Rapid climatic variability of the North Atlantic Ocean and global climate: a focus of the IMAGES program

Quaternary Science Reviews 19 (2000) 227}241

Rapid climatic variability of the North Atlantic Ocean and global climate: a focus of the IMAGES program Elsa Cortijo!,*, Laurent Labeyrie!,", Mary Elliot!, Estelle Balbon!, Nadine Tisnerat! !Laboratoire des Sciences du Climat et de l'Environnement, CNRS/CEA, Domaine du CNRS, 91198 Gif-sur-Yvette cedex, France "De& partement des Sciences de la Terre, Universite& d'Orsay, 91405 Orsay cedex, France

Abstract In the mid-latitudes of the North Atlantic Ocean, six large iceberg surges, called Heinrich events, have been recognised between 60 and 10 kyr BP. They are characterised by meltwater events associated with iceberg discharges from both the Laurentide and the Fennoscandian-Greenland ice sheets. The hydrological conditions associated with these surges show changes in sea surface temperature (2}63C drop) and in sea surface salinity (1}3& decrease). Carbon isotopic analyses tend to show that the thermohaline circulation is a!ected by such discharges with a signi"cant decrease in the ventilation of deep waters. During the same period, ice cores record large climatic #uctuations, called Dansgaard-Oeschger events, which have been recognised in the North Atlantic Ocean and in the Norwegian Sea, and more recently, in various sedimentary environments world wide. Interglacial periods, with reduced continental ice sheets, allow us to investigate rapid climatic variability in the absence of large Northern Hemisphere ice masses. There are no major instabilities during the Eemian period but the transition from a full interglacial period into glacial time is abrupt, in less than 400 years. ( 1999 Elsevier Science Ltd. All rights reserved.

1. Introduction Among the largely unresolved issues in Earth Sciences are the origin and mechanisms of the rapid climatic variability. The study of this rapid climatic variability is one of the fundamental objectives of the IMAGES program (International Marine Global Change Studies), created under IGBP-PAGES. Rapid climatic changes are recorded in the marine sediments or in glacial archives of the last glacial period. They may also have occurred during the last interglacial period. The comparison of these two climatic states can help to understand the evolution of the climatic system on such short-time scales: what is the role of the continental ice sheets, are there thresholds, how are they controlled, how do internal and global variabilities compare 2 In this paper, we will address three main issues. First, we will present some of the methods used to reconstruct paleoclimate from ocean sediment cores and the di!erent ways to build chronostratigraphical frameworks. We will

* Corresponding author. E-mail address: [email protected] (E. Cortijo)

then draw up a non-exhaustive state of the art of the rapid climatic variability during the last glacial period. This period is characterised by a series of huge iceberg discharges recognised in the sediment cores and called the Heinrich events (Heinrich, 1988): what are the sources of the icebergs and what can we learn from the detrital fraction, what are the impacts of the iceberg discharges on the hydrology, the mechanisms involved and the global impact? Finally, we will focus on the climate variability during the last interglacial period, a period of minimum ice volume in the high latitudes. Cores used in this paper are summarised in Table 1 and Fig. 1.

2. Rapid climatic variability during the last glacial period The rapid climatic variability of the last glacial period, between 10 and 60 kyr BP, has been extensively studied during the past few years. The "rst discovery of millennial climatic variability was done in the d18O ice record of Greenland (Dansgaard et al., 1982). In the marine realm, several authors have shown that, in the North Atlantic Ocean, the last glacial period is punctuated by several levels particularly rich in detrital minerals (Heinrich, 1988; Pastouret et al., 1975). These levels,

0277-3791/99/$ - see front matter ( 1999 Elsevier Science Ltd. All rights reserved. PII: S 0 2 7 7 - 3 7 9 1 ( 9 9 ) 0 0 0 6 3 - 3

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Table 1 Latitude, longitude, depth and references of the di!erent cores used in this paper Core

Latitude

Longitude

Depth (m)

Data

Reference

ODP644

66340@N

04334@E

1227

d18O

SU90-24 ENAM93-21 HU75-42 SU90-33 90-013-013

62340@N 62344@N 62339@N 60334@N 58312@N

37322@W 03359@W 53354@W 22305@W 48322@W

2085 1020 2403 2400 3380

d18O d18O d18O d18O, temp. d18O

SU90-16 BOFS17K V23-23 NA87-22 V23-81 SU90-39 SU90-12 91-045-094 ODP 609 BOFS5K SU90-44 T88-9P

58313@N 58300@N 56305@N 55330@N 5432@N 52334@N 51352@N 50312@N 503N 50340}N 50306@ 48323@

45310@W 16329@W 44333@W 14342@W 1638@W 21356@W 39347@W 45341@W 243W 21352W 17354@W 25305@W

2100 1150 3292 2161 2393 3955 2950 3448 3900 2547 4255 3193

d18O temp. d18O d18O, d18O, d18O, d18O, d18O d18O temp. d18O d18O

SU90-11 SU90-08 CH69-K09 SU90-05 SU90-03

44304@N 43331@N 41345@N 41338@N 40330@N

40301@W 30324@W 47321@W 32315@W 32303@W

3645 3080 4090 3285 2213

d18O d18O, d18O, d18O, d18O,

(Fronval, Jansen, Bloemendal & Johnsen, 1995) This study (Rasmussen et al., 1996b) (Fillon and Duplessy, 1980) This study (Hillaire-Marcel, de Vernal, Bilodeau & Wu, 1994) This study (Maslin, Shackleton, P#aumann, 1995) (Mix and Fairbanks, 1985) (Rasmussen et al., 1996b) (Bond et al., 1993) This study This study (Hillaire-Marcel et al., 1994) (Bond et al., 1993) (Maslin et al., 1995) This study (van Kreveld, Knappertsbusch, Ottens, Ganssen & van Hinte 1996) This study This study This study This study (Chapman and Shackleton, 1998)

called Heinrich events after (Broecker et al., 1992) are attributed to major iceberg discharges from the Hudson strait and the Norwegian Sea into the North Atlantic Ocean (Andrews et al., 1994; Bond et al., 1992). Many questions are still pending about the causes and consequences of the millennial scale climatic changes, the relationship between the di!erent ice sheets of the northern hemisphere and the dynamics of the ocean and atmosphere systems.

2.1. Methods and chronostratigraphical framework The oxygen isotopic ratio of the carbonate of the foraminifera shells is a function of the growth temperature and the water isotopic ratio (Emiliani, 1961). Planktic and benthic oxygen isotopic compositions are used to monitor the hydrological changes of surface and deep water, respectively. Neogloboquadrina pachyderma left coiling (polar form) and Globigerina bulloides (subpolar form) are the most common planktic species used to follow the changes in the sea surface temperature and salinity (Duplessy et al., 1991). In subtropical waters, Globigerinoides ruber white variety is analysed. Among the benthic species, Cibicides wuellerstorx, Uvigerina peregrina and Oridorsalis tener are used to monitor changes

temp. temp. temp. temp.

temp. temp. temp. temp.

in d18O of the deep sea water. Only C. wuellerstorx species can be used to follow the changes in the isotopic composition of the carbon. In all the cases, the d18O values are given versus PDB after calibration to NBS19 (Coplen, 1988). The mean external reproducibility of powdered carbonate standards is $0.06& for oxygen. Sea surface temperatures (SST) can be estimated by di!erent methods, based on micropaleontological tracers (dino#agellates see for example de Vernal et al., 1994, or diatoms, see for example Pichon et al. (1987), or on biomarkers (Villanueva et al., 1998). The SST reconstructions presented here are based on planktic foraminifera counts. The modern analogue technique is applied on countings of at least 300 individuals, to estimate paleotemperatures by identifying the "ve most similar core top samples in the North Atlantic data base (615 core top samples between 0 and 803N, modi"ed from P#aumann, Duprat, Pujol and Labeyrie, (1996)). Summer and winter SSTs are then estimated by averaging the summer and winter SSTs associated with the most-similar core tops (Prell, 1985). Similarity between sample and core-top assemblages, using 32 planktic taxa, is calculated using the chord distance dissimilarity measure. Uncertainty in the SST reconstructions corresponds to the root mean square error of

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Fig. 1. Location of the cores used in this paper

the top 5 analogue temperatures. In each case, all of the 5 most similar core-top samples are valid modern analogues for the studied fossil sample (i.e. having a dissimilarity coe$cient (0.2). The error bars on temperature reconstructions are between 0.5 and 23C for the majority of the cores but may reach up to 3.53C in the subtropical cores because of the paucity of nearby core tops in the reference data base. Sea surface salinity (SSS) is estimated from paired SST and planktic d18O results following the method detailed by Duplessy et al. (1991). In order to estimate changes in SSS, the variations of the d18O of the sea water are calculated by using the measured planktic foraminiferal d18O and the reconstructed summer SST to solve the paleotemperature equation (Shackleton, 1974), taking into account the relationship between the growth temperature of the foraminifera and the summer SST. SSS is reconstructed from sea-water d18O using the global average d18O: S relationship for modern surface water of 0.5 : 1 (Craig and Gordon, 1965), or the glacial relationship 1 : 1. The changes in the global sea water d18O due to ice volume variations are reasonably estimated ($0.1& (Labeyrie et al., 1992)) for the last glacial period and can be removed from the sea water d18O signal. Salinity calculations are associated with many corrections and uncertainties for which the major ones are linked to the uncertainties in the SST reconstructions. Taking into account all the sources of errors (SSTs estimates, isotopic disequilibrium e!ect for 18O in calcite, global ice volume variations, calci"cation temperature, 18O-salinity relationship), the uncertainty on a given SSS

reconstruction will be of the order of magnitude of 0.5}1& depending on the slope chosen for the 18Osalinity relationship (with an average error bar in the SST estimate of $23C). For the last interglacial period, the interpretation of the SSS variations are restricted to the period of minimal ice volume because the global sea water d18O is not well constrained before and after the isotopic substage 5e. The study of the rapid climatic variability needs a common chronostratigraphic frame between the cores. Di!erent methods can be used to construct the age scales: f Radiocarbon datings by accelerated mass spectrometry (AMS) are used as often as possible to constrain the stratigraphic scale for the last 40 kyr. In these cases, the 14C ages are obtained on planktic foraminifera and corrected for a mean ventilation age of surface waters of 400 yr. This reservoir age correction can have been larger in the past particularly at high latitudes during cold stages. Estimation of this duration is still under discussion and a uniform correction of 400 yr is applied in the age scale of the cores presented here. The two ash layers, ash zones 1 and 2, identi"ed in the North Atlantic Ocean and dated to respectively 10.6 and 55 kyr are also used along with the 14C dates (Bard et al., 1994; Ruddiman and Glover, 1972) to build a polynomial or a linear regression between the datings, taking into account the error-bar for each dating point which increase from less than 100 yr for ages younger than 10 ka to over 1 ka for ages older than 40 kyr.

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f As the 14C dates are di$cult to obtain in all the cores, the age scale of the non 14C dated cores is obtained by correlating the benthic or planktic isotopic signals between cores. 2.2. Sources of the icebergs and synchroneity of the discharges The Heinrich events (HE) are very well imprinted in the sediment between 40 and 553N and easily recognisable with sedimentological and magnetic parameters (Fig. 2A and B). The magnetic susceptibility record, the water content and the coarse ('150 lm) detrital fraction (ice rafted detritus, IRD) provide a "rst tool to recognise these levels (Grousset et al., 1993). As the sediment is rich in continental minerals, and particularly in magnetite and titano-magnetites, the magnetic susceptibility is higher than in the surrounding carbonated sediment (Weeks et al., 1994). Furthermore, the grains transported by the icebergs are from di!erent grain sizes and the porosity of these levels is lower than in the foraminiferal-rich sediment (Cortijo, 1995). Nevertheless, all the HEs do not have the same imprint in the sediment. While HE1, HE2, HE4 and HE5 are very well identi"able by a sharp increase of the magnetic susceptibility signal, HE3 and HE6 have no clear signatures (Fig. 2B) within the majority of the sediment cores located in the `Ruddiman belta (Ruddiman, 1977). Where are the icebergs coming from? Why do all the events not have the same signature were some of the "rst question addressed within the scienti"c community. The sources of the icebergs can be constrained by the study of the detrital fraction. (Grousset et al., 1993) using the magnetic susceptibility signal and the neodymium/strontium isotopic analyses have determined a principal pattern from the Laurentide ice sheet for the major events, HE1, HE2, HE4 and HE5, and from the Norwegian sea for, HE3 and HE6 (Fig. 2C). The major role of the Laurentide ice sheet during HE1, HE2, HE4 and HE5 was con"rmed by (Gwiazda et al., 1996) using lead isotopic measurements. The separation of the IRD in di!erent mineralogical classes allows a "rst rapid estimation of their origin (Bond and Lotti, 1995) and can also help to determine synchroneity between sources. Detrital carbonate is Transported mainly from the Laurentide ice sheet, while volcanic ashes indicate a control from an Icelandic source (Bond and Lotti, 1995). These authors have shown using mineralogical studies that icebergs originating from Iceland and the Gulf of Saint Laurent (i.e. hematite stained grains and volcanic ashes) precede the icebergs originating from the Laurentide (i.e. detrital carbonate). These results have been interpreted as an indicator of the existence of precursor events of the major ice sheets instabilities. In a recent study Elliot et al. (1998) traced the pathway of icebergs containing dark volcanic ashes

Fig. 2. (A) magnetic susceptibility record in core SU90-08 (403N, 303W) (Grousset et al., 1993), (B) percentage of the lithic fraction in the same core, (C) sources of the icebergs as de"ned by (Grousset et al., 1993). The shaded bars in panels A and B underline the Heinrich events (noted H).

from Iceland and showed that ice sheet instabilities operate on millennium timescales from the Nordic regions. These precursor events seem to be part of a more rapid system of icebergs instabilities which operates within the Nordic regions. 2.3. Hydrological impact The iceberg discharges associated with the Heinrich events have an e!ect on the surface hydrology of the northern Atlantic. Each of the layer rich in ice-rafted

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debris in the 40}553N latitudinal band is associated with a lowering of the d18O of the planktic foraminiferal species Neogloboquadrina pachyderma left coiling. These isotopic minima correspond to an input meltwater that need to be quanti"ed to evaluate the impact of the iceberg discharge on the thermohaline circulation and then on the global climate. The hydrological impact of HE4 has been studied, using SST and SSS reconstructions from about 25 cores through the North Atlantic Ocean (Fig. 1) (Cortijo et al., 1997). Since the maximum lithic concentration is reached during the period of lowest SST, three time-slices around the HE4 have been determined: f 37 kyr BP corresponds to the period before the discharge, in the middle of isotopic stage 3 f 35 kyr BP is the maximum of the iceberg discharge and corresponds to the period of maximum IRD f 33 kyr BP corresponds to the end of the event and the return to `before the eventa conditions. The summer SSTs decrease by 2}63C during the iceberg discharge at 35 kyr BP and, at, return to values comparable to the ones before the discharge at 33 kyr (Fig. 3A and B). The SSSs estimated using the method described by (Duplessy et al., 1991) show a decrease by 1}3& during the iceberg discharge between 40 and 503N (Fig. 3C). Benthic oxygen and carbon isotopic studies have shown that the oceanic circulation was a!ected by such discharges (Vidal et al., 1997). This is very clearly shown in the benthic d13C record of core NA87-22 at 2100 m (Fig. 4). Fig. 5 summarises the surface hydrological evolution during Heinrich event 4. The air temperature variations in the Greenland ice core, as reconstructed by the d18O of the ice, exhibit a lot more oscillations than the records of the North Atlantic Ocean mid latitudes (Dansgaard et al., 1982). These events have been called Dansgaard-Oeschger events and show a pseudo-periodicity of 1.5}2 kyr. This has led to study more northern sediment cores in order to follow the evolution of the hydrology in the high latitudes and revealed high frequency changes of the sea surface hydrology (Elliot et al., 1998; Rasmussen et al., 1996a). The sedimentological and magnetical parameters have very di!erent signatures than in the mid-latitudes of the North Atlantic Ocean (Rasmussen et al., 1996b). The detrital fraction of the sediment is more abundant than farther south and the HE are identi"able by minima in the magnetic susceptibility signal (Fig. 6). The most remarkable characteristic of the northern magnetic susceptibility signal is the similarity with the ice d18O in GRIP (Fig. 6A and B). A correlation to the planktic d18O signal in the same core shows that each minimum in magnetic susceptibility is associated with a light peak in the d18O record. This magnetic signal is interpreted in terms of deep ocean circulation changes (Kissel, person. com.) induced by the

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melting of large quantities of icebergs. (Keigwin and Boyle, 1999) identi"ed repetitive changes in the d18O signal in the northern Sargasso sea related to stadialinterstadial changes in Greenland, pointing out to a large impact of both phenomenon.

2.4. Mechanisms The mechanisms invoked to explain these two kinds of variability are still uncertain. (Heinrich, 1988), based on an approximate time-scale (pseudo-periodicity of 10}11 kyr), suggested that these oscillations were linked to a half-precessional cycle and thus to insolation variations. But, as the time scale and the datation of the HE were improving, this hypothesis was rejected and glaciological models have been proposed. The HE have then been attributed to an internal oscillation of the Laurentide ice sheet (MacAyeal, 1993). Using a simple glaciological model, MacAyeal (1993) proposed that the Laurentide ice sheet could collapse with an internal periodicity of 7 kyr, close to the one observed in the data. But, this is not su$cient to explain the DansgaardOeschger events in the sediments nor why the small ice caps of the high latitudes are reacting "rst, and thus, seem to lead the big Laurentide debacle (Elliot et al., 1998). These results would be more in favour of an internal process of the climatic system, independent of any orbitally driven phenomenon. One possibility to better understand the role of the orbital processes in the rapid climatic variability is to study a glacial period under di!erent orbital forcing. Such a study was done in the subpolar North Atlantic and shows that such processes are inherent to the glacial periods of the Late Pleistocene, regardless of the orbital con"guration (Raymo, 1998). Oppo, McManus and Cullen (1998) have shown in isotopic stage 9, between 350 and 400 kyr, that HE-like events were present with a periodicity comparable to that of the last glacial cycle even though the precession cycle was of smaller amplitude.

2.5. Global impact The glacial rapid climatic events, best documented in Greenland and in the North Atlantic Ocean, have smoothed counterparts in Antarctica (Bender et al., 1994). The detailed record of atmospheric changes in oxygen isotopic ratio and in methane concentration of ice cores allows the study of the link between northern and southern hemisphere climates over this time period. Blunier et al. (1998) have shown that the atmospheric temperature over the Antarctic ice sheet was warming when coolings were recorded over the Greenland ice cores in the northern hemisphere, thus pointing to a complex relationship between both hemispheres.

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Fig. 3. (A) Summer SST reconstructions before Heinrich event 4 at 37 kyr BP, (B) Summer SST reconstructions during Heinrich event 4 at 35 kyr BP, (C) SST and SSS di!erences (during the event minus before the event) versus latitude. The major melting area between 40 and 503N is shaded. Salinity is calculated using a slope of 1 (glacial relationship) and a slope of 2 (modern relationship).

The HE have been extensively studied in the North Atlantic Ocean and studies are now showing that rapid climatic changes also occurred in other oceanic basins. There is now strong evidences that the rapid climatic variability was not restricted to the North Atlantic Ocean. In the North East Paci"c Ocean, (Behl and Kennett, 1996) have shown that the variations of the ocean oxygenation (and thus, of the ocean circulation) of the Santa Barbara basin correlates well with the Greenland ice-core

records. This may be explained by rapid changes in the location and intensity of the production of intermediate waters, in#uencing the ventilation of the Paci"c Ocean. Changes in deep or intermediate water production can be rapidly communicated to distant parts of the ocean. Similarly, the rapid climatic variability in the North Atlantic Ocean may in#uence changes in the low-latitude monsoonal variability. The variations of the total organic carbon concentration and of the planktic foraminifera oxygen isotopic composition in cores o!

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Fig. 4. (A) lithic fraction '150 lm in core NA87-22, (B) C. wuellerstor" d13C in the same core. Heinrich events have been underlined.

Pakistan show striking similarities with the d18 record of the d18O of the ice from the Greenland ice core (Schulz et al., 1998). All these results suggest that these changes, "rst observed in the Northern Hemisphere, may have been widespread around the world and that broad areas are sensitive to rapid changes in the ocean}atmosphere} cryosphere systems. The continents have then probably recorded climatic changes in accordance with these general climatic reorganisations. Although correlations of continental and oceanic paleoclimatic proxies are hampered by our ability to construct absolute age scales, temperature and/or precipitations changes have been observed in North America and Europe continental records. (Phillips et al., 1994) have shown that the Searles Lake in Southern California was experiencing low water levels during the last glacial period because of variations in the precipitations that could be related to the Dansgaard-Oeschger events in Greenland. On the other side of the american continent, the wet episodes of Lake Tulane in Florida are related to the HE (Grimm et al., 1993). Because the European continent is more sensitive to changes in the North Atlantic SSTs, the connection between the rapid climatic variability in North Atlantic

and changes in the European climate seems more direct. (Guiot et al., 1993) show by the study of continental records from La Grande Pile and Les Echets that the annual air temperature was experiencing a 2}53C decrease during the HE. Nevertheless, the comparison between marine and ice records still su!ers some basic problem of chronology. None of the results presented above is su$cient to "rmly establish that there is a phasing between a given signal (changes in oxygenation, concentration of organic matter 2) and the d18O of the ice in Greenland. The global and full understanding of the system lays on the construction of absolute time-scales between the di!erent kinds of records, marine sediments, continental sediments or ice.

3. Hydrologic changes during the last interglacial period Rapid climatic variability is quite easy to understand under glacial climate since the large continental ice sheets are susceptible to undergo important self-induced volume variations. Rapid climatic changes are less easy to understand during times of minimal ice sheet extension:

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Fig. 5. Summary of the hydrological changes recorded by Cortijo et al. (1997) for the Heinrich event 4 at about 35 ka. The light arrows show the iceberg discharges, the dark arrow the main path of the warm water current coming from the south. The 2& isoline is the limit of the maximum d18O anomaly during Heinrich event 4 iceberg discharge. The 83C isotherm is shown in the position before the iceberg discharge (at about 37 ka) and during the icebergs melting (at about 35 ka). The potential convection areas are located south and north of the maximum melting zone as de"ned by Vidal et al. (1997).

interglacial periods. Initial results from the deeper portions of the GRIP ice core, showing that the climate of Eemian was not uniformly warm as thought from CLIMAP Project Members, (1984) but experienced rapid climatic coolings (Dansgaard et al., 1993), stimulated renewed interest in the Last Interglacial period. Even if it is now well established that ice originally thought to correspond to the Eemian in GRIP has been disturbed by ice #ow (Fuchs and Leuenberger, 1996), some doubts persist about what happened exactly during the so-called Eemian part of the GRIP ice core (Johnsen et al., 1997). Furthermore, several subsequent studies in continental and marine environments have shown that Eemian climate was not as stable as previously suggested (Cortijo et al., 1994; Field et al., 1994; Fronval and Jansen, 1996; Seidenkrantz et al., 1995) and is punctuated by a rapid transition to the glacial period, less than 400 yr, as demonstrated by Adkins et al., (1997) using Thorium excess measurements to estimate the duration of the hydrographic change (Fig. 7). 3.1. Identixcation of isotope substage 5e All studied cores show the large decrease in benthic d18O associated with the transition between isotope stage 6 and substage 5e. However, the substage 5e}5d

transition is much more di$cult to identify with con"dence. When possible, the interval of minimum ice volume is de"ned within marine isotope 5e by identifying local isotopic minima at the beginning and end of the interval of low d18O values (events 5.53 and 5.51, respectively, of (Martinson et al., 1987)). This interval is referred as the `5e plateaua by SPECMAP. The temporal resolution of the cores presented here is larger than the resolution of the cores used in the SPECMAP stack and the structure of the signal is not exactly the same. Moreover, the SPECMAP stack was built using planktic foraminifera records. In some cases, these 5.53 and 5.51 events were not obvious. In these cases, the 5e plateau is de"ned based on the minimum d18O value plus or minus 0.2& (to account for natural variability during the interval of minimum ice volume). Sediment thicknesses for stratigraphic sections corresponding to the 5e plateau range from 24 to 192 cm. Using ages of 125.2}122.6 kyr for events 5.53 and 5.51, respectively, these thicknesses equal to a range in average sedimentation rate of 9}73 cm/kyr. In this orbitaly tuned age scale, the 5e plateau lasts around 3 kyr. This length is very short in comparison to other interglacials like the Holocene for example (Broecker, 1998). We choose then to use in this paper, the age scale proposed by Adkins et al. (1997) who give a duration for the 5e plateau of about 10 kyr.

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Fig. 6. From top to bottom panel: (A) d18O of the ice at GRIP site (Greenland) showing the Dansgaard-Oeschger oscillations (Dansgaard et al., 1993). The Heinrich events are underlined by grey rectangles, (B and C) Magnetic susceptibility and N. pachyderma s. d18O record for core ENAM93-21 (623N, 33W) from (Rasmussen et al., 1996a).

3.2. Results and discussion This `5e plateaua interval is de"ned by minimum benthic d18O values de"ning a plateau with little or no trend. An exception is core CH69-K9, where values increase by 0.2}0.3& during the interval (Fig. 8A). This trend is di$cult to explain, but may be due to a progressive change in deep water masses at this location during the ice volume minimum. The planktic d18O records in the di!erent cores display a variation of plus or minus 0.5& during the same interval, probably linked to changes in surface temperature and salinity values. However, G. bulloides d18O record in core CH69-K9 shows a peculiar positive event, with d18O values increasing by 1&, in the middle of isotopic substage 5e (Fig. 8B). Reconstructed summer SSTs follow di!erent time-dependent trends depending on the latitude of the core (Fig. 8C). Northern cores V27-60 and NA87-25 show a timedependent decrease, primarily within the second half of the substage 5e plateau, during which summer SST values decline by 2}43C. The southern cores CH69-K9 and JPC37 show a time-dependent increase during the same interval, with summer SST values rising by 1}43C. The error bar associated to the SSTs reconstructions are

of the order of magnitude of 0.5}23C and 1}33C in the low latitudes. These general trends in the SST and in the planktic d18O record, are overprinted with higher local variability, like in core CH69-K9. In most cores, the "rst half of the 5e plateau is separated from the second half by an abrupt event. This step like event is particularly clear on the SST records of the cores NA87-25 and CH69-K9, where it accounts for most of the general increasing or decreasing trends. The data therefore suggest, at least in some locations, a rapid change in the North Atlantic surface hydrology, superimposed on a more gradual trend. The observed temperature and salinity changes during the 5e isotopic plateau cannot be attributed to global ice volume changes, that are minimal or non-existent during this time interval. The insolation forcing has to be involved in some way (see Fig. 9). If the summer insolation decreases at all northern latitudes by about 4}10 W/m2, depending on the age scale used during this time period, the annually averaged insolation shows a slight decrease (4 W/m2) at high latitudes and a small increase (2 W/m2) at low latitudes. Even if these annually averaged changes are small in amplitude, it can be argued that, in the absence of ice-sheet changes, the upper ocean integrates the direct insolation forcing over one or several years,

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Fig. 7. Variations of the d18O of benthic foraK mifera, re#ectance of the sediments, Cd/Ca ratio, percentage of fragments, clay #uxes in core MD95-2036 (333N, 573W), from (Adkins et al., 1997).

and the corresponding temperature changes are therefore linked to annual insolation changes, not only seasonal ones. The associated salinity changes would then be caused by a direct evaporation}precipitation mechanism, because higher temperatures enhance evaporation while lower ones increase precipitations. The seasonal insolation changes may also have a net impact on the sea surface temperatures, through the associated changes in atmospheric circulation. At the beginning of substage 5e, the summer insolation in northern latitudes was indeed at an exceptionally high level, while the winter insolation was very low. The seasonallity was at a maximum and consequently, the thermal gradients between continents and sea were also higher. The atmospheric activity was therefore stronger, and the monsoon was particularly strong (De Noblet et al., 1996; Prell and Kutzbach, 1987). Low latitudes were thus experiencing stronger winds, which could lead to stronger upwellings in coastal areas and possibly also along the equator, and could also enhance the evaporation over the low latitude ocean. This enhanced atmospheric activity,

at the beginning of substage 5e, would therefore tend to cool at least some portions of the low latitude oceans. The following decrease in the seasonallity during the Eemian, by reducing the low latitude winds, could therefore induce a slight increase in low latitude SST. Actually, the same mechanisms were also at work during the Holocene period, and the SST reconstructions made at 6 kyr BP have shown that high latitudes in the North Atlantic were slightly warmer than today, while lower latitudes were cooler (Ruddiman & Mix, 1993), though the di!erences with the present climate are almost always within the error estimates. Simulations using atmospheric general circulation models with a slab ocean can compute SSTs by assuming that the ocean heat #ux, and thus the ocean circulation, remains unchanged. Experiments for 6 kyr BP (Liao et al., 1994) and 9 kyr BP (Mitchell et al., 1988) tend to present similar tendencies. The Holocene thus probably experienced a slight cooling over the ocean at high latitudes and a slight warming at lower latitudes. Our data suggest this was also the case for the Eemian, possibly with a larger amplitude.

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Fig. 8. (A) benthic d18O, (B) planktic d18O, (C) SST reconstructions using the modern analog method during the isotopic substage 5e and trend observed for the 5e plateau (Cortijo et al., 1999).

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Fig. 9. Insolation changes on annual mean, January and August average calculated for 103N and 653N. The dark shaded area is the 5e plateau as de"ned by Martinson et al. (1987), while the light shaded area is the 5e plateau as de"ned by Adkins et al. (1997).

A change in the strength of the Atlantic ocean northward heat and salt transport could also explain the observed trends in temperature and salinity. Isotopic stage 5e was at least as warm as the modern period and slightly warmer than today in some areas (White, 1993). Our data furthermore indicates that low latitude temperatures did increase throughout this time period. Several models have shown that the hydrological cycle intensi"es as the climate warms (Manabe and Stou!er, 1993; Weaver and Hughes, 1994). High SST resulting in strong evaporation in low latitudes will give increased precipitations in high latitudes (Arctic sea and surrounding continents) and then, this will induce small variations in the salinity and density of surface water at these latitudes. The increase in fresh water input will contribute to an enhanced formation of sea ice and an increased run-o! input in Arctic seas. The #ow of exported fresh water will then be larger and contribute to destabilise the

deep water formation. Model simulations indicate that a slight decrease in surface water salinity could have slackened (or even stopped) deep-water formation in the Norwegian Sea within a few decades (Bryan, 1986). The deep convection and the thermohaline circulation would then be a!ected as shown by (Oppo and Lehman, 1995), and a positive feedback will contribute to reduce the temperature and salinity of the Norwegian Sea. As the rate of deep water formation is reduced in the middle of substage 5e, the heat and salt not transported to the high latitudes contribute to further increase the SST gradient between low and high latitudes, acting as a positive feedback. As also pointed out by (Keigwin et al., 1994), changes in the balance between evaporation and precipitation over the high latitudes of the North Atlantic Ocean can be a good candidate to explain variations in deep-water production. This scenario, however, is not entirely supported by benthic 13C and Cd/Ca

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measurements made in cores MD95-2036 for no signi"cant variation is observed during the 5e plateau (Adkins et al., 1997). During the minimal ice volume time interval, the northern temperatures and salinities show an increasing trend, while the southern ones decrease. These changes are mainly gradual, but appear to lead to rather rapid changes in the middle of substage 5e, at least in some locations. Such changes in sea surface hydrology are somewhat similar to the changes already observed during the Holocene. They may be linked to changes in insolation and in atmospheric circulation. They are also consistent with a reduction of the North Atlantic thermohaline circulation, though more high-resolution deep-water records would be required to reach such a conclusion.

4. Conclusion All these studies show that abrupt climatic changes are inherent to the past climate. This points to the evidence of a threshold e!ect in the climatic oscillations linked to changes in the freshwater budget and heat transport a!ecting the ocean surface. During these last years, important improvements have been made in paleoceanography by increasing the time resolution of the studies. It is now very important to continue these e!orts by providing further quantitative studies, SST and SSS reconstructions as well as deep ocean circulation proxies. This will lead to develop more quantitative reconstructions to compare with the model results and thus help to understand the di!erent equilibriums of the thermohaline circulation in the North Atlantic Ocean.

Acknowledgements Thanks are due to two anonymous reviewers who greatly improved the manuscript with their review, and to Jess Adkins for his various comments. B. Le Coat and J. Tessier are thanked for processing of the isotopic analyses. This work was supported by CNRS, CEA, INSU (PNEDC) and EEC Environment Programme.

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