marine realm, the last interglacial period is defined by the isotopic stage 5.5, also called ... the isotopic composition of oxygen in the ice recorded at GRIP, which ...
6 GLACIAL AND INTERGLACIAL HYDROLOGICAL CHANGES IN THE NORTH ATLANTIC OCEAN Elsa Cortijol, Estelle BalbonI, Mary Elliot l , Laurent Labeyrie 1,2, and Jean-Louis Turon3 lLaboratoire des Sciences du Climat et de l'Environnement CNRS/CEA, Gif-sur-Yvette, France 2Departement des Sciences de la Terre Universite d'Orsay Orsay, France 3Departement de Geologie et d'Oceanographie Universite de Bordeaux I Talence, France
ABSTRACT The north Atlantic ocean is involved in the rapid climatic changes observed during glacial times. Indeed, many paleoclimatic indicators, from surface hydrological reconstructions, to deep water chemistry proxies, clearly indicate that major reorganizations of the Atlantic happened together with, and to some extent are probably responsible for the abrupt climatic shifts recorded in Greenland ice, but also in many marine and continental sites around the world. We have performed a detailed study of the last interglacial period, or Eemian, on several marine cores along a latitudinal transect in the North Atlantic ocean in order - to investigate the role of the Atlantic ocean during warm periods - to understand to what extent similar mechanisms were also at work at these times. Though the North Hemisphere ice volume was minimal and global ice volume constant throughout this time period, some trends in the sea surface temperatures and salinities are clearly recorded in the cores we have studied. These results are probably documenting the response of the ocean-atmosphere mostly to insolation changes, and represent an opportunity to better understand the inception of the last glaciation. In particular, the abrupt event recorded by Adkins et al. (1997) in the deep Atlantic, just at the Reconstructing Ocean History: A Window into the Future edited by Abrantes and Mix, Kluwer Academic I Plenum Publishers, New York, 1999.
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end of this period of minimal ice volume, illustrates that the role of Atlantic ocean reorganizations are also probably quite important for the understanding of warm period variability, as well as the variability of cold periods.
1. INTRODUCTION The stability of climate during the last interglacial period has attracted much attention during the past few years in order to better understand the natural variability of warm climatic intervals. This time period is also called Eemian on land. In the marine realm, the last interglacial period is defined by the isotopic stage 5.5, also called isotopic stage 5e. The boundaries are dated by (Martinson et aI., 1987) from the stages 6 to 5.5 transition at 128kyr to the 5.5 to 5.4 transition at 118kyr, though this timing is still somewhat controversial. Interest for this time period has emerged from the first results of the isotopic composition of oxygen in the ice recorded at GRIP, which showed that the climate of the Eemian was not uniformly warm (CLIMAP Project Members, 1984) but experienced rapid cooling events (Dansgaard et al., 1993). This initial evidence of rapid climatic oscillations led to a renewal of interest of Eemian climate. It is now well known that the GRIP record for this time period is disturbed by ice flowing (Fuchs and Leuenberger, 1996), but nevertheless, several studies in continental and marine environments have shown that the climate during the Eemian period was not as stable as was previously expected (Cortijo et al., 1994; Fronval and Jansen, 1996). The large shifts observed during glacial periods in the marine and ice records can probably be explained by switches in the operational mode of the Atlantic thermohaline circulation, caused by iceberg discharges from the surrounding large ice-sheets (Broecker, 1994; Vidal et al., 1997). Potential instabilities during interglacial periods, defined as periods of minimum continental ice volume, are therefore quite puzzling. Furthermore, several studies in continental and marine environments have shown contradictory results. Evidence of climatic instability has been found in Northwest Europe, Norwegian Sea and Labrador Sea (Field et aI., 1994; Fronval and Jansen, 1996; Seidenkrantz et al., 1995). In this paper, we will first present a summary description of the climatic variability recorded in marine sediments during glacial times, between approximately 60 and 10 kyr BP. Then, we will use the results of six sediment cores spanning late marine isotopic stage 6 (MIS 6) to MIS 5.4, from around 135 to 115 kyr BP, in order to follow the hydrological changes in the surface waters of the North Atlantic Ocean during the last interglacial period. We will thus follow in detail the inception of the last glacial period using different cores on a north-south transect in the North Atlantic Ocean.
2. VARIABILITY DURING THE LAST GLACIAL PERIOD It is now well established that the climate of the last glacial period, between 10 and 60kyr BP, was not stable in the North Atlantic region (Fig. 1). This period was punctuated by large and abrupt switches of the air temperature over Greenland (Dansgaard et al., 1993). Massive iceberg discharges over the North Atlantic ocean (Bond et al., 1992; Heinrich, 1988) were able to induce drastic changes of the hydrological conditions (Cortijo et al., 1997) and consequently·in the thermohaline circulation (Vidal et al., 1997). The large and rapid temperature fluctuations during the last glacial period were first observed in the Greenland ice records (Dansgaard et al., 1982). These oscillations have been found in all Greenland ice records, and the most detailed ones available now
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are from the GRIP site (Dansgaard et al., 1993; Johnsen et al., 1992) and the GISP2 site (Grootes et al., 1993). These so-called Dansgaard-Oeschger oscillations occur every 1.5 to 3kyr. But the most striking feature is the abruptness of the switches, since the climatic change occurs only within, at most, a few decades (Dansgaard et al., 1989). These events have always been suspected to be linked to changes in the Atlantic thermohaline circulation (Broecker et al., 1985). Marine records have a lower resolution, and the first records of glacial climatic variability were published somewhat later (Heinrich, 1988; Bond et al., 1992). Marine sediments from the North Atlantic between 40 0 N and 55°N have indeed several well identified layers of ice-rafted debris (IRD) which indicate episodes of massive iceberg discharges from the Laurentide and Fennoscandinavian ice sheets. These so-called Heinrich layers are therefore associated with a partial disruption of these two ice-sheets. The associated climatic events, usually called Heinrich events, are characterized by an intense decrease of the sea surface temperature, between 2 and 5°C, and a lowering of the surface salinity, from 1 to 3%0 (Cortijo et al., 1997). The impact on the deep ocean geochemistry, as recorded by l3C/ 12 C isotopic ratios (Vidal et aI., 1997), is very large. It clearly indicates a considerable slow-down of the North Atlantic thermohaline circulation during the meltwater event, which can explain in a large part the observed cooling. These Heinrich events occur approximately every 7 to 10 kyr. The sea surface temperature (SST), as recorded by the abundance of the polar planktic foraminifera Neogloboquadrina pachyderma left coiling in North Atlantic Ocean sediment (cores ODP609, V23-81) also exhibits the same large amplitude oscillations than those observed in the Greenland ice records (Bond et aI., 1993). Smaller and more frequent IRD layers are also found in some North Atlantic cores (Bond and Lotti, 1995), especially around the Irminger Basin (Elliot et al., 1998). Using the similarity of the ocean and ice records, Bond et al. (1993; 1995) proposed a common chronostratigraphic framework, in which the Heinrich events are linked to the largest coolings in the GRIP and GISP2 ice records, and in which the smaller IRD inputs are synchroneous of the other cold oscillations in the Dansgaard-Oeschger cycles. The climate system variability during glacial time and the relationships between the ice records and the oceanic records are summarized on Fig. 1. Together with the Greenland ice GRIP record, we presented data from several sediment cores from the North Atlantic Ocean for which high sedimentation rates made it possible to observe climatic oscillations with a centenial temporal resolution. The figure shows the results obtained in two different areas: the Faeroes-Scotland ridge documented by core ENAM93-21 located at 62°N and 3°W; and the Rockall plateau documented by core NA87-22 located at 54°N and 15°W (Fig. 2). The chronologies of the ice and ocean records are independant: The first one is based on layer counting in the first part of the ice core and an ice accumulation model when the ice layers are not visible (Dansgaard et al., 1993). GRIP and GISP2 age scales are different (Alley et aI., 1995), but, in this paper, we are only considering the morphology of the ice signal and we do not intend to make precise correlations between glacial and marine records. The sediment cores are AMS 14C dated, with error bars increasing from 50 years in surface samples to 1,000 years at about 40 kyr B.P. Between 40 and 60 kyr, the marine record is constrained by the age of the ash zone II, around 55 kyr (Ruddiman and Glover, 1972), and by the transition between isotopic stages 4 and 3 at 59 kyr (Martinson et aI., 1987). These two chronological pointers have an associated error bar of plus or minus 5kyr, and allowed us to constrain the lower part of the age-depth relationship. The final age scale then mixes a 14C age scale between 0 and 40 kyr, which is the period of interest, with two astronomical dates at the end of the record.
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Glacial and Interglacial Hydrological Changes in the North Atlantic Ocean
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The magnetic susceptibility record from core ENAM93-21 bears a noticeable similarity with the 818 0 of the ice at GRIP (Rasmussen et al., 1996). Each of the DansgaardOeschger events between 10 and 50 kyr is identified in the magnetic susceptibility record of this sediment core. The lowest magnetic susceptibility values correlate with cold temperatures of the air above Greenland and, in the opposite, the highest values are associated with warmer air temperatures. This record is characteristic of the whole Northern Atlantic Ocean deep sea sediments, from the Faeroes to the Southern Greenland Sea. This parameter appears to trace the variations in the size and concentration of magnetic minerals from the northern area and thus the activity of the bottom water circulation (Kissel et al., this issue). The Dansgaard-Oeschger oscillations are also correlated to variations in the 8180 of the planktic foraminifera, with temperature minima corresponding to a lowering of the surface water salinity of the North Atlantic and Southern Norwegian Sea. Based on the isotopic composition of the carbon of the benthic foraminifera Cibicides wuellerstorfi
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picked from the North Atlantic core NA87-22, Vidal et al. (1997), showed that reduction of deep water ventilation occured during the largest Heinrich events, and were therefore synchronous with sea surface drops in salinity and temperature, and reduced transport of magnetic particles by bottom current. These Ol3 C lowerings are recorded in other North Atlantic cores but we still do not have reliable data because of the paucity of C wuellerstorfi in sediments from this area. Understanding the role of the ocean regarding the climatic variability recorded in the Greenland ice cores has been one of the great challenges in the last few years for the paleoceanographic community, and is one of the major questions addressed by the IMAGES project (International Marine Global Change Study). These oscillations, or at least the largest ones during Heinrich events, are linked to iceberg discharges as recorded by the IRD in marine sediments. Several explanations have been given to understand how these iceberg discharges could occur (Denton et al., 1986; Paillard, 1995), amongst them internal instabilities of the Laurentide ice sheet (MacAyeal, 1993). Based on this binge and purge model from McAyeal, simple box model experiments indicate that the Laurentide ice sheet internal instabilities could lead to a slow-down of the thermohaline circulation and a subsequent cooling during Heinrich events, then followed by an intense and abrupt warming when the iceberg discharge spontaneously stops (Paillard and Labeyrie, 1994). The other cool phases of the smaller Dansgaard-Oeschger events seem to be also associated with iceberg discharges and surface water 0180 anomalies but with smaller amplitudes. At this stage, the l3C signal of marine records does not give any clear evidence of an equivalent reduction of the thermohaline circulation associated with the more frequent IRD events even if faunal benthic evidences seems to follow the overflow water reduction in the Norwegian Sea according to Rasmussen (1996). They could correspond to the rapid oscillations of smaller marine-based ice sheets around the northern Atlantic and the Norwegian sea, at the margins of the Fennoscandian, Iceland and Greenland ice sheets (Elliot et aI., 1998). Spontaneous oscillations of the thermohaline circulation are also a possible explanation for the Dansgaard-Oeschger cycles (Sakai and Peltier, 1997). But the precise chronology of environmental changes during DansgaardOeschger cycles as well as the exact relationship between Heinrich events and DansgaardOeschger events still remains unclear in many respects.
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3. THE INTERGLACIAL PERIOD 3.1. What Is Isotopic Substage 5.5? Before considering the climatic evolution of the last interglacial period, it is necessary to precise the duration and boundaries of the studied period. The duration of the last interglacial period is somewhat controversial and varies from IOkyr in the marine records (Martinson et al., 1987) to 24 kyr in glaciological records (see for example, Lorius (1985)). Benthic foraminifera 018 0 records as a proxy for global ice volume and sea level is the most reliable method to determine the isotopic stage 5.5 in marine sediments. In the age scale established by Martinson and others (1987), the isotopic substage 5.5 lasted roughly 10kyr, from 128 to 118kyr BP, from transition 6/5.5 to transition 5.5/5.4. However, as shown in Fig. 3, these transitions may sometimes be difficult to determine using the oxygen isotopic composition of the benthic foraminifera and can therefore be subject to various interpretations (Broecker, 1998). The only feature characteristic from
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one core to another is the plateau defined by the constant and minimum ice volume period in the isotopic composition of the benthic foraminifera. Then, to compare the climatic evolution of the isotopic substage 5.5, we decided to correlate, from one core to another, the start and the end of this plateau. This allowed us to define the "5.5 plateau" interval during which we will follow the hydrological changes of the surface waters.
3.2. Correlation and Age Scale The stratigraphy is defined using benthic 8180 records, and all the records are put on a common depth scale, using the core with the highest sedimentation rate, core JPC8 (around l8cmlkyr, (Oppo et al., 1997». When possible, the 5.5 plateau was defined using the first and last isotopically light events (respectively 5.53 and 5.51 events following (Martinson et al., 1987), Fig. 4). But in some cases, these events were not obvious and the trend present in the record did not allow us to identify them. We chose, in these cases, to use the average value of benthic 8180 of the 5.5 plateau plus or minus 0.2%0 to take into account the slight variations due to natural variability. Using this definition of the 5.5 plateau, sediment thickness of this interval ranges from 24cm (core MD95-2036) to 192cm (core JPC8), which represents an average sedimentation rate of 9cmlky to 73cmlky, using the ages defined by the Martinson et al. time scale (125.2 to l22.6ky for the 5.5 plateau, (Martinson et al., 1987». It is worth noticing that this time scale probably underestimates the duration of the 5.5 plateau. Indeed, according to this scale, the plateau would last only about 3ky or less, while the Holocene plateau is more than 9ky long and still counting (Broecker, 1998). Other timescales suggest that the duration of the 5.5 plateau could be more comparable to the Holocene, like the estimation made by Adkins and others (1997) using thorium measurements in the sediment. Another study by Winograd (1997) showed that the 5.5 plateau defined in the continental record of Devils Hole lasts around 10 kyr. We then decided to present our data versus the age scale defined by Adkins et al. (1997). It is worth noticing that, as long as we are comparing data versus the same reference, the age scale is not really important.
3.3. Methods and Results 3.3.1. Methods. In each core (Table 1), we used measurements of the 8 180 of benthic foraminifera: Cibicides wuellerstorfi, Oridorsalis tener, and Uvigerina peregrina, corrected
Table 1. Location of the cores discussed in this study Core
Latitude
Longitude
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time period
Reference
NA87-22 ENAM93-21 SU90-08 V27-60 JPC8 NA87-25 CH69-K9 MD95-2036 JPC37
55°29'N 62°N 400 N 72°11 'N 61°N 55°II'N 41°N 33°41'N 31 0 41'N
J4041'W 3°W 300 W 8°35'E 25°W 14°57'W 47°N 57°34'W 75°25'W
2,161 1,000 3,080 2,525 1,917 2,320 4,100 4,461 2,972
glacial glacial glacial interglacial interglacial interglacial interglacial interglacial interglacial
I 2 I 3 4 3 5 5 5
References given on the table: 1, (Vidal et al., 1997); 2, (Rasmussen et al., 1996); 3, (Cortijo et al., 1994); 4, (Oppo et aI., 1997); 5, (Cortijo et al., 1999).
Glacial and Interglacial Hydrological Changes in the North Atlantic Ocean
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to take into account the isotopic fractionation and the 8 180 of planktonic species: Neogloquadrina pachyderma left coiling in core V27-60, N pachyderma dextral coiling in core JPC8 (Oppo et al., 1997), Globigerina bulloides in core NA87-25 and CH69-K9 and Globigerinoides ruber white variety in cores MD95-2036 and JPC37. A complete description of these records is available in (Cortijo et al., 1999). Data are reported versus PDB (Pee Dee Belemnite standard) after calibration with NBSl9 (Coplen, 1988). Sea surface temperatures (SST) are based on planktonic species counts of at least 300 individuals using the modern analog technique. Paleotemperatures are estimated by identifying the five most similar core top samples in the North Atlantic data base (615 core top samples between 0 and 80 0N, modified from (Ptlaumann et al., 1996)). Summer and winter SSTs are then estimated by averaging the summer and winter SSTs associated with these most similar core tops (Prell, 1985). Dissimilarity between sample and coretop assemblages, using 32 planktonic taxa, is calculated using the squared chord distance. Uncertainties in the SST reconstructions correspond to the root mean square error of the top 5 analog temperatures. In each case, the 5 most similar core-top samples are accepted as valid modern analogs for the studied fossil sample, with dissimilarity coefficient lower than 0.2. The error bars on temperature reconstructions are between 0.5 and 2°C. However, in core MD95-2036 and JPC37 the standard deviation can exceed 2.5°C and reach 3.5 °C because of the low number of available reference core tops in the data base in the western tropical Atlantic. Sea water 8 180 was estimated following the method described in Duplessy et al. (1991). The overall glacial-interglacial amplitude between isotopic stages 6 and 5.5 in the benthic 8 180 records corresponds to the value of 1.2%0 found in core V27-60 (nON), which can be attributed solely to global ice volume variations (Fairbanks, 1989; Labeyrie et al., 1987). For the same transition, North Atlantic cores (JPC8, NA87-25, CH69-K9, MD95-2036, JPC37) record a benthic 8 180 shift of 1.7 to 2%0 (Fig. 4). The 0.5 to 0.8%0 excess may be attributed to a warming of the deep waters of approximately 2°C to 3.5 0C. The glacial-interglacial amplitude of the 8 180 planktic records ranges from 2%0 to 3%0, from which 0.8%0 to 1.8%0 must be attributed to temperature and salinity changes. Considering that this whole change is due to temperature, the glacial-interglacial temperature shift recorded by planktic 0180 will be around 30 to 8°C, which in good agreement with the amplitude of the estimated summer SST changes, from around 4°C in cores JPC37 and V27-60 to around 9°C in cores CH69-K9 and JPC8. 3.3.2. Results.
3.3.3. The 5.5 Plateau.
This interval is defined by minimal variations of the benthic 8 180 records and, in most cases, this signal does not show any significant variations, except for the core CH69-K9 where there is a slight increasing trend inside the plateau (Fig. 4). This trend could be caused by a progressive change in deep water masses at this location. The planktic 8180 records in the different cores show a variability of plus or minus 0.5%0 probably linked to changes in surface temperature and salinity values (Fig. 4). However, the G bulloides 8180 record in core CH69-K9 shows a peculiar positive event, with 8180 values increasing by I%0, in the middle of isotopic substage 5.5. Within isotopic substage 5.5, the summer SSTs follow two different profiles (Fig. 5). Northern cores V27-60, JPC8 and NA87-25 show a decreasing trend during the second half of the 5.5 with summer SST values decreasing by 20 to 4°C. The southern cores CH69-K9, MD95-2036 and JPC37 show an increasing trend during the same time period, with summer SST values increasing by 10 to 4°c. Other published records show
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