates

. Cl

Val

ues

show

n ar

e th

e cm

of

fset

s th

at

give

the

high

est

corre

latio

n co

effic

ient

(in

pa

rent

hese

s)

betw

een

6’sO

an

d SS

T ac

ross

th

e la

st

inte

rgla

cial

inter

val.

b Fo

ram

inife

ral

tem

pera

ture

s we

re

obta

ined

us

ing

a ca

lend

ar

equa

tion:

T,

is

sy

nony

mou

s wi

th

Augu

st

and

T,

is

syno

nym

ous

with

Fe

brua

ry.

c Us

ed

only

Cibi

cides

wu

elle

rsto

rfi

to

defin

e st

age

Se 6

’“O

value

. d

Fora

mini

fera

l te

mpe

ratu

res

were

ob

tain

ed

usin

g a

cale

ndar

eq

uatio

n;

T,

is s

ynon

ymou

s wi

th

Febr

uary

and

T,

is s

ynon

ymou

s wi

th

Augu

st.

oc


FIG. 60. (a) Observed lead or lag (in lo3 yr) of estimated local February SST relative to 6i*O (-ice volume) across the entire isotopic Stage 6/5e/Sd interval; derived by cross-correlation analysis. Values of lead-lags followed by * have a correlation coeffkient <OS. (b) Observed lead or lag (in lo3 yr) of estimated local August SST relative to S’s0 (-ice volume) across the entire isotopic Stage 6/5e/5d interval; derived by cross-correlation analysis. Values of lead-lags followed by * have a correlation coefficient <0.5.


equator. Other scattered cores also show SST leads: RClO-65 (Fig. 12) and Y71-6-12 (Fig. 53) in the eastern South Pacific, Y72- 11-l (Fig. 54) in the northeast Pacific, V28- 345 (Fig. 46) in the eastern Indian Ocean, and both RC12-339 (Fig. 18) and V34-88 (Fig. 52) in the nothern Indian Ocean. The very large lead in Y71-6-12 (Fig. 53) may be partly an artifact of the poorly defined 5e level in the benthonic isotopic curve. In these areas, however, there are fewer cores to define the regional response and in some cases adjacent cores show different responses. We regard these other cores only as tentative evidence of a regional trend. Not all Southern Hemisphere cores show the SST lead; some eastern boundary current cores and most equatorial Pacific cores have either indeterminant trends or else negligible leads or lags (Figs. 58-60).

The North Atlantic Ocean between 30” and 70”N is the region of the other major reproducible trend: a lag in SST response behind the al80 changes (Ruddiman and McIntyre, 1979, 1981a). The lag is most pronounced in cores north of 40”N, but is also present in the Canaries Current (core M12392-1; Fig. 8) and in the Gulf of Mexico (cores TR126-23, Fig. 25; TR126-29, Fig. 26). The SST lags are as large as 6000-7000 yr on the 5e/5d transition, but are generally smaller (1000-2000 yr) on the 615 transition and in the cross-correlations (Tables 5, 6, 8).

Clear definition of the North Atlantic SST lag dies out toward the center of the subtropical gyre, where SST vary within very narrow limits (core RC8-145, Fig. 11; Crowley (1981)). Equatorial Atlantic cores are either indeterminant or show locally conflicting leads and lags.

Although the general lead-lag pattern for equatorial cores is not so clear as at higher latitudes, several cores do show a systematic response. Cores V29-29 (Fig. 47) and V34-88 (Fig. 52) (equatorial Indian Ocean) have warm SST peaks on both the Stage 6/ 5 and 5e/5d isotopic transitions. This pattern is also visible to some extent in foraminiferal estimates from eastern South At-

lantic cores RC13-228 (Fig. 20) and RC13- 229 (Fig. 21), although both peaks are not evident in the other biotic estimates. Our methodology for quantifying leads and lags can miss this kind of pattern because the SST lead that would arise from having an SST maximum on the Stage 615 transition preceding the minimum in ice volume tends to be cancelled by the lag due to the other SST maximum on the 5e/5d transition.

Other low to middle-latitude cores show some suggestion of this pattern: RCS-145 (Fig. ll), RC12-294 (Fig. 17), V22-174 (Fig. 34), V28-127 (Fig. 43), V28-345 (Fig. 46), and Y72-11-l (Fig. 54).

PROBLEMS AND COMPLICATIONS

The results discussed previously are based on direct observations of 6180 and SST signals in 52 deep-sea cores. To decide whether these results are significant, we must address three problems: (1) the validity of the assumption that the 6i80 curves used as standards of reference actually reflect ice volume; (2) the validity of the local SST estimates in each ocean region; and (3) the impact of bioturbation on both the 6i80 and SST records. Each problem introduces complications in the interpretations of al80 and SST records; these complications are compiled in Table 9.

Validity of the 61a011ce-Volume Assumption

The first-order assumption that al80 in foraminiferal shells reflects global ice volume is complicated by temperature, dissolution, meltwater, and the mixing time of the oceans. We will argue here that there probably are temperature and other overprints on the dominant ice-volume signal in the 6180 records, but that in most cases these overprints do not appear to alter significantly the timing of the ice-volume signal in the al80 record from region to region.


TABLE9. EVIDENCEOFCOMPLICATIONSINTHE INTERPRETATIONOFCLIMAT~CRECORDSFROMTHE LAST INTERGLACIAL LEVEL

Core

ZY80 complicationsN

Abundance Discordant changes in plank./ben. #*O carrier curves

No-analog biota

SST complications

Discordant multiple

SST estimates

Abundance changes in SST signal

carrier

A180-73 D117 E49-18 K-11 K708- 1 M12392-1 MD73025 RCS-39 RC8-145 RClO-65 RCl l-86 RCll-120 RCll-210 RCll-230 RC 12-294 RC12-339 RC13-205 RC13-228 RC 13-229 RC15-61 RC17-69 RC17-98 TR126-23 TR126-29 v12-122 V18-68 V19-29 v19-53 V21-146 V22-38 v22-108 V22-174 V22-182 V22-196 V23-82 V25-59 V27-20 V27-86 V28-14 V28-56 V28-127 V28-238 V28-304 V28-345 V29-29 V29- 179 v30-97 V32-126 V32-128 V34-88

X

X

X X X X X

J

J

X X X

XII XB X X X

X

X

xB

X X

X X X

X X X X X

X X X X X

X

X X

X X

J

J

X X

X

xF X

xc JF

JF’ xC

J

dF> xC

JF

X

J X

J X

xc J xc JF, xC X J

J

X

X X

X


TABLE 9-Continued

VO complicationsa SST complications

Core

Abundance changes in 61s0 carrier

Discordant plank./ben.

curves No-analog

biota

Discordant multiple

SST estimates

Abundance changes in SST signal

carrier

Y71-6-12 X X J Y72-11-I X X J

Note. X = no basis for an assessment; J = probable effect; ? = possible effect; C = coccolith SST signal carrier, F = foraminiferal SST signal carrier; B = oxygen isotopic curve based on benthonic foraminifera; P = oxygen isotopic curve based on planktonic foraminifera.

(1 All 13~~0 curves are potentially subject to temperature and other effects noted in text. All &I80 curves based on planktonic foraminifera are subject to dissolution effects.

Temperature and dissolution overprints. Temperature affects @O values through the thermodynamic relationship; the gradient of 0.22%0 per “C, equates to a 4S”C temperature change for a shift in S’*O values of 1.0%0. For this reason, we generally avoided relying on planktonic foraminifera in regions where estimated surface temperature variations exceed 3°C.

Also we generally did not use species of Globorotalia which live at subsurface depths in low-latitude areas and in some cases record signals unlike those from the surface-dwelling planktonic species and the bottom-dwelling benthonic population. These anomalous trends are probably caused by glacial/interglacial depth migra- tions and by changes in vertical T/S profiles at depths of 100 to 300 m.

changes in the oxygen isotopic composition of their shells. The evidence for this lies in the very large glacial/interglacial changes in &I80 across the Stage 6/5 boundary. Glacial/ interglacial 6l*O differences (A@O) of up to 1.6- 1.7%0 can be reconciled reasonably well with the upper range of plausible estimates of glacial/interglacial ice-volume differences (Flint, 1971; Denton and Hughes, 1981a) and with a mean oxygen isotopic composition of the additional glacial-world ice sheets of -30 to - 35%0 (Shackleton, 1967; Dansgaard and Tauber, 1969; Broecker, 1975).

We relied mainly on the oxygen isotopic signals provided by surface-dwelling planktonic species and by benthonic foraminifera. In cores where both signals were obtained, the curves are generally similar in shape and the isotopic stage boundaries coincident (e.g., Figs. 18, 19, 25, 26, 32, 35, 38, 41, 43, 46, 49). In a few cores, the two kinds of isotopic curves are markedly dis- similar (Figs. 13, 42).

Our results, however, show numerous cores with AS’*0 signals in excess of 2.%0. This includes both low-latitude planktonic foraminifera and benthonic foraminifera (Table 10). The mean A&l80 across the Stage 6/5 transition is 1.91%~, for 30 cores wth reasonably complete benthonic foraminiferal signals and 1.64%0 for 32 cores with comparably complete planktonic foraminiferal records (Table 10). Recent detailed studies of cores with high sedimentation rates have often found A6l8O values in excess of 2.0%0 (Shackleton, 1977; Streeter and Shackleton, 1979; Duplessy et al., 1981; Ruddiman and McIntyre, 1981a).

Even the low-latitude planktonic forami- These large A6180 values, if combined nifera and the benthonic foraminifera on with the assumed mean isotopic composi- which we placed the greatest reliance may tion of excess glacial ice over that present have recorded significant temperature today, lead to implausibly large calculations


TABLE 10. GLACIAL/INTERGLACIAL CHANGE IN OXYGEN ISOTOPIC VALUW ACROSS THE STAGE 615e TRANSITION (IN %o)

Ocean Core Benthonic Planktonic

Indian

Pacific RC lo-65 RCll-210 RCll-230 RCIS-61 Vl9-29 v19-53

Antarctic E49-18 MD73025 RC8-39 RCl l-120 Vl8-68 v22-108 Mean

North Atlantic Al80-73 Dl17 K-11 K708- 1 Ml2392-1 RC8-145 TR126-23 TRl26-29 v12-122

V22-196 V23-82 V25-59 V27-20 V27-86 V28-14 V28-56 V28-127 V29-179 v30-97 Mean

South Atlantic RCl l-86 RC 12-294 RCl3-205 RCl3-228 RCl3-229 V22-38 V22-174 V22-182 Mean

RCl2-339 RCl7-69 RC17-98 V28-345 v29-29 V34-88 Mean

1.86

1.86 ” 0.00, n = 1

1 .26a

2.13

1.99 1.73”

2.06 2.07” 2.03a 2.05

1.90a

1.70 2.22 2.52 1.97 If: 0.31, n = 12

I .92” 2.16b 2.38 1.63” 1.39” 1.72O 1.77 1.85 + 0.33, n = 7

2.50

1.96

1.66 2.04 2 0.43, n = 3

1.50 2.29O 1.34 1.75

1.26” 1.63 ? 0.41, n = 5

1.63

1.64 1.61” 1.58“ 2.19 (G. ruber)O 2.10 (G. snccutifer)O

1.83

1.76

1.79 r 0.23, n = 8

3.43a 1.30

1.22” 1.17 2.09” 1.84 2 0.96, n = 5

1.99 0.95” 1.70 1.91 2.33 2.47 1.89 zt 0.54, n = 6

1.26 1.27 1.53

2.31 1.73

1.04


TABLE IO-Continued

ASi

Ocean Core Benthonic Planktonic

V21-146 V28-238 V28-304

V32-126 V32-128 Y71-6-12 Y72-11-t Mean

Global Global Mean

2.12 1.92 1.37

0.80 (G. dut.)”

1.28 (G. tong.) 0.91”

1.40” 1.46O 1.76 1.81 -c 0.33, n = 7 1.18 2 0.24, n = 8

1.91 * 0.32, n = 30 1.64 5 0.54, n = 32

a Possibly missed Stage 6 isotopic maximum and/or Stage 5e minimum, thus underestimating total glacial/ interglacial change.

b Used only Cibicides wuellerstorfi to define Stage 5e 6’*0 value.

of ice volume on the continents, as well as unrealistically large sea-level lowerings (see Broecker (1975) for a discussion of some of the constraints on these calculations). This contradictory result has thus forced a reexamination of isotopic records for effects other than ice volume.

There is a weak consensus that glacial bottom waters were colder than today, especially in the Atlantic Ocean (Streeter and Shackleton, 1979; Duplessy et al., 1980; Broecker, 1981). If correct, the most likely effect of bottom-water temperature changes on oxygen isotopic signals in benthonic foraminifera will be to increase the total amplitude by shifting glacial values toward more positive readings and interglacial values toward more negative readings. The size of this shift is difficult to assess. Most modern bottom-water temperatures near our core sites are in the range of 0” to 3.X. The upper limit of additional cooling (to the freezing point of sea water) is thus about 2” to YC, which corresponds to isotopic shifts of 0.35 to 1.2%0. Bottom temperature effects could range from being small (-15% of the mean observed shift) to substantial (-50% of the largest observed A#*0 change in shallow cores potentially

subject to largest deep-water toolings). There is no way for us to assess bottom temperature effects on a core-by-core basis.

Comparable uncertainties exist in the planktonic foraminiferal S1*O records. The mean A6180 value for all cores with complete records is 1.64%0 (Table lo), the upper limit attributable to ice-volume effects alone. Ocean-to-ocean differences in planktonic foraminiferal AV*O values are also large. Lowest ASI shifts occur in the Pa- cific Ocean, partly because of low density sampling in cores with low sedimentation rates. Preferential dissolution during interglacials and preservation during glacials of foraminifera with low @*O values could also reduce the mean Aal amplitude in planktonic foraminifera from the Pacific (Erez, 1979). In contrast, the mean of In- dian Ocean, North Atlantic and South At- lantic Ati’*0 shifts for planktonic foraminifera are very close to the average value of 1.91%0 for benthonic foraminifera (Table 10). Erez (1979) suggests that the larger At- lantic amplitude is due to a dissolution effect of 0.4%0 that is additive to a basic ice- volume signal of about 1.5%0. Antarctic A@0 shifts are somewhat lower, despite

the likelihood of temperature overprints at regarded analyses from three Norwegian high latitudes. Sea cores (K-11, Fig. 6; V27-86, Fig. 40;

The large average A&l80 values and wide V28-56, Fig. 42). In general, we regard range of shifts from core to core in plank- cores located poleward of the glacial-age tonic foraminifera (Table 10) imply that tem- polar fronts (--40”N, -403) the most sus- perature and other factors have altered the pect for meltwater overprints on the plank- basic ice-volume message to a degree com- tonic foraminiferal al80 records. parable to that in benthonic foraminifera. In Some scientists contend that low-latitude principle, the independently determined planktonic foraminiferal species can record SST estimates could be used to remove the meltwater overprints on the isotopic signal. temperature effect from the planktonic for- Berger and associates (Berger et al., 1977; aminiferal al80 signal (e.g., Imbrie et al., Berger, 1978) proposed that a low-salinity 1973). In practice, however, this requires meltwater lid blanketed the entire world complex decisions on past seasonal and ocean about 12,000 yr B.P. If true, this depth preferences of the species used. Each could bias low-latitude al80 records based decision increases the uncertainty of the on surface-dwelling planktonic foraminif- correction. Still remaining in the &I*0 signal era. Recently, Jones and Ruddiman (1982) are other factors that cannot be indepen- have argued that the global meltwater spike dently estimated with sufficient accuracy, inferred by Berger can be more plausibly such as the effect of the local precipitation- interpreted as an artifact of overestimating evaporation balance on local surface-water the mixing (and thus the “unmixing”) op- salinity. Dissolution effects remain a pos- erator. This overestimation causes the sible complication in all #*O curves from small-scale structure actually present on planktonic foraminifera. the isotopic curves (and reflecting a com-

Meltwater complications and oceanic bination of dissolution effects and the ac- mixing time. Meltwater and melting ice- tual ice-volume history) to be blown up in bergs deliver 160-enriched water directly to the unmixing process into false meltwater the mid-latitude oceans and marginal seas spikes. during deglaciations (Kennett and Shack- The majority of isotopic Stage 615 tran- leton, 1975; Emiliani et al., 1975). This sitions based on planktonic (or benthonic) creates the possibility of a superimposed foraminifera shows no obvious spikes or isotopic signal on the Stage 615 boundary, more muted negative excursions that could either as discrete “spikes” if the delivery be attributed to sudden meltwater influxes occurs in sudden catastrophic pulses, as (Figs. 3-54). The negative excursions that broader “shoulders” if the delivery is are observed appear to be scattered ran- spread over a longer period of time, or as domly across different parts of the transi- a change in rate of the al80 transition if the tion. This suggests that these excursions influx is more subtle. It also raises the pos- are artifacts of heterogeneous mixing, sibility of increased mixing time of the deep bottom-current redistribution, or analytical oceans due to reduced bottom-water for- error. mation caused by increased salinity strati- More realistically, meltwater could cause fication in high-latitude oceans. a translational offset of the position of the

We have avoided the worst effects of isotopic transition. Two opposed scenarios meltwater on planktonic isotopic curves by are possible. A sufficiently rapid influx of minimizing the use of planktonic species in meltwater could, by lightening the isoto- the high-latitude northern oceans adjacent pic composition of ocean water sufficient- to the great Northern Hemisphere ice ly, create a negative aI80 “ledge” or sheets. Meltwater effects in this area were shoulder on the Stage 6/5 isotopic suggested by Kellogg et al. (1978). We dis- boundary. This effect could blend smoothly



into the sigmoidal shape of the deglacial isotopic transition, offsetting it deeper into the core from where the true ice-volume signal would otherwise lie.

The converse effect-a delay in the reg- istering of deglacial ice-volume changes by the isotopic curve-could also be caused by meltwater. Because of its low salinity and hence low density, meltwater can increase the near-surface stratification in oceans near the melting ice sheets. This stability could slow or stop bottom-water formation in the very areas where it primarily occurs today. As a result, overturning of the deep oceans might take considerably longer than the 1000 yr characteristic of modern circulation. Thus, recording of deglacial ice-volume changes in deep ocean waters and in benthonic foraminiferal shells might be delayed.

Summary of 6’*0. We conclude that oxygen isotopic records from benthonic and planktonic foraminifera contain temperature (and other) effects that could disrupt simple ice-volume interpretations (Table 9). The central question of the impact of these effects on the 6’*0 curves is timing. De- pending on their phasing relative to the true ice-volume portion of the 6i80 signal, they may have no impact or a potentially harmful effect on the assumption of a tight 6l*O/ice-volume linkage.

Consider the Stage 6/5 transition in any of our cores. If the deglacial 6180 shifts measuring true ice-volume change are matched increment-for-increment by local cold-to-warm temperature changes across the entire transition, then the oxygen isotopic curve will still record the timing of ice-volume changes correctly. If, however, the local temperature change occurs only in a brief interval confined to a small portion of the isotopic transition, then the 6**0 transition will not give an accurate picture of the timing of ice-volume changes. In short, local temperature, meltwater, or dissolution overprints could be either benign or malignant depending on their timing relative to true ice-volume changes.

Seemingly this question could be evaluated by examining the most recent deglacial transition for evidence of synchroneity or nonsynchroneity in the isotopic signals from region to region. However, the 14C and 6i*O data bases are inadequate. Most records have too widely spaced 14C dates and 6i*O analyses or the sedimentation rates are too low. Cores with higher sedimentation rates tend to occur in coastal- margin or high-latitude areas where many potential 14C contamination problems exist.

The midpoint ages of the best-dated, best-resolved oxygen isotopic curves across the last deglaciation are: 10,600 yr B.P., planktonic foraminifera, Caribbean core P6304-9 (Rona and Emiliani, 1969); 11,800 yr B .P., planktonic foraminifera, equatorial Indian Ocean core V 19-188 (Peng et al., 1977); 13,000 yr B.P., planktonic foraminifera, western equatorial Pa- cific box core ERDC92 (Berger and Kil- lingley, 1977); 14,500 and 11,500 yr BP, planktonic and benthonic foraminifera, respectively, equatorial Atlantic core CH22KW31 (Pastouret et al., 1978); and 13,200 yr B .P., benthonic foraminifera, North Atlantic core CH73139 (Duplessy et al., 1981).

This 4000-yr range (14,500 to 10,500 yr B.P.) includes dates older than the commonly accepted deglacial midpoint of 11,000 yr B.P. proposed by Broecker and Van Donk (1970). Duplessy et al. (1981) suggested that the last deglaciation occurred in two discrete phases, one from 16,000 to 13,000 yr B.P. and the other from 10,000 to 8000 yr B.P. Independent evidence also suggests a major phase of Northern Hemisphere deglaciation prior to 13,000 yr B .P. (Ruddiman and McIntyre, 1981b, c). On the other hand, Berger (1982) recently summarized evidence for a midpoint transition closer to 11 ,OOO- 10,000 yr B.P. in the Pacific.

With both the basic form and the midpoint age of the last isotopic termination still substantially in doubt, little can be learned from the Stage 2/l boundary that is


relevant to the question of isotopic synchroneity across the Stage 615 and 5e/5d transitions. This insufficient evidence pro- hibits detailed evaluation of the effects of two major factors- temperature and meltwater-on the fidelity of the 6r80/ice- volume relationship at the last interglacial level. Significant misrepresentations of the true ice-volume signal could arise from un- critical reliance on the oxygen isotopic records.

It is, however, encouraging that most low- to mid-latitude cores with both planktonic and benthonic foraminiferal isotopic records have similarly phased #*O signals. It could be argued that both signals are substantially altered by temperature effects in a similar and malignant way; more plausibly, the signal alterations are spread fairly evenly across the transition and thus are largely benign. We see no compelling reason to reject the first-order assumptions of isotopic synchroneity and isotopic fidelity to the ice-volume signal made earlier in this paper and used generally by the paleoclimatic community.

Validity of the SST Estimates

Numerous sedimentologic factors could affect the reliability of the SST estimates used in this study, including, primarily, differential solution of calcareous organisms (Ruddiman and Heezen, 1967; Berger, 1968, 1970, 1973; McIntyre and McIntyre, 197 1)) differential preservation of siliceous organisms (Johnson, 1974; Mikkelsen, 1978; Takahashi and Honjo, 1981), and differential transport by bottom currents (Berger, 1971; Berger and Piper, 1972).

Evaluating each of these factors in each core is impossible. Instead we have relied on other indicators of significant complications. We began with the assumption that the standard errors of estimate characteristic of each transfer function (Table 3) are valid representations of the errors in downcore estimates. Two lines of evidence can warrant the conclusion that various other factors have rendered the SST estimates in-

valid: (1) indications of a no-analog fauna or flora and (2) discordant SST estimates between two biotic (fauna1 or floral) groups.

No-analog conditions. The most detailed treatment of the no-analog problem is that in Hutson (1977), who suggested that either low communalities or environmental estimates lying outside of the range of the calibration data set are the clearest indication that a transfer function has encountered a fauna1 or floral assemblage not encompassed by modern assemblages. This can be caused by (1) post-depositional altera- tion of originally “normal” assemblages, (2) the presence of an oceanic environment unlike those known today, or (3) evolution.

Communalities are low (0.62 to 0.70) for foraminiferal assemblages in nine samples of South Atlantic core RC12-294 (250 cm, 256-266 cm). Although the percentage of Globigerina falconensis at 264 and 266 cm (37 and 42.5%) exceeds the maximum percentage in the calibration data set (36%), factor loadings for the core samples do not exceed those for the core tops.

Only one core in this study yields temperature estimates even slightly outside the range of the relevant transfer function: core V22-196 at 13”50’N along the Saharan coast of Africa in the North Atlantic (Fig. 36). The summer SST estimate in one Stage 5e sample of this core reaches a value of 29.3”C, whereas the modern Atlantic Ocean atlas values used in the equation are no- where warmer than 28.8”C. At this and several surrounding levels in the same core, the communalities fall to low values (0.4 to 0.7). This indicates that the foraminiferal fauna1 composition in these levels is not encompassed by the modern biota used in the transfer function (Table 3 ; FA13 of Kipp, 1976).

Inspection of the fauna1 composition in core V22-196 reveals the probable source of the no-analog condition. In all levels with anomalously warm temperature estimates and/or low communalities, the percentage abundance of the Globorotalia menardiil


Globorotalia tumida complex reaches values larger than those normally found in the core-top data set of transfer function FA13 (Kipp, 1976). One sample in the core- top data set did have higher abundances of this group than any last interglacial count in V22-196. But this single sample appears to have been insufficient to stabilize the transfer function completely against the impact of high abundances of these species. Thus the moderate warmth normally estimated by these low-latitude species is ex- aggerated into slightly unrealistic warmth during the last interglaciation.

There are two possible explanations of these marginally no-analog percentages. The species G. tumida and G. menardii are resistant to dissolution and could be resid- ually concentrated by differential destruction of other, more fragile species. This could mean that intense dissolution contrib- utes to the anomalously warm temperatures in this core. Gardner and Hays (1976) noted strong interglacial dissolution in core V22- 196, particularly at isotopic Substage 5e.

The second explanation involves a no-analog surface-water condition in the last interglaciation. Thiede (1975) found unex- pectedly high percentages of G. menardii and G. tumida in surface planktonic tows taken within 100 km of the African coast from 15”N to 23”N. He interpreted this surface occurrence of normally deeper-habitat species (Jones, 1967) as indicative of upwelling from a warm subsurface counter- current flowing northward along the upper continental slope and outer continental shelf. Unusually strong upwelling at the last interglaciation could thus explain the no- analog fauna.

Thiede (1975) also noted that increased dissolution would favor higher abundances of these two resistant species. Among the possible causes of increased dissolution in coastal areas, upwelling is a likely candi- date. Berger (1970) inferred that increased coastal upwelling causes a stronger rain of organics to the sea floor, the organic components subsequently decay, the oxidiza-

tion yields CO,, the CO, increases the acid- ity of the interstitial water, and this en- hances dissolution. Thus, both of the likely causes of no-analog G. menardiilG. tumida abundances are probably linked to increased upwelling.

We conclude from these observations that upwelling appears to have been inten- sified over core V22-196 at and near the peak interglaciation of oxygen isotopic Substage 5e, causing unrealistically warm SST estimates. Because upwelling normally brings cooler waters to the surface, we infer that the last interglacial SST estimates in core V22-196 may actually be anomalously warm by several “C relative to true temperatures.

Discordant multiple estimates. Disagree- ment of SST estimates obtained from more than one biotic indicator (radiolaria, foraminifera, and coccoliths) provides a second indicator of invalid temperature estimates. We used multiple groups to test for discordant estimates in two regions of the last interglacial ocean: the South Atlantic and the Indian Ocean sector of the Antarctic Ocean.

At South Atlantic lower latitudes, estimated SST differences between the biotic groups are minor. Foraminiferal estimates in three out of four cores are slightly warmer (lo-2°C) than those from the radiolaria (Fig. 19) or coccoliths (Figs. 32, 34). The exception is core V22-182 (Fig. 35), where the trend is for foraminiferal estimates to be lo-3°C cooler than those from coccoliths, especially during Substage 5e.

In the mid-to-high latitude South At- lantic, the reverse sense of difference occurs: foraminiferal temperatures tend to be colder than those from both radiolaria (Figs. 20, 21) and coccoliths (Figs. 13, 17, 20). The differences are largest in the mid- latitude cores situated in the Agulhas Cur- rent region and along the eastern boundary current (RCll-86, Fig. 13; RC13-228, Fig. 20; RC13-229, Fig. 21), and smallest in the high-latitude core located near the present position of the subtropical convergence


(RC12-294, Fig. 17). For this latter core, foraminiferal and coccolith estimates are in close agreement during Substage 5e but have differences of up to 5°C during cold periods (Stage 6 and Substage 5d). Moltino et al. (1982) showed similar discordancies for February temperatures in this part of the South Atlantic during the last glacial maximum (18,000 yr B.P.) and attributed them to a possible increase in water mass stratification along the subtropical convergence, with the foraminifera missing the warmest near-surface temperatures. An- other possibility is that this region (37” 16’S) was invaded during cooler intervals by subpolar waters for which coccoliths overestimate temperatures due to the lack of adequate cold end-member factors (Mol- fino et al., 1982). Such waters now lie only about 4” of latitude to the south of RC12- 294.

Of the three mid-latitude cores containing strikingly discordant multiple estimates, only the discordancies in RCll-86 from the Agulhas Current region seem to be readily explainable. The Agulhas Cur- rent region today is subject to one of the largest seasonal and annual variations in water masses known in the world ocean (e.g., the South Equatorial Current, the Mozambique Current, shoaling of subtropical water, and various eddies). These waters range in annual temperature from 14”- 17°C to 24”-26°C (Darbyshire, 1964, 1966; Harris, 1972; Harris and van Foreest, 1978). The differences in biotic estimates may be indicative of different water bodies seasonally occupying the Agulhas region and of differing responses of the biotic groups to those water masses.

Cores RC13-228 and RC13-229 located in the Benguela upwelling region show differences of up to 10°C among biotic estimates. RC13-228 has estimates from all three biotic groups, with temperature estimates of coccoliths and radiolaria the most similar (differences <3”C). The foraminiferal estimates are considerably colder. In RC13- 229, foraminifera are again colder than ra-

diolaria. The radiolaria, in turn, differ in their overall temperature response between these two adjacent cores: in RC13-228 their Substage 5e warming leads the isotopic 5e level; in RC13-229, it lags.

Which biotic group is more valid in estimating temperature in this mid-latitude region? Molfino et al. (1982) have shown that all three transfer-function equations work well for the modern world and the last glacial maximum, although discordant estimates are present at 18,000 yr B.P. in the region represented by the cores in this study. Coccolith estimates beyond 73,000 yr B.P. are based on a smaller fraction of the population due to substantial evolu- tionary changes that resulted in a reversal of species dominance (McIntyre et al., 1972). This dominance reversal creates a form of no-analog problem, because the species dominating the coccolith population during the last interglaciation cannot be part of the transfer-function analysis. The radiolarian estimates are based on the smallest surface-sediment calibration set of any of the three groups (Morley, 1979; Mol- fino et al., 1982). The ability of this equation to recreate a wide range of past environments from a small range of modern environments is thus suspect. Finally, the fact that the foraminiferal estimates are offset from the more concordant (and warmer) radiolarian and coccolith estimates suggests that the foraminiferal estimates could be in error. Dissolution may be driving the winter foraminiferal estimates toward warmer values at selected levels (above 685 cm in RC13-228; from 298 to 312 cm in RC13- 229); this occurs because some downcore intervals are better preserved than the nearby surface-sediment calibration data set. If true, this further aggravates the discordancy between the biotic groups.

Temperature discordancies across the oxygen isotope Stage 6/5e/5d interval in these three eastern South Atlantic cores can, however, be viewed from a different per- spective. The coccolith/radiolarian results have SST values rarely cooler than today


(1’C at most) and generally 2” to 7°C warmer even in glacial isotopic stages. The extreme example of this view is the radiolarian pattern in RC13-229, which never cools to within 1°C of the modern values. The foraminiferal results are somewhat more typical of those found in many other oceanic regions, in that the warmest temperatures are relatively close to those today. The coldest temperatures are lower than those today by as much as 4” to 8”C, usually during glacial isotopic stages but in some interglacial (Stage 5e) pulses as well. The discordancies thus create two distinct alternative choices: a last interglacial sequence consistently far warmer than today (radiolarian/coccolith data) or one rather similar to the glacial/interglacial changes over the last 15,000 yr (foraminiferal data). There is no basis for choosing between the two alternatives.

A second example of discordancy occurs between the radiolarian and foraminiferal SST estimates in cores RC8-39 (Fig. 10) and RCll-120 (Fig. 14) from the subantarctic sector of the Indian Ocean at roughly 40”s. The basic sense of difference is not due to temperature range. Both biotic groups yield very similar SST ranges (for both seasons in RCll-120 and for winter in RC8-39), although the warmest radiolarian summer SST estimates are considerably warmer than the foraminiferal estimates in RC8-39. The major difference between the radiolarian and foraminiferal estimates is in the timing of SST changes. In both cores, the radiolarian estimates rise to peak interglacial values on the Stage 615 boundary, whereas the foraminiferal estimates peak more or less at the Stage 5e isotopic minimum. Similarly, the radiolarian estimates cool to full-glacial values within the isotopic Stage 5e minimum, whereas the foraminiferal estimates cool along the Stage 5e/ 5d transition. In summary, radiolarian estimates indicate an oceanic response that leads alsO and ice volume by several thousand years; the foraminiferal estimates are nearly in phase with the 6i80 response or

lead by much smaller amounts (Tables 5, 6, 8).

Howard and Prell (1984) have reexam- ined the relationships of modern surface- sediment assemblages to surface and subsurface oceanographic properties near these cores. They note that the two most prominent foraminiferal assemblage boundaries coincide with prominent oceanographic boundaries (the subtropical and Antarctic convergences), in accordance with the idea that the foraminifera respond to surface conditions. The single prominent radiolarian assemblage boundary, however, does not coincide with the surface and near-surface oceanographic boundaries. In- stead, it appears to match more closely a deeper salinity boundary (3 150 m) that sep- arates saltier waters of the subtropical gyre to the north from less-saline Antarctic polar waters that override from the south. Howard and Prell suggest that radiolarian fauna1 changes in this region may thus be partly decoupled from sea-surface conditions and more indicative of expansions and contractions of a subsurface core layer of subtropical water.

In the Howard and Prell interpretation, the Southern Hemisphere oceanic lead may be centered not in the circumantarctic polar waters but in the subtropical central waters, at least in the Indian Ocean. In this view, the Antarctic sea-surface response at 40”s is nearly in phase with 6i80 and global ice volume.

These explanations of SST discordancies between different biotic estimates are not the final word. There is still much to be learned about the basic depth and seasonal preferences of all species in each biotic group. This information will be critical to the refinement of paleoenvironmental reconstructions in the future.

Bioturbational and Bottom-Current Mixing

Deep-sea sediment mixing by bottom- dwelling fauna and by bottom currents is potentially the largest hindrance to direct


AMPLITUDE REDUCTION

ORlGtNAL RECORD

GRADIENT REDUCTION

NIXED RECORD

TRANSLATION

FIG. 61. Major categories of bioturbation effects typical in deep-sea paleoclimatic records.

paleoclimatic interpretation of data from the last interglacial interval.

Bioturbational mixing can have several major effects on original climatic records. Several are summarized schematically in Figure 61: (1) amplitude suppression, (2) gradient reduction, and (3) translational offsets. The first two are achieved by the simple smoothing effect normally associated with mixing. The third is an important example of the complications that can arise from the problem of absolute abundance variations (see section on Data Bases, Ab- solute Abundance Counts).

The three effects in Figure 61 relate to that part of the bioturbation process which acts as a low-pass filter and smooths original climatic records. If we knew explicitly the strength of the mixing process in each of these records, we could exactly remove mixing effects and restore the original climatic record. Unfortunately, radioisotopic tracers cannot be used to quantify mixing in these older parts of the record, and we have no dispersed ash zones to use as proxy mixing tracers. Without this knowledge, our assessment of the effects of

mixing cannot achieve an explicit restora- tion of any original 6l*O or SST signals from the data shown in Figures 3-54. In view of the large variability of mi;ting effects in time and space (Jones-and Rud- diman, 1982; Ruddiman and Glover, 1982), a mean value of the mixing process cannot be utilized for all records.

Furthermore, mixing is not entirely a low-pass filter; it can also create “noise” in climatic signals. Infrequent but large bur- rowers penetrating well below the surface may carry and deposit lumps of sediment far away from their original‘level of depo- sition. The resulting burrows have a character quite different from that of the surrounding sediments. This k&d of mixing is usually called “lumpy” or “heterogeneous” mixing (Hanor and Marshall, 1971; Berger and Johnson, 1978). It in part ac- counts for the commonly observed decrease in analytical 6180 reproducibility across abrupt isotopic transitions like the Stage 6/5 boundary. Heterogeneous mixing, because it is a random process, also impedes attempts at explicit signal resto- ration.

Bottom currents are still another complicating factor. They may leave short ero- sional gaps in some records while at the same time thickening other sequences with redeposited sediments that are not wholly contemporaneous with the rain of pelagic material from above. Thus, sections thick- ened by bottom currents may yield climatic signals containing a false component. Rud- diman and Glover (1982) note that bottom currents may be more important than bioturbation in mixing many paleoclimatic records. At present, we cannot judge the impact of bottom currents on the records in this study.

Despite these complications, we can make some assessment of the kind of impact that bioturbation has had on many of our most critical climatic records. We do know the approximate range of mixing effects in near-surface sediments from radioisotopic evidence. Berger ‘and Johnson

THE LAST INTERGLACIAL OCEAN 19.5

(1978) conclude that bioturbation can be modeled by a well-mixed layer some 4 to 8 cm thick. Data from late Quaternary ash zones ,-Mggest that the thickness of the well- mixed surface layer can range from roughly 8 cm $0 as little as l-2 cm (Ruddiman and Glover, 1982). Jones and Ruddiman (1982) found evidence for well-mixed surface layers less than 2 cm thick in late glacial sediments, of the equatorial Atlantic. Jones (1980) found that the shape of the average ash zone in 1he late Quaternary North At- lantic can be modeled by a homogeneously mixed upper layer 4 cm thick and a lower layer marked by an exponential decrease in mixing intensity toward negligible values 15 cm deep. Thus, with empirical knowledge of the reasonable limits of late Quaternary mixing intensity, we can at least assess which of our critical climatic signals could have been significantly altered by mixing.

The only climatic signals (@O or SST) which can be evaluated for bioturbation effects in a .meaningful way are those for which we have absolute abundance data (Fig. 61; Appendixes 1, 3). The difference between the lower two drawings in Figure 61 suggests,the critical impact of these absolute abundance changes. If there is no change in absolute abundance of the climatic signal carrier, bioturbation may simply smooth the signal. If there is an upward increase in abundance of the signal carrier in excess of a factor of 3, the same amount of mixing will also achieve a downward translation of the smoothed climate signal (bottom drawing, Fig. 61). The converse (upward translation of the 6180 SST signal due to downward increases in abundance of the signal carrier) also can occur (Hutson, 1980).

Of the 52 cores in this study, 7 have no absolute abundance data for either the 6180 or SST record. Another 18 lack sufficient abundance data only for evaluation of the 6180 analyses, and 2 lack abundance data only for evaluation of the SST signal carrier. No direct abundance data exist for any of the six coccolith SST curves in the South

Atlantic, but we have included percentage CaCO, curves as first-order indicators of absolute coccolith abundance. These six cores have foraminiferal SST estimates and abundance data. This leaves roughly half the cores in this study with a full array of absolute abundance data (Table 9).

As stated previously, no attempt was made to restore explicitly any al80 or SST signals. Instead, we made numerous model runs to test the sensitivity of original climatic signals to these various mixing effects (Fig. 61), particularly absolute abundance changes. Published examples of our work are given in Hutson (1980), Ruddiman et al. (1980a), Ruddiman and McIntyre (1981b), and Jones and Ruddiman (1982). From this experience, we have deduced the following possible impacts of mixing on this data set.

Impact of mixing on Stage 5e SST estimates. We first examined the last interglacial interval to assess the impact of mixing on the estimated SST at the 5e level (Table 11). We primarily looked for likely in- stances of amplitude reduction and translational displacements (Fig. 61, top and bottom).

Some SST records with short intervals of interglacial warmth appear likely to have suffered peak-amplitude reduction to cooler values (e.g., foraminiferal SST in cores A180-73, Fig. 3; M12392-1, Fig. 8; RCS-145, Fig. 11; V19-29, Fig. 29; V21-146, Fig. 31). On the other hand, many cores have sufficiently broad regions of stable SST values at the isotopic Stage 5e level that peak-amplitude reduction is not a problem. This is particularly true of several Indian Ocean, subantarctic, and North At- lantic cores (K708-1, Fig. 7; MD73025, Fig. 9; V23-82, Fig. 37; V28-345, Fig. 46; V30- 97, Fig. 49; V34-88, Fig. 52). In some cores, we find the converse situation: the isotopic 5e level coincides with an SST minimum which may actually have been reduced in amplitude by mixing (that is, toward warmer values). In the Southern Hemi- sphere, the preserved Stage 5e SST min-

TABL

E 11

. IM

PACT

OF

COM

PLIC

ATIN

G F

ACTO

RS O

N ST

AGE

5e S

ST

Red

uctio

n of

Tr

ansla

tion

Tran

slatio

n R

educ

tion

of

Tran

slatio

n Tr

ansla

tion

Core

pe

ak a

mpl

itude

of

SST

cu

rve

of P

O

curv

e Co

re

peak

am

plitu

de

of S

ST

curv

e of

6’*0

cu

rve

A180

-73

- X

X V1

9-29

-

X X

D117

X

v19-

53

7 E4

9-18

+

X V2

1-14

6 -

X K-

11

- V2

2-38

+?

K7

08-

1 v2

2-10

8 X

M12

392-

1 -

X X

V22-

174

+ M

D730

25

X V2

2-18

2 RC

S-39

‘F

X

V22-

196

X RC

S-14

5 -

X X

V23-

82

+ RC

lO-6

5 X

V25-

59

RC

ll-86

V2

7-20

+

RC

ll-12

0 +

V27-

86

+?

RC

ll-21

0 V2

8-14

-+

? R

Cl

l-230

V2

8-56

RC

12-2

94

- F

V28-

127

+?

RC12

-339

+

V28-

238

X X

RCl3

-205

-

V28-

304

RC13

-228

-F

V2

8-34

5 RC

13-2

29

-F

X V2

9-29

-c

RC

15-6

1 X

V29-

179

RC17

-69

v30-

97

RC17

-98

V32-

126

TR12

6-23

X

V32-

128

- X

TR12

6-29

+

X V3

4-88

v1

2-12

2 X

X Y7

1-6-

12

X V1

8-68

-

X Y7

2-11

-l X

Nof

e.

- =

Obs

erve

d es

timat

e m

ay b

e to

o co

ol;

+ =

obse

rved

SS

T es

timat

e m

ay b

e to

o wa

rm;

X =

ther

e is

no

data

ba

se

for

a de

cisio

n on

ab

unda

nce-

rela

ted

trans

latio

nal

mix

ing;

?

= po

ssib

le

impa

ct;

F =

SST

curv

es b

ased

on

fora

min

ifera

.


imum may be clearly defined by numerous values, as in cores E49-18 (Fig. 5) and RCll-120 (Fig. 14).

Some of the low-latitude SST minima at or near the 5e level are very small in amplitude and duration, as in cores RCll-86 (foraminiferal record, Fig. 13), V22-38 (coccolith record; Fig. 32), and V22-182 (foraminiferal record; Fig. 35). Toward North Atlantic higher latitudes, these minima often appear as single-point, low-SST estimates embedded in the prevailing interglacial warmth (e.g., V27-20, Fig. 39; V27-86, Fig. 40; V28-14, Fig. 41). Attempts to correct these small-amplitude, high-frequency peaks for bioturbational mixing are not worthwhile for several reasons. Those minima based on single analyses may simply reflect variations within the 1.0” to 1.5”C standard errors of estimate and may be meaningless. Second, our selection of the 5e level is subject to some uncertainty due to sampling gaps, analytical error on the oxygen isotopic measurements, and statistical limitations reflecting the small number of benthonic foraminifera used for analysis. If we moved our 5e stratigraphic choice up or down by one or two samples, it would no longer coincide with the brief SST minimum but would coincide with levels showing warmer SST values. Thus the inferred effects of mixing on a brief SST minimum are counterbalanced by the possibility of opposite effects due to misplace- ment of the isotopic 5e level.

Translational offsets (Fig. 61, bottom) could not have affected most of the SST values at the isotopic Substage 5e level. Most of the major absolute abundance changes of the SST and 6i80 signal carriers occurred above or below, rather than within, the 5e interval. In several cores noted in Table 11, however, translational offsets of the SST curve or of the 6180 curve are likely to have reduced the estimated temperatures relative to those actually occurring during the last interglaciation: RC12-294 (Fig. 17), RC13-205 (Fig. 19), RC13-229 (Fig. 21), and VlB-68 (Fig.

28). In other cores noted in Table 11, Stage 5e SST values would probably be cooler if translational offsets due to bioturbation were eliminated: RC12-339 (Fig. 18), RC13- 228 (Fig. 20), TR126-29 (Fig. 26), V19-53 (Fig. 30), V25-59 (Fig. 38), and V29-29 (Fig. 47).

Core RC13-205 (Fig. 19) shows with un- usual clarity one kind of impact mixing can have. The 6i80 values from Uvigerina indicate a much smaller (and differently po- sitioned) Stage 5e isotopic minimum from that shown by Cibicides. We infer that Uvi- gerina did not live at this core site during the mid-Se interval and was mixed in from lower (and higher) levels. Cibicides did live at this site during Stage 5e and thus record the 3.0%0 hi80 values characteristic of the 5e level (dashed line, Fig. 19). This may also occur in core Y71-6-12 (Fig. 53). More commonly, however, analyses of multiple species at any one level in a core agree closely enough to suggest that both lived at roughly the same time and were not juxta- posed by mixing after death.

Impact on lead-lag relationships. Trans- lational offsets of large-amplitude &I80 or SST transitions (Fig. 61, bottom) can be as large as 5- 10 cm (Hutson, 1980; Ruddiman et al., 1980a). Even for cores with sedimentation rates as high as 4-5 cm/1000 yr, this means that leads or lags of 1000-2000 yr may be mixing artifacts. At rates as low as 2 cm/1000 yr, offsets as large as 5000 yr may be mixing artifacts. Translational offsets this large will occur if several conditions are met: the downcore changes in signal- carrier abundance must be large (a factor of 3 or more), must originally have occurred over a short interval (~20 cm), and must have occurred close to a pronounced change in the actual climatic signal. Under these conditions, the actual climatic signal can be translated as much as 5-10 cm within a core. This would have a very large effect on the validity of our lead-lag maps (Figs. 58-60).

Many of the al80 curves based on benthonic foraminifera are cushioned against


TABLE 12. IMPACT OF TRANSLATIONAL OFFSETS ON LEA&LAG RELATIONSHIPS

Core

Stage 6/5 transition

Increase in

tP0 (dec.) SST, SST,

Stage 5el5d transition

Decrease in

fP0 (inc.) SST, SST,

Also-73 D117 E49-18 K-11 K708- I M12392-1 MD73025 RCS-39 RC8-145 RClO-65 RCll-86 RCll-120 RCl l-210 RC I l-230 RC 12-294 RC12-339 RC13-205 RC13-228 RC13-229 RC15-61 RCl7-69 RC 17-98 TR126-23 TR126-29 v12-122 V18-68 V19-29 v19-53 V21-146 V22-38 V22- 108 V22-174 V22-182 V22-196 V23-82 V25-59 V27-20 V27-86 V28-14 V28-56 V28-127 V28-238 V28-304 V28-345 V29-29 V29-179 v30-97 V32-126 V32-128 V34-88

x

X X X X X

t X X X

X X

xB X X X X X X X

xB

X X

X X X

4 X X

t X

I X

4.

X X X X

T 1‘ X X X X X X

t t xF xF X X

X

JF

X

JF,XC

1

X X X

X X

J

X

L X X X X

4

xc X X X X

1

X

1

X

I X X

t X

J

; X

X

JF

X

L F’ xC

1

X X X

X X

J

X

1

X X X X

J

xc

I‘ X X

J

X

1

X

c

X X

t

X

t

t X

X

X X

X X

1 X X X

X X

xB X X X X X X X

xB

X

X

1

X

1

X X X X

1

X X

X X

X

t X

xF X X X

TF X X

TF? xC

r XF X X

X X X X

-1

X X X X

xc X X

X

f

X

t X

X X X X

X X

X X

X

t X

xF X X X

TF X X

fF,XC

r XF X X

X X X X

1

X X X X

xc X X

X

t

X

t X

X X X X

X X


TABLE 12-Continued

199

Core

Stage 6/5 transition Stage 5elSd transition

Increase in Decrease in

@O (dec.) SST, SST,,, PO (inc.) SST, SST,,,

Y71-12-6 X X X X X X Y72-11-l X 1 1 X J 1

Note. X = no data base for assessing any offsets or there was no lead-lag determination made from the observed data; t and J = sense of dislocation of a’s0 and SST curves in cores; B = oxygen isotopic curve based on benthonic foraminifera; F = SST curve based on foraminifera; C = SST curve based on coccoliths.

large-scale translational offsets by the use of multiple species. Each species tends to have a different absolute abundance pattern, with maximum concentrations in different parts of the record. In the Atlantic, Cibicides wuellerstorfi is usually most abundant within interglacial Substage 5e, whereas Uvigerina tends to occur mainly in the surrounding glacial Stage 6 and Sub- stage 5d. Translational offsets due to one species tend to be counterbalanced by op- posing offsets due to the other. In some cases, a mismatch occurs at the “splice” between two species (e.g., core V25-59, 330-332 cm, Fig. 38; V29-179,480-490 cm, Fig. 48).

The most likely translational offsets of al80 signals in our records (Table 12) are as follows: downward translation of the Stage 615 6l*O boundary in cores RCll-120 (Fig. 14), V28-127 (Fig. 43), V32-126 (Fig. 50), and V34-88 (Fig. 52); upward translation of the Stage 6/5 al80 boundary in cores RCl l- 210 (Fig. 15), RC13-205 (Fig. 19), V28-345 (Fig. 46), and V30-97 (Fig. 49); downward translation of the Stage 5e/5d 6180 boundary in cores RC13-205 (Fig. 19), V28- 14 (Fig. 41; planktonic 6i80 curve only), V28-127 (Fig. 43), and V30-97 (Fig. 49); and upward translation of the Stage 5e/5d 6i80 boundary in core RCI L-120 (Fig. 14).

The most likely translational offsets of SST signals (Table 12) are: downward translation of deglacial SST warming in cores RCl l-120 (Fig. 14; foraminiferal estimates),

RC12-294 (Fig. 17; foraminiferal estimates), RC12-339 (Fig. 18), TR126-29 (Fig. 26), V18-68 (Fig. 28), V22-108 (Fig. 33), V27-20 (Fig. 39), V28-14 (Fig. 41), V28-127 (Fig. 43), V29-179 (Fig. 48), and Y72-11-l (Fig. 54); upward translation of deglacial SST warming in cores E49-18 (Fig. 5), MD73025 (Fig. 9), V28-345, (Fig. 46), V30-97 (Fig. 49), and V32-126 (Fig. 50); downward translation of subsequent SST cooling in cores V18-68 (Fig. 28) and Y72-11-1 (Fig. 54); and upward translation of SST cooling in cores K708-1 (Fig. 7), RCll-120 (Fig. 14, foraminiferal estimates), RC12-294 (Fig. 17, foraminiferal estimates), RC12-339 (Fig. 18), RC13-205 (Fig. 19, foraminiferal estimates), V27-20 (Fig. 39), and V28-14 (Fig. 41).

In some cases, these possible translations of S’*O and SST curves are in the same direction and thus probably had little or no impact on the relative lead-lag relationships (e.g., cores RCll-120, V27-20, V28-127, and V30-97). In other cases, the likely translations were insignificantly small relative to the observed lead or lag (e.g., the Stage 5e/5d SST lag in cores K708-1 and V23-82).

Based on reasonable assessments of the possible amount of translational offset, we conclude that the following SST leads or lags could be largely or wholly explained by translational bioturbation offsets: the (foraminiferal) SST lead on the Stage 6/5 transition in South Atlantic core RC12-294;


roughly half of the SST lead on the Stage 6/5 transition in Indian Ocean core RC12- 339; the small SST lag on the Stage 6/5 transition in Gulf of Mexico core TR126-29; the SST lag on the Stage 5e/5d transition in Norwegian Sea core V28-14; and the SST lag on the Stage 5e/5d transition in Carib- bean core V28-127. On the other hand, the SST lead across the Stage 6/5 transition in Indian Ocean core V28-345 would be even larger if bioturbation effects were removed.

These results suggest that several of the low-latitude lead-lag patterns are equiv- ocal, particularly in the Atlantic Ocean (Figs. 58-60). Many of the observed SST leads in the equatorial South Atlantic and the observed SST lags in the equatorial North Atlantic are <2000 yr. We conclude that SST leads or lags relative to #*O in equatorial Atlantic cores are ambiguous because the observed SST/6180 offsets are well within the range of probable mixing effects.

Summary of Problems and Their Impacts

It is clear from the preceding that many of our initial conclusions based on observed @*O/SST trends are compromised by one or more of the complicating factors listed in Tables 9, 11, 12. Yet, in the aggregate, several patterns remain coherent.

Stage 5e SST reconstruction. The most striking observation about the last interglacial ocean remains its similarity to the modern ocean. The negative SST anomalies based on radiolaria in the subantarctic ocean and calculated by comparison with modern atlas temperatures (Fig. 56a, b) largely disappear if the comparison is made against core-top temperatures (which consistently underestimate atlas values; Table 4). Foraminiferal estimates also argue against negative anomalies (Figs. 10, 14).

The region of positive SST anomalies in the mid-latitude North Atlantic is harder to assess. Generally, the positive anomalies are within the standard error of estimate and some cores even show small negative anomalies. The general tendency toward positive SST anomalies leads us to conclude

that the evidence permits, but does not prove, a slightly warmer mid-latitude North Atlantic during the last interglaciation.

The 1.5”-2°C negative winter SST anomalies in the Gulf of Mexico cores (TR126- 23, -29) could be invalid only if our Stage Se picks were significantly in error. This is unlikely because benthonic foraminifera in this region are well protected from the meltwater incursions that probably affected surface waters (Emiliani et al., 1975; Kennett and Shackleton, 1975). However, the negative SST anomalies in the Caribbean (also noted by Prell and Hays, 1976) are probably invalid. One occurs in V12-122 where the last interglacial peak may have been missed due to its low sedimentation rate and wide sampling interval; the other anomaly (in V28-127) would disappear if we adjusted the Stage 5e pick upward or downward by a small amount.

Finally, the mixture of warm and cold SST anomalies in the South and equatorial Atlantic, particularly along the African coast, are shown to be suspect by no-analog SST estimates and by discordant multiple SST estimates. The implication is that circulation in these regions was significantly different from that today, but the nature of the difference is not clear. Up- welling is clearly implicated in the no-analog conditions in northern equatorial Atlantic core V22-196 and may be involved in the three South Atlantic cores with discordant multiple estimates (RCl l-86, RC13-228, and RC13-229).

Lead-lag relationships. The two major lead-lag trends in the last interglacial ocean are the tendency for a Southern Hemisphere SST lead and a mid-latitude North Atlantic SST lag.

The North Atlantic SST lag remains valid despite complicating factors. Bioturbation could not produce the 40-cm offset of the El80 and SST transitions at the 5e/5d boundary in cores K708-1 and V23-82. Changes in bottom-water temperature could complicate the al80 response, but to explain the full 1.1%0 change from Stage 5e to 5d would require an impossible 5°C de-


crease in bottom temperature. The trends in both cores show that SST values were still near full-interglacial levels when maximum Stage 5d 6180 values were reached. Even if only half of this 6i80 shift was ascribed to true ice-volume changes, it nev- ertheless occurred before a significant SST decrease took place.

There is clear evidence of downward bioturbational translation of the SST values in high-latitude North Atlantic cores at the Stage 6J5 boundary (Ruddiman er al., 1980a). This, and the scarcity of benthonic foraminifera for El80 analysis, has, however, probably had the effect of reducing an otherwise substantial lag of SST behind 6180 across the Stage 615 transition in cores K708-1 and V23-82.

In the subantarctic region, the 3000-yr SST lead based on radiolaria represents up to 50 cm of core length. No mixing effect could cause this large an offset; however, factors discussed in the section on Discor- dant Multiple Estimates leave in doubt whether this oceanic lead is actually a sea- surface or a subsurface phenomenon in the subantarctic ocean.

Several low-latitude cores show an SST lead at the Stage 6/5 boundary which per- sists despite bioturbational complications: Indian Ocean cores V28-345, RC12-339 and V34-88; equatorial Atlantic core V22-174; eastern South Pacific core Y71-6-12; equatorial Pacific core RClO-65; and eastern North Pacific core Y72-1 l-l. The SST leads at the 5e/5d boundary in these same cores are generally less convincingly developed and more within the range of bioturbational complications. In summary, some cores in the low latitudes of each ocean and others scattered across the world ocean appear to carry the basic Southern Hemispheric SST lead pattern at the Stage 6/5 transition.

OTHER EVIDENCE OF 6’80/SST PHASING

Our results imply that SST changes during the last interglaciation were time transgressive. This conclusion can be com-

pared against other oceanic evidence of two kinds: (1) lead-lag relationships across other major isotopic boundaries and (2) phase relationships determined from cross- correlation of long 6180/SST time series.

Evidence from Other Transitions

Climatic changes during the last deglaciation lie within the range of 14C dating and can be independently compared from region to region without referring to 6180 records. Hays et al. (1976) show a peak value of estimated SST dated at 9400 yr BP., in subantarctic core RCll-120, based on radiolarian estimates. The warming into this maximum began at an extrapolated age of 14,000 yr B.P., and a subsequent cooling occurred at an interpolated age of 6000 yr BP This positions the peak sea-surface or subsurface warmth of the subantarctic ocean during the present interglaciation within the earliest part of the Holocene and prior to the global ice-volume minimum reached after 6000 yr B.P. Evidence for a similar early Holocene SST spike occurs in South Atlantic cores RC13-229 and RC13- 228 (Morley and Robinson, 1980; Morley and Hays, 1981; Morley, unpublished data) and in the Ross Sea Ice Shelf region (Kel- logg and Truesdale, 1979).

In many parts of the mid-latitude North Atlantic, coldest SST values occurred just prior to 13,000 yr B.P. (Ruddiman et al., 1977). A substantial warming followed at 13,000 yr B.P., affecting a large part of the east-central North Atlantic (Ruddiman and McIntyre, 1981b, c). After a brief but intense oceanic cooling from 11,000 to 10,000 yr BP, a major warming at and after 10,000 yr B .P. reached most areas unaffected by the previous warming. The remaining cold areas- the Labrador and Greenland seas-then reached full interglacial warmth by 7000 to 6000 yr BP. Because the next glaciation has not yet begun, it is impossible to test whether the North Atlantic SST lags the future 6i80 volume increase.

It is, however, possible to compare the subantarctic and North Atlantic responses.


At the Stage 2/l boundary, the subantarctic warming indicated by radiolaria leads the North Atlantic response in most areas, but may be nearly synchronous with the eastern North Atlantic response off the coasts of Portugal, France, and Great Britain. More critically, the significant mid- Holocene cooling of the subantarctic prior to the 6’*0 and ice-volume increase leads the North Atlantic response, since no cooling has yet been observed. The foraminiferal response in the subantarctic is more nearly in phase with the North At- lantic changes.

At the Stage 514 isotopic boundary, Rud- diman et al. (1980b) found that the SST cooling lagged several thousand years behind the 6’*0 (-ice-volume) increase. The area thus affected encompassed most of the subpolar North Atlantic between 40” and 60”, including cores at the mouth of the Labrador Sea. This represents a substantially larger area than that which showed a SST lag at the last interglaciation (Figs. 58- 60). Ruddiman et al. (1980b) also noted that the amount of SST lag seemed to increase toward maximum values in cores at -4O”N.

These lead-lag results at other transitions are thus basically compatible with those discerned at the last interglacial interval.

Comparison with Long-Term Phase Relationships

Several studies of longer climatic records argue that the results found in the last interglacial interval also reflect significant long-term trends; these studies use the techniques of spectral analysis, frequency filtering, coherency analysis, and cross- correlation analysis to determine the mean phase relationships between the SST and 8*8O components across long lengths of record (alO yr). These phase relationships can be determined either directly for the complete downcore record of each component or for selected frequencies filtered from the original record because they show prominent spectral power.

Hays et al. (1976) examined SST and 6’*0 signals over the last 300,000 yr in the combined record of two subantarctic cores used here: RCll-120 and E49-18 (Fig. 1, Table 1). Their results show estimated SST leading al80 by 1000 to 2000 yr in various measures of the phase relationship. These results match in sign, and approach in magnitude, the circumantarctic SST lead relative to global ice volume found in this study (subject to the uncertainties of problems discussed in the section on Discordant Mul- tiple Estimates).

Ruddiman and McIntyre (1981a) found that SST in the 250,000-yr record of North Atlantic core V30-97 (Fig. 1, Table 1) lags behind 6i*O by 6000 2 1500 yr in the 23,000-yr cycle. In this case, the SST/Zi*O phase relationship was determined only for the 23,000-yr period because of its dominance in the SST power spectrum. A cross- correlation analysis between the total 8l*O and SST records in that core (unpublished data) gave the same SST lag, confirming the dominance of the 23,000-yr rhythm in the phase relationships. These results match in sign but exceed in magnitude the mid-latitude North Atlantic SST lag relative to global ice volume found in this study.

Moore et al. (1977) published cross-correlation analysis results for relatively short sections of two Pacific Ocean cores. In eastern equatorial Pacific core V19-29, SST lags #*O by 4000 ? 2000 yr over the last 80,000 yr of record. We also found a SST lag in V19-29 although somewhat smaller in magnitude. In eastern North Pacific core Y6910-2, Moore et al. (1977) found no phase offset between SST and S’*O over the last 90,000 yr. In this area, however, we found an SST lead rather than no phase offset. Neither of the two long-term phase relationships in the Pacific Ocean is necessarily definitive due to the short length of record analyzed. This did not permit clear definition of any orbital periodicities to which the phase leads and lags may be causally linked.

Imbrie has evidence of a prominent


11 ,OOO-yr SST period in an equatorial South Atlantic core not included in this study (V25-56 at 3”33’S, 35”14’W; see also BC et al. (1976)). An 11,500-yr rhythm is also evident in the SST record from subantarctic cores RCll-120 and E49-18 (Hays et al., 1976). Such a periodicity might explain the systematic response in some equatorial cores discussed earlier (section on Lead- Lag Relationships: SST and Ice): those cores with SST maxima on both the Stage 615 and Stage 5e/5d transitions. Because these transitions are separated in time by 11,500 yr, the SST maxima surrounding Stage 5e may be additional manifestations of periodic behavior at that frequency. Os- cillations at this high frequency may be particularly difficult to detect because of the low amplitude of SST changes at low latitudes, compounded by the effects of mixing in equatorial cores with low sedimentation rates. This pattern could thus be far more widespread than is evident in our data.

Ultimately, it would be desirable to obtain a long SST time series from dozens of cores across the world ocean and to define the phase relationship between local SST and global ice volume in each core for each of the orbital periodicities. This kind of coverage would fully characterize the phase response of the late Quaternary (and modern) ocean. In the meantime, our efforts here may point the way to some of the basic global phase relationships that will emerge: the prevalent Southern Hemi- sphere/Indian Ocean SST lead, the more lo- calized mid-latitude North Atlantic SST lag, and the more complex, possibly high- frequency response of SST signals in the equatorial oceans.

Implications of Diachronous Regional Responses

Our evidence from the last interglaciation suggesting that the ocean recorded major climatic changes at different times in different areas is thus borne out by comparison with other studies.

These results imply that continental rec-

ords could be subject to a chronostratigraphic constraint even more stringent than the basic goal of correlating with the oxygen isotopic stages or substages. It is ob- viously not sufficient in all areas to use the simple assumption that intervals of maximum local warmth are synchronous with minimum global ice volume; in some areas it will even be necessary to know which portion of an isotopic substage the continental section represents.

Some confirmation of this view is available from 14C-dated pollen records from the Holocene. The pollen record of New Zea- land, which has a maritime climate, indicates maximum warmth from 10,000 to 8000 yr B.P. (J. Salinger, unpublished data). Kellogg and Truesdale (1979) also inferred that the warmest Holocene temperatures in the region of the Ross Sea Ice Shelf ended about 8000 yr ago. This interval is synchronous with the early Holocene SST maximum in the subantarctic at 9400 yr B.P. (Hays et al., 1976; Hays, 1978). Pollen spectra in southern Chile also place the maximum warmth at - 11,000 and 9000 yr B.P. (Heusser and Streeter, 1980). On Ker- guelen Island, the maximum Holocene warmth ended by 5000 yr B.P. (Young and Schofield, 1973).

In the Northern Hemisphere, the clas- sical view places the maximum mid-latitude Holocene warmth of the “climatic op- timum” at 7000 to 6000 yr B.P. (Deevey and Flint, 1957; Iversen, 1973; Davis et al., 1980), with, at most, minor subsequent cooling in some areas. This picture is, however, subject to marked geographic variations. In the higher latitudes of central Northern America, the maximum northward extent of continuous forest was not reached until 5500 yr B.P. and was short- lived (Nichols, 1967). In coastal Labrador- Ungava to the east, the maximum northward advance of the boreal forest did not occur until 4000 to 3000 yr B.P., despite the disappearance of ice from the area several thousand years earlier (Short and Nichols, 1977).


In contrast, evidence from beetle remains suggests that some parts of Great Britain experienced summer temperatures warmer than today between 13,000 and 12,000 yr B.P. (Coope, 1970). Areas of Si- beria appear to have reached the maximum warmth at 9000 to 8000 yr B.P. (Khotinskii, 1977). These data show almost as much variation in response within the Northern Hemisphere as that between the Northern and Southern Hemispheres, probably in large part because the climatic effects of the late-lingering ice sheets varied geographically.

The Holocene thus displays a diachronous pattern of maximum continental warmth on land. Our data suggest that the last interglaciation, and hence previous interglaciations as well, should also have this attribute.

This constraint also applies to paleoceanographic studies. Several papers have attempted regional reconstructions of the last interglacial ocean based on relatively limited or no oxygen isotopic control (Rud- diman and McIntyre, 1976; Giraudet et al., 1976; Alvinerie et al., 1978; Kellogg, 1980; Crowley, 1981). Most of these studies shared a common, but unproven, under- lying assumption: that the warmest SST values at the approximate level of the Stage 5e region marked the interglacial ice- volume minimum. Our results show that this assumed synchroneity does not always hold and that maximum SST values lie near, but not necessarily on, the 6t80 minimum.

A heavy burden of chronostratigraphic proof is thus placed on any scientist who attempts a synoptic reconstruction of the last interglaciation or of any other level.

Figure 62 summarizes the results of the @*O/SST lead-lag study from the last interglaciation. It distinguishes between areas in which cores show clear evidence of leads and lags, including the confirmation provided by phase relationships, and those with somewhat uncertain, or geographically more scattered, lead-lag rela-

tionships. The designation “possible” also includes regions which may show a lead or lag on only one of the two isotopic transitions.

The clearest oceanic lead (surface or subsurface) is associated with the subantarctic ocean (4W5O”S), but much of the Southern Hemisphere core coverage displays this trend. The clearest SST lag is centered in the subpolar North Atlantic (40”-SOON), but much of the Atlantic north of 20” favors this trend. The equatorial trend is mixed, with many cores showing a SST lead on the Stage 6/5 transition, and a smaller number showing a SST lag on the Stage 5e/5d transition.

The reasons for these geographic patterns of leads and lags are not yet entirely clear. The subantarctic pattern may be linked to an inherently fast response of the very large high-albedo surface of sea ice around Antarctica to orbital forcing (Hays et al., 1976; Hays, 1978). Alternatively, the radiolarian lead could reflect changes in the subtropical gyre, with rapid propagation of these changes to higher latitudes of the subantarctic at subsurface depths (Howard and Prell, 1984). The delayed response of the subpolar North Atlantic may be due to the large thermal inertia of Northern Hemi- sphere ice sheets and to the substantial impact of icebergs and meltwater on the North Atlantic (Ruddiman et al., 1980a, b; Ruddiman and McIntyre, 1979, 1981a-c). Explanation of the various equatorial responses (high-frequency and other) must await detailed downcore studies of long time series in each area.

ADDITIONAL EVIDENCE OF LAST INTERGLACIAL CLIMATE

The diachronous timing of prominent climatic transitions and of maximum interglacial SST values during the last two interglaciations underscores the difficulty in linking our oceanic results from the Stage 5e level in the deep-sea record to records on land. In this review of other evidence from the last interglaciation, we accord-


FIG. 62. Final assessment of oceanic regions showing SST lead, lag, or both, relative to 6”O transitions. “Persistent” lead or lag means occurrence on the Stage 615 and 5e/5d transitions, in cross- correlation analysis, and in long-term phasing (large circles). “Possible” lead or lag means occurrence on only one transition and/or uncertain status due to factors complicating climatic records.

ingly restrict our attention to continental sequences with firm radiometric dating or with direct correlations into marine records by means of oxygen isotopic control.

Extent of Last Interglacial Ice Sheets

Most direct evidence for the extent of last interglacial ice sheets has either been eroded or now lies buried beneath younger sediments or present-day ice sheets. Infer- ences about last interglacial ice sheets come largely from indirect evidence.

There is a wide, but not universal, consensus that less ice existed during the last interglaciation than exists today. This consensus derives primarily from the evidence of coral-reef terraces, dated by uranium series and lying above modern sea level on

islands considered to be tectonically stable and on continental margins or tectonically unstable regions for which uplift histories are inferred. Across a wide range of geographic areas and tectonic environments, there is recurrent evidence of eustatic sea levels roughly 6 m higher than today (see references mentioned in section on Defini- tion of the Last Interglaciation; and summarized in Mercer (1968) and Moore (1982)). Estimates of the actual level vary within the range 2 to 7 m, in accord with models predicting regional variations in the mean sea-level difference (Walcott, 1972; Clark et al., 1978).

Isotopic evidence bearing on the question of relative ice volume at Stage 5e versus today is ambiguous. Detection of a


6-m sea-level difference in oxygen isotopic records from deep-sea cores is effectively impossible; 6 m would equate to an isotopic shift of only 0.066%0 by the calibration used in Fairbanks and Matthews (1978). In contrast, the standard error on replicate analyses at a given stratigraphic level is usually about +0.10%0. In cores with the highest sedimentation rates and most detailed S180 records, many signals show slightly lighter isotopic values in Stage 1 than 5e (V19-28 and V19-29 in Ninkovich and Shackleton (1975), MD73025 in Duplessy et al., (1980), V19-30 from Shackleton (unpublished data)), while others show slightly lighter values in stage 5e (M12392-1 in Shackleton (1977), RCll-120 in Hays et al. (1976). Gen- erally, the observed isotopic differences between the two levels lie within the standard analytical error. They could also be ascribed to temperature differences of <l”C between the Substage 5e and Stage 1 levels. Thus, the strongest evidence for an ice- volume difference at isotopic Substage 5e is from dated coral reefs.

Six meters of global sea-level equivalent is roughly the volume of ice contained in each of two ice sheets today: the Greenland Ice Sheet and the West Antarctic Ice Sheet. Previous explanations of the higher sea level have suggested essentially complete disappearance of either the Greenland Ice Sheet (Emiliani, 1969) or the West Ant- arctic Ice Sheet (Mercer, 1968). A third possibility is partial disappearance of one (Emiliani, 1969) or both ice bodies, and/or the much larger East Antarctic Ice Sheet (Denton et al., 1971).

Assessing the stability of the Greenland Ice Sheet during terminations and associated warm peaks is a difficult problem without a clear answer. Either of two extremes still remains as possibilities for the last interglacial. The ice sheet may have melted significantly or even disappeared entirely due to precession-induced warmth (Emiliani, 1969). Alternatively, it may have been larger than now due to increased precipitation. We do not even have an answer

to the fundamental background problem of why the Greenland Ice Sheet survived nearly intact during the Stage 2/l deglaciation. The margins receded up to several hundred kilometers, as shown by Weidick (1976), and north Greenland ice elevations dropped about 800 m, as deduced by Ray- naud and Lebel(1979), while all other large Northern Hemisphere ice bodies, including the adjacent Laurentide Ice Sheet, suc- cumbed completely.

Hughes et al. (1977) and Denton and Hughes (1981b) inferred that marine disintegration mechanisms accomplished much of the Stage 2/l deglaciation. They suggested that the Greenland Ice Sheet survived the most recent deglaciation primarily because it was unique among Northern Hemisphere ice sheets in being largely terrestrial and hemmed in by coastal uplands. Thus it was immune to the marine disintegration mechanisms that were instru- mental in destroying other Northern Hemi- sphere ice sheets. In this view, the Green- land Ice Sheet was the most stable of the late Quaternary Northern Hemisphere ice domes. And it is unlikely to have collapsed by marine disintegration mechanisms during previous deglaciations, including the Stage 615 boundary.

Other views emphasize ablation by the sun, and by warm air; these suggest that prolonged interglacial warmth higher than the Holocene peak would be the most likely cause of significant ice-surface lowering. Ablation rates faster than during the Stage 2/l deglaciation could also lead to increased shrinkage of the Greenland Ice Sheet during earlier deglaciations.

Ambach (1980) calculated that a climatic warming of 1.5”-3.O”C could cause a negative mass balance for the present Green- land Ice Sheet which, if maintained for a sufficiently long time, could result in a lowering of surface elevation. His analysis stresses ablation and accumulation area changes due to temperature variations, but ignores possible variations in cloudiness, precipitation, and flow dynamics. Am-


bath’s calculation suggests that the mass balance of the Greenland Ice Sheet is sen- sitive to further climatic warming. How- ever, the results of Funder and Hjort (1973) from east Greenland imply that increased ablation might be counterbalanced by enhanced precipitation during phases of the deglaciation hemicycle when the North At- lantic was ice free. Flohn (1977) cautioned that sea-ice reduction associated with climatic warming might allow increased precipitation over Greenland. Conceivably, this could counterbalance summer melting and allow the ice sheet to remain stable or even expand.

In addition, air temperatures warm enough to cause surface melting on the Greenland Ice Sheet need not lead to a large decrease in ice mass, because much of the meltwater percolates into the firn and refreezes to become a broad band of superimposed ice that lies upslope from the true ablation zone (Bader, 1961; Benson, 1962). This may invalidate the assertion by Amback (1980) that the Greenland Ice Sheet would acquire a strongly negative mass balance with only moderate climatic warming. The most complete treatment of this problem, both in terms of atmospheric circulation associated with Greenland and ice-sheet dynamics, is by Radok et al. (1982). They concluded that, during the present interglaciation, the Greenland Ice Sheet does not have a mass balance significantly different from zero.

Questions concerning the effect of atmospheric warming on the Greenland Ice Sheet must consider its central and southern domes separately. The dilemma is simply stated. Snow precipitation over the ice sheet originates as evaporation from open seas. The more distant these seas, the greater the amount of precipitation that occurs before convective storm systems reach the ice sheet. A cover of floating ice (sea ice or ice shelves) keeps these evaporation sources at a distance.

Today, the low southern dome of the Greenland Ice Sheet has no sea-ice fringe

in the summer and has a fringe averaging about 100 km wide in the winter. Climatic warming would have virtually no effect in reducing the distance to source areas. We might then expect that the major conse- quence of climatic warming would be to en- hance ablation associated with the southern dome, both in terms of the rate and area. With no compensating increase in snowfall, the southern dome might then vanish completely during a prolonged interglaciation.

A quite different response may apply to the high central dome of the Greenland Ice Sheet. Today, along its eastern flank, a summer sea-ice fringe at least 100 km wide extends southward to 70”N latitude in the Greenland Sea and expands to some 500 km in width during the winter. Along its western flank, Baffin Bay is filled with sea ice in the winter; along its northern flank, sea ice covers the Arctic Ocean throughout the year. Climatic warming that removed any substantial part of this sea-ice cover would put sources of snowfall, especially winter sources, much closer to the central dome. The central dome of the Greenland Ice Sheet might then expand as a result of climatic warming. Expansion would be particularly vigorous toward the north, and might well cause a northward displacement of the high central dome during a prolonged interglaciation.

Sea level would increase by only 1 m if climatic warming during the Stage 5e level melted the low southern dome of the Greenland Ice Sheet. Increased ablation around the periphery of the high central dome might raise sea level an additional 1 m, but increased accumulation over this dome might lower sea level by 1 m or more. The net effect of prolonged interglacial climatic warming on sea level cannot be pre- dicted for the Greenland Ice Sheet. It could range from no change to a drop of 1 to 2 m, depending on the yet unknown interactions between the ice sheet, the atmosphere, and sea ice.

The surface elevation of the Greenland Ice Sheet during Substage 5e should even-


tually be determined from total gas content and 6180 values in deep ice cores from central Greenland. Also, improved modeling of factors affecting mass balance should indicate whether warming greater than that attained during the Holocene could cause significant ice-sheet shrinkage.

For the present, we think that plausible arguments can be made for increased ablation of the Greenland Ice Sheet during the Stage Se level due to higher summer insolation values and increased air temperatures. The ice sheet was susceptible to greater ablation due to the apparently increased warmth of the Northern Hemi- sphere during the Stage 5e level (sections on Correlation to Pollen Records, and Other Terrestrial Sequences). Also, the summer insolation maximum leading into the Stage 5e level was substantially stronger over the Greenland Ice Sheet than that leading into the present interglaciation (Berger, 1978). It is likely that these factors led to a somewhat smaller Greenland Ice Sheet during the Stage 5e level than during the Stage 1 level, because of increased ablation during both the Stage 6/5 deglaciation and the Substage Se interglaciation. We do not think that this increased ablation led to the demise of the Greenland Ice Sheet during Substage 5e because apparently ice of this age still exists at the base of the ice sheet in northern Greenland (Dansgaard et al., 1971). However, it is quite possible that the difference in ice-sheet volumes at the last two interglaciations could translate into a 1- or 2-m change in eustatic sea level. This could account for part or all of the positive eustatic sea-level range, 2-7 m, inferred from coral reefs.

In the Antarctic, Mercer (1968, 1979) suggested that a summer temperature increase of 5”-10°C in high, southern latitudes would cause recession or disappearance of the huge Ronne and Ross ice shelves in the Weddell Sea and Ross Sea embayments of Antarctica, leading to disintegration of the marine-based West Ant-

arctic Ice Sheet that is in dynamic equilibrium with these shelves. A temperature rise of this magnitude is assumed to have had little effect on the East Antarctic Ice Sheet, which is more terrestrial than the West Ant- arctic Ice Sheet and does not depend on fringing ice shelves for its existence (Mercer, 1968). Nor is a warming of this magnitude sufficient to form extensive peripheral ablation zones on the East Ant- arctic Ice Sheet (Denton et al., 1971). Thus the marine-based West Antarctic Ice Sheet is thought to be uniquely vulnerable to climatic warming through instabilities induced by removal of peripheral ice shelves, whereas the East Antarctic Ice Sheet is generally thought to be stable.

Mercer (1968) specifically suggested that a warming of 5”-10°C over present-day values caused the demise of the West Ant- arctic Ice Sheet during the Stage 6/5 deglaciation, leading to the higher-than-present sea level inferred at the Stage 5e level.

Recent modeling efforts (Thomas and Bentley, 1978; Stuiver et al., 1981; Fastook, 1981, 1984) and geologic field work (Stuiver et al., 1981) emphasize the sensitivity of the marine-based West Antarctic Ice Sheet to changes in both sea level and fringing ice shelves in the Ross Sea and Weddell Sea embayments.

Stuiver et al. (1981) emphasized the im- portance of ice streams, ice shelves, and calving bays in the disintegration mechanisms of marine ice sheets in West Antarc- tica during the Holocene and Stage 6/5 deglaciation. According to this model, ice- sheet recession in West Antarctica occurs when the buttressing effect of fringing ice shelves is reduced. Diminished buttressing occurs when (1) a sea-level rise lifts fringing ice shelves from their pinning points, causing ice thinning and recession and/or (2) climatic warming puts an ice-shelf surface into an ablation regime, causing the ice shelf to disintegrate by warming or by frac- turing along weak links coupling it to ice streams and bedrock. Removal of the large


West Antarctic ice shelves would cause surging of West Antarctic grounded ice (Hughes, 1975, 1977). Surging would result in downdraw of interior drainage basins and eventually lead to grounding line recession. Slow-mode disintegration would occur when the grounding line and calving front of an ice shelf retreated at comparable rates, so that a degree of ice-shelf buttressing could be retained and possibly increased, if isostatic rebound repinned the ice shelf. Fast-mode disintegration would occur when the calving front retreated faster than the grounding line, thereby removing the buttressing ice shelf. Isostatic rebound would have no effect in this case. Figure 63 shows the minimum Holocene West Antarctic Ice Sheet for the two disintegration modes. These extremes of deglaciation should also represent the extremes of Substage 5e deglaciation.

These modeling results afford support for Mercer’s (1968) hypothesis. However, significant uncertainty still exists as to whether ocean/air temperatures during the Stage 6/5 transition were warmer than during the Stage 2/l transition, as is needed to explain why West Antarctic ice would entirely disintegrate in the earlier deglaciation and not in the most recent one. In this paper we have concluded that the SST of the last interglacial ocean was in general similar to that in the present interglacial. In particular, Hays et al. (1976) estimated a SST maximum in subantarctic latitudes during the Stage 615 deglaciation that was not significantly warmer (0.5”C) than that at 9000 yr B.P. at the height of warmth during the Stage 2/l deglaciation. Without clear evidence of sustained excess warmth during the Stage 6/5 deglaciation, the jus- tification for a Stage 5e destruction of the West Antarctic Ice Sheet by Mercer’s (1968) hypothesis is not complete.

Bentley (1983) examined the mass balance of the West Antarctic Ice Sheet during the present interglaciation. He concluded that the ice sheet is, if anything, growing

and that climatic warming would not directly reverse this condition because surface meltwater would percolate into the firn to refreeze as superimposed ice. On the other hand, if climatic warming produced warmer ocean currents that caused rapid melting beneath the Ross and Ronne ice shelves, West Antarctic ice could be evac- uated in a matter of centuries, perhaps only 500 yr, by marine instability mechanisms. Gordon (1983) analyzed the oceanic response to climatic warming and concluded that rapid melting beneath Antarctic ice shelves was possible. Hughes (1982, 1983a, b) emphasized the possibility that surface meltwater produced by climatic warming would attack weak links of Antarctic ice shelves to the ice streams that feed them and to the ice rises that pin them, so that ice shelves might disappear by fragmentation instead of by uniform basal melting in situ. He argued that the fragmentation process would be much more efficient in removing ice shelves and proposed 200 yr as a minimal time for evacuating ice from West Antarctica.

In view of the fact that results from the Substage 5e subantarctic ocean do not clearly justify removing the huge ice shelves in the Ross Sea and Weddell Sea embayments, we suggest several alternate possibilities for disintegration of the West Antarctic Ice Sheet during the Stage 6/5 deglaciation.

Our first alternative follows a suggestion made by Hughes and Denton in 1975 (Hughes, 1982) that Pine Island Bay at the head of the Amundsen Sea now constitutes the most vulnerable part of the West Ant- arctic Ice Sheet. The Amundsen Sea forms a third large embayment in the West Ant- arctic Ice Sheet. Unlike the Ross Sea (Ross Ice Shelf) and Weddell Sea (Ronne and Filchner ice shelves), however, the Amund- sen Sea embayment is not covered with an ice shelf. The Amundsen Sea sector of the West Antarctic Ice Sheet is mostly drained by the fast-moving Thwaites and Pine Is-


a

WlEO’E

FIG. 63. (a) West Antarctica after slow-mode collapse. Dotted areas denote floating ice shelves. East Antarctic ice elevations are shown in 500-m contour intervals. From the modeling results of Stuiver et al. (1981). (b) West Antarctica after fast-mode collapse. Dotted areas denote floating ice shelves. East Antarctic ice elevations are shown in 500-m contour intervals. From the modeling results of Stuiver et al. (1981).

land glaciers, two huge ice streams that inland migration of grounding lines. The ab- flow into Pine Island Bay and therefore are sence of bedrock sills assumed in the model not buttressed by a confined and pinned ice has been confirmed (Drewry, 1980; Crab- shelf. Inland from Pine Island Bay, the ice tree and Doake, 1982). Associated down- saddle separating flow into the Weddell and draw of interior ice divides would shrink ice Amundsen seas is the lowest in West Ant- drainage basins of ice streams feeding the arctica and is underlain by a deep subglacial huge Ross and Ronne ice shelves, causing trench. The modeling results of Stuiver et similar recession of their grounding lines. al. (1981) show that surges of the Pine Is- Thus the ice sheet could collapse because land and Thwaites glaciers would promote unstable stream flow into Pine Island Bay


b

Wl

FIG. 63-Continued.

draws down the more stable sheet flow in the interior and further destabilizes shelf flow along the Ross Sea and Weddell Sea margins. The modeling results from Stuiver et al. (1981) leave open the possibility that this process is now underway, triggered by late Holocene removal of an ice shelf in Pine Island Bay, and that it could lead to the demise of the West Antarctic Ice Sheet in the present interglacial without increased warming. These results suggest that collapse of the West Antarctic Ice Sheet during the Stage 6/5 deglaciation could have been initiated through Pine Island Bay

without climatic warming greater than that of the Holocene. The results are also consistent with a recent analysis of diatoms in two sediment cores taken on the outer continental shelf of the Amundsen Sea, beyond Pine Island Bay (Kellog et al., 1982). The conclusion by Crabtree and Doake (1982) that Pine Island Glacier has a positive mass balance at present has been challenged by Hughes (1983b).

The second alternative is that the sea- level rise during the Stage 6/5 deglaciation caused disintegration of the West Antarctic Ice Sheet. There is a widespread consensus


that Stage 6 ice sheets in the Northern Hemisphere (Illinoian in North America and War-the in northern Europe) were more extensive than those of Stage 2. Recent studies in the McMurdo Sound area of the Ross Sea indicate that the same was probably true for the West Antarctic Ice Sheet (Dagel et al., unpublished data). There- fore, isostatic depression may have been greater beneath the Stage 6 than the Stage 2 ice sheets, causing an increased fraction of Stage 6 ice sheets to be marine-based and therefore susceptible to marine disintegration mechanisms during the Stage 6/5 deglaciation. Ruddiman and McIntyre ( 198 1 b , c) and Ruddiman et al. (1980a) suggested that rapid decreases of Northern Hemi- sphere ice-sheet volumes occurred early in both the Stage 6/5 and the stage 2/l glacial/ interglacial transitions, largely by downdraw and calving. An even more abrupt and rapid sea-level rise at the beginning of the Stage 6/5 deglaciation, followed by continued sea-level rise of a greater amplitude than that during the Stage 2/l transition, could have caused more rapid recession of West Antarctic ice than the slower retreat documented by Stuiver et al. (1981) during the Stage 2/l deglaciation. Rapid Stage 6/5 deglaciation could have outstripped isostatic rebound so that grounding-line recession would not be slowed by repinning of fringing ice shelves, as was postulated in the Ross Sea during the slower Stage 2/l grounding-line recession (Thomas, 1976). Thus, it is possible that the speed and amplitude of the Stage 6/5 sea-level rise caused enhanced recession, or perhaps even caused the demise, of the marine-based West Antarctic Ice Sheet.

The third alternative is that increased ablation of the Greenland Ice Sheet during the Stage 6/5 deglaciation and/or at the Sub- stage 5e maximum contributed the excess sea-level rise that caused increased recession or demise of marine ice in West Ant- arctica. It seems unlikely, however, that a l- or 2-m increase in sea level from melting Greenland ice alone could have caused

widespread Antarctic deglaciation. It probably would have to have acted in concert with sea-level rise due to the concurrent recession of other Northern Hemisphere ice sheets.

In summary, there is somewhat ambiguous evidence for slightly smaller ice volumes in the Substage 5e interglaciation than in the present interglaciation. Small differences (of 2 m or less) could possibly be explained solely by increased ablation of the Greenland Ice Sheet by summer insolation values and temperatures in the lower atmosphere that were higher than those of the present interglacial. Larger differences (6 m or so) would require partial or complete destruction of marine ice in West Antarctica. This destruction could possibly be accomplished by several mechanisms that do not require prolonged warmth higher than that reached during the Stage 2/l deglaciation.

Correlation to Pollen Records

Convincing correlations to last interglacial pollen sequences can be made through pollen records in marine cores with isotopic control and through terrestrial records with adequate radiometric control.

Heusser and Shackleton (1979) pioneered the technique of directly linking pollen- bearing sediments on coastal margins with the deep-sea oxygen-isotope stratigraphy. Both records exist in core Y72-11-l located at 43”15’N off the coast of Oregon (a core included in this study, Fig. 54). The isotopic stratigraphy extends into Stage 6 to an age of about 150,000 yr B.P. Use of an oceanic core raises some question about the fidelity of the pollen record due to selective fluvio- marine transport processes, but a clear pollen signal emerges. Heusser and Shack- leton (1979) note that there are synchronous fluctuations in 6i8O and in temperature-related pollen; this implies that local (northwest USA) atmospheric temperatures react synchronously with global ice volume at the major climatic oscillations .

We have made no attempt to link isotopic


Substage 5e with many continental sections considered to be last interglacial in age and given local or regional names (for example, see Frenzel, 1973). Most prominent among these are the Eemian pollen sections of Eu- rope that have been commonly equated with isotopic Substage 5e by both paleo- ceanographers and palynologists since the work of Shackleton (1969).

Most such land sections are fragmentary in nature and lack the chronological support of direct radiometric dating; correlations are ultimately based on curve matching. Kukla (1977) criticized the clas- sical continental terminology and attempted to align several European land sections with the deep-sea isotopic stages. He proposed that Eemian sections in three of the best-known European areas actually correlate with three different isotopic interglacial stages (5e, 7, and 9); the original type section Eemian was proposed to be correlative with Stage 9. This debate (see also Bowen, 1979) suggests the unresolved status of correlating European and other land sections with the more continuously deposited deep-sea record.

Three pollen-bearing sections from the continents have been published that are significant because of (1) their long, probably continuous, records and (2) independent stratigraphic control from 14C or paleomagnetic analyses.

Van der Hammen et al. (1971) published a 120-m long pollen section from Mace- donia, Greece, and assigned it a time scale based on extrapolation from 14C dating near the core top, with allowance for compac- tion. Ruddiman and, McIntyre (1976) later attempted to correlate the pollen curve into North Atlantic fauna1 climatic oscillations in core K708-7. The assumed correlation between the North Atlantic fauna1 cycles and the aI80 cycles over most of the Brunhes interval represented in core K708- 7 was later borne out by Thierstein et al. (1977). Recently, extended coring beyond the limits of the original section in Mace- donia retrieved the Brunhes/Matuyama pa-

leomagnetic reversal at a depth consistent with this proposed correlation (Kukla, unpublished data, 1980). The local name for the interval correlative with the last interglacial in the Macedonian section is Pan- gaion (Van der Hammen et al., 1971).

Woillard (1978, 1979a) published results from several borings in the Grande Pile peat bog of northeast France. She proposed correlating the arboreal/nonarboreal pollen curve with several deep-sea climatic signals, including a SST curve from the adjacent northwest Atlantic (Sancetta et al., 1973). Although based at first largely on curve matching, this correlation of the last 140,000 yr of record has been recently up- held by 14C dating back to 70,000 yr B.P. (Woillard and Mook, 1982) and seems to be widely accepted. Woillard uses the Euro- pean name Eemian for the presumed pollen equivalent of isotopic Substage 5e.

Adam et al. (1981) published a 115-m long pollen section taken from Clear Lake in northern California. Numerous 14C dates covering the last 30,000 yr controlled the late-glacial, deglacial, and Holocene sections. The deeper record was correlated by curve matching with the deep-sea oxygen- isotope record. The record was estimated to penetrate the last 130,000 yr to slightly beyond a level (pollen zone I) correlated to isotopic Substage 5e and showing warm (oak) pollen percentages substantially higher than any reached in the Holocene.

None of these four pollen records (one marine, three terrestrial) has yet been subjected to transfer-function analysis for estimation of atmospheric temperatures or precipitation. This critical step has been hindered by the difficulty in obtaining a modern pollen calibration set that is free of the influence of disturbances from civili- zation, especially in Europe. In the western United States, the complex climatological interrelationships of temperature, precipitation, and altitude have slowed this kind of analysis.

All four records lie between 39”N and 47”N and have a common pollen compo-


nent that is typical of extreme interglacial warmth-oak (Quercus). The common pattern among these cores is for oak pollen (and related warm indicators) to be roughly equivalent in abundance during the last interglacial (isotopic 5e equivalent) to the values in the mid-Holocene. This implies that temperatures in the middle of the last interglacial were equivalent to, or slightly warmer than, those in the mid-Holocene. In the Clear Lake record, oak was considerably more abundant in the presumed last interglacial level than anywhere in the Ho- locene, possibly implying greater last interglacial warmth than today.

In several of the records, the percentages of oak and other warm species diminish toward the core top. This is generally ascribed to interference from human activi- ties rather than to climatic deterioration. In Europe, delayed migration of certain tree types during the Holocene also may have affected changes in pollen spectra. For these reasons, it is not possible to infer un- equivocally environmental differences between the last interglaciation and modem conditions in these four pollen records.

Other Terrestrial Sequences

Most terrestrial sequences are short, dis- continuous fragments from the long span of Quaternary history. This forces an almost complete reliance on U-series dates to es- tablish correlations with oxygen isotopic Substage 5e.

Gascoyne et al. (1981) published 12 U- series dates on speleothems from Victoria Cave in northern Great Britain. Eliminating 3 initial values at levels for which redeter- minations were made, the 9 dates averaged 124,000 yr B.P. Interstratified with the speleothems are numerous components of the distinctive “hippopotamus fauna” assigned to the Ipswichian interglacial and thought to be the British equivalent of the Eemian. In addition to hippopotamus, this fauna includes rhinocerus, lion, and ele- phant. It appears to be virtually absent on the European mainland. Gascoyne et al.

(198 1) draw no climatic inferences from these remarkable deposits, noting only that rare occurrences of this fauna in France may trace its dispersal route from the Med- iterranean area. It is associated elsewhere with fully interglacial pollen spectra indicating warmth equal to, or greater than, today.

Gaven et al. (1981) measured the age of a brackish lake in a now hyperarid region of southeastern Libya with U-series analyses. Numerous dates cluster around a mode at 130,000 yr B.P.; this suggests a moist climatic phase in North Africa centered on the Stage 6/5 transition.

Shallow Marine Sequences

Shallow-marine records are also fragmentary and require U-series dates to con- firm correlations.

Keen et al. (1981) dated a travertine from the Isle of Jersey in the English Channel at an age of 121,000 yr B.P. The travertine is interbedded with beach gravel containing a mollusk, Astralium rugosum (Linne), which has a present northern limit some 340 km to the south. Keen et al. (1981) infer that SST along the coast may thus have been some 3” to 4°C warmer than those today.

Mangerud et al. (1979, 1981) claim, with no supporting U-series dates, a correlation of a coastal sequence in southern Norway both with the European Eemian type section and with Stage 5e of the deep-sea oxygen isotopic stratigraphy. The first correlation is based on the generally similar pollen succession to that of the type section in the Netherlands, although some differences in pollen spectra are noted and ascribed to the latitudinal separation. The correlation to the deep sea is based on a unique correlation: (1) mollusks in the section show that Norwegian coastal waters were slightly warmer than today (- 1°C); (2) isotopically constrained records from Nor- wegian Sea cores suggest that the only can- didate for an ocean with temperatures comparable to, or warmer than, those today is


the upper portion of the isotopic Stage 5e level (Kellogg, 1976; Kellogg et al., 1978). Thus, two coastal records suggest an eastern North Atlantic warmer than the modern ocean.

Other Oceanic Evidence

Due mostly to the lack of tine-scale #*O control, previous studies in the North At- lantic are inconclusive about the last interglaciation. Ruddiman and McIntyre (1976) inferred a 3” to 4°C SST warming based on the northward displacement of the polar front relative to its modern position; however, the northern and western limits of the polar front in that study were very poorly constrained.

Giraudet et al. (1976) and Alvinerie et al. (1978) concluded from planktonic foraminiferal studies that during the last interglaciation the northeast Atlantic and Norwe gian Sea were essentially identical to modern conditions.

Kellogg (1980) found that winter SST values were generally lo-5°C cooler in Stage 5e than today, whereas summer SST values ranged anywhere from 4°C cooler to 3°C warmer. The increased summer warmth was concentrated in the Norwegian Current close to the Norwegian coast, whereas colder values in summer and winter occurred near the center of the Norwegian Sea at 70”N and in a second region northwest of the Faeroe Islands. Three Norwe- gian Sea cores with 6*80 control used by Kellogg are also included in this study (K- 11, V28-14, V28-56); we used, however, a different transfer function with less restricted surface-sediment coverage. We also selected different atlas SST values for comparison. Our choice of equation and atlas values generally reduced or eliminated the positive SST anomalies that Kellogg found for summer, but had no consistent effect upon the winter values, some remaining positive, some negative.

Crowley (1981) reconstructed last interglacial SST values for the eastern North At- lantic between 15” and 45”W. He concluded

that average annual temperatures were within 0.5”C of modern values north of 35”N but 1 .o” to 1.5”C warmer in a narrow band across 30” to 35”N. In some cores, Crowley selected the warmest SST values available from anywhere within the interval thought to correlate with isotopic Substage 5e as the last interglacial temperature. All five cores with SST anomalies >l”C fell into this category; this implies lesser (or no) positive anomalies at the exact Stage 5e level.

In general, all these studies tend to support the conclusion that the North Atlantic during the last interglacial was very similar in temperature to today’s ocean or possibly slightly warmer, with most differences no larger than the standard error of estimate of the equation. Positive anomalies in many cores may be due to the choice of maximum SST values at levels not precisely coincident with the 6t80 minimum observed (or inferred by correlation). Colder than modern values may be due largely to amplitude reduction of warm SST peaks by mixing in cores with slow sedimentation rates. Overall, the SST anomalies in these studies are neither larger nor geographically systematic.

Thompson (1981) estimated last interglacial temperatures in several cores from the western equatorial and North Pacific. The equatorial temperatures were not significantly warmer than those today; little temperature change occurred in these cores even during full glaciations. Thompson noted Stage 5e SST estimates warmer than today in three cores between 30”N and 40”N (RC13-17, V32-126, and V32-128), with cooler estimates in one core (V28- 304). Because of complications from dissolution and limited surface-sediment coverage, the transfer function equation used by Thompson has large standard errors of estimate (l.S’C in August, 3°C in Feb- ruary). These factors, and the mixed pattern of SST anomalies at the stage 5e level, lead to a conclusion similar to that in the North Atlantic: the data permit, but do not


conclusively prove, a last interglacial ocean slightly warmer than today.

There is a large body of work on cores collected off the west coast of Africa in the North Atlantic by the German research vessel Meteor. Several of these cores span an interval reaching to and beyond the last interglaciation and are stratigraphically controlled by al80 or by correlations uti- lizing percentage CaCO, or microfossil bio- zones (Diester-Haass et al., 1973; Thiede, 1977). Investigators in this region are still working out disagreements over the basic glacial vs interglacial sense of changes in factors such as wind-strength and moisture balance in the eolian source areas (Parkin and Shackleton, 1973; Sarnthein and Koopman, 1980; Diester-Haass, 1976). They have made no attempt to discriminate significant differences between interglaciations.

In the vicinity of core V22-196 at 13” 50’N with the no-analog fauna during the last interglaciation, it is not possible to ex- tract a clear oceanic upwelling signal from productivity indicators in cores because river nutrients may actually control the productivity changes (Diester-Haass, 1975). Parkin and Padgham (1975) suggest that wind strength was stronger in interglacials than in glacials at 8”N latitude. This implies that extremely strong winds in an interglaciation even more intense than today (i.e., Substage 5e) would promote more vigorous upwelling than found in the modern ocean.

Thunell and Williams (1982) investigated the last interglacial interval in the eastern Mediterranean. They found a thick sapro- pelic mud on the isotopic decrease into Stage 5 (see also Vergnaud-Grazzini et al., 1977), preceded by a rise in SST, and ac- companied by an increase in diatom productivity and suppression of foraminiferal productivity. Thunell and Williams argue against Nile runoff as a factor in the increased salinity-induced stratification and favor deglacial meltwater routed through the Aegean Sea as the primary cause. Other workers favor African pluvial maxima as

the cause of deglacial sapropels (Rossignol- Strick et al., 1982). This origin would accord with the U-series dating of moist lacustrine conditions in North Africa around 130,000 yr B.P. (Gaven et al., 1981).

Summary

The first-order conclusion about last interglacial climate is that it was extremely similar to the current climate. The total volume of ice was probably equal to or slightly less than today’s, with any decrease due to calving of West Antarctic ice and atmospheric ablation of Greenland ice.

SST were generally similar but may have been slightly warmer in the North Atlantic and North Pacific. There is some evidence that marine temperatures along the Euro- pean coast and atmospheric temperatures over the European continent were significantly warmer than those today at some point in the Stage 5e interval.

It might seem contradictory that the last interglacial ocean shows so little change from today, whereas there is abundant (although mostly undated) evidence for markedly warmer summers during the last interglaciation. This difference may, however, be ascribed to the different climatic responses of land and water. Land has very little thermal inertia and responds very directly to climatic forcing. The higher summer insolation levels at 122,000 yr B.P. (Berger, 1978) would cause increased summer heating of the continental interiors (e.g., Kutzbach, 1981). Reduced winter insolation during the last interglaciation would presumably lead to cooler winters on the continents as well, but the many kinds of mid-latitude vegetation critically sensi- tive to summer warmth would only reflect summer-season conditions, In contrast, the oceans have a much higher thermal inertia and can to a larger extent integrate seasonal (and annual) changes. Thus the relatively similar condition of the last interglacial oceans may reflect the fact that mean annual insolation levels were very similar to those today.


The maximum warmth in and near Eu- rope was apparently not centered on the middle of the Stage 5e interval. Data in Kel- logg et al. (1978) show maximum oceanic warmth only in the upper half of Stage 5e in the Norwegian Sea. Ruddiman and McIntyre (1979) and our results here emphasize the late-lingering corridor of warmth in the eastern North Atlantic on the isotopic 5e/5d transition. Mangerud et al. (1979) found that sea level had begun to fall along the Norwegian coast before air or sea temperatures decreased significantly, implying ice buildup in North America or Ant- arctica during lingering warmth in Europe. Drawing on evidence published by Woillard (1979b), Kukla (1980) concluded that the rapid spread of the boreal forest coincided with the beginning of ice-volume increase on the Stage 5e/5d transition, but that the still-faster transition to taiga coincided with the significant oceanic and atmospheric cooling in the European area.

Farther south, in the Mediterranean and North Africa, there is corroborative evidence both on land and in the marine record of a major pluvial episode on the Stage 6/5 boundary.

CONCLUSIONS

(1) The last interglaciation, centered at -122,000 yr B.P., was the last time when global ice volume was as small as, or smaller than, that today. Any decrease in ice volume relative to today, apparently in the range of 2-7 m, probably represents minor ablation of the Greenland Ice Sheet by increased summer insolation and higher air temperatures and major disintegration of the West Antarctic Ice Sheet by calving.

(2) SST at the ice-volume minimum were little different from those today, with suggestive but inconsistent evidence of a slightly warmer North Atlantic and North Pacific and a cooler Gulf of Mexico. U- Series dates on shallow marine and land sequences on the European margin of the North Atlantic provide stronger evidence

of warmer conditions during the last interglaciation.

(3) The surface and/or shallow subsurface part of the Southern Hemisphere ocean generally led the global ice-volume response into and out of the last interglaciation; the mid-latitude North Atlantic lagged. As a result, different parts of the world ocean, and presumably of adjacent maritime portions of the continents, regis- tered the full warmth of the last interglaciation at times differing by as much as several thousand years. The ultimate quantification of these complex lead-lag relationships should be done in frequency-specific terms, with the relative phasing tied to specific (probably orbital) frequencies.

(4) A number of complicating factors limit the conclusions we can draw from these data. The oxygen isotopic records clearly contain local temperature (and other) factors in addition to the dominant ice-volume signal; however, the agreement in relative timing between planktonic and benthonic foraminiferal al80 curves for most cores in which both were analyzed argues that local temperature effects do not markedly alter the phasing of the global ice- volume signal in 6180 records.

(5) Discordancies among the estimates of local SST by various biotic groups probably reflect inadequate knowledge of the basic ecologic preferences (particularly depth and season) of the component species. One danger is that inaccurate sea-surface estimates may be made based on biota living below, and possibly decoupled from, the sea surface, although these may still reg- ister deeper changes adequately. Changes in the seasonal range of oceanic change may also bias the SST estimates.

ACKNOWLEDGMENTS The last interglacial study was supported by National

Science Foundation grants funded jointly by the In- ternational Decade of Ocean Exploration and the Cli- mate Dynamics Research Section to Brown Univer- sity, OCE 77-22888; Lamont-Doherty Geological Observatory of Columbia University (LDGO), OCE 77-22893; University of Maine, OCE 77-23253; and the


University of Rhode Island, OCE 77-23472. We thank the core curators at Florida State University, LDGO (currently maintained under NSF Grant OCE 81- 22083), Oregon State University, and the University of Rhode Island for their cooperation and aid. All of our CLIMAP cooperative researchers, who have will- ingly and freely discussed with us the many problems involved in the production of this work, are gratefully acknowledged. We thank B. Taylor and N. Katz for drafting the numerous figures; and J. Guadagnini for her perseverance in processing the ever-revised drafts. And to Ann Esmay, who kept tight and accurate control of the raw data resulting in the tables and microfiche, goes our loving gratitude. This manuscript could not have been assembled or published without her. LDGO Contribution No. 3583.

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The last interglacial ocean