DEGLACIAL WARMING OF THE NORTH ATLANTIC Over the last million years, the Earth's climate experienced major changes, with alternation of glacial and interglacial periods. These climatic variations are recorded in the variations of the fossil planktonic foraminiferal shells, which accumulate on the sea floor : On the one hand, as the composition of the fauna in ocean surface water depends mainly on the temperature, the composition of the fossil fauna in a deep sea core may be used to estimate the past sea surface temperature (2). On the other hand, continental ice-volume changes may be reconstructed by oxygen isotopic analysis of the same foraminiferal shells (3). The timing of the deglaciation has important consequences relating to mechanisms of climate change. However, details of the climatic evolution of the ocean and of the continental ice-volume during the last glacial to interglacial transition are not recorded in most sediment cores because of the short duration of this event. Althrough the sediment material is deposited at a rate on the order of a few centimeters per thousand years, a major problem encountered when attempting to record such a rapid climatic change in deep sea records is created by the natural process of bioturbation: The activity of benthic organisms disturbs the original stratification of the deposited shells and mixes the upper few centimeters of sediment. As a result, the variations of the paleontological and isotopic compositions in the sedimentological record are both smoothed and disturbed. Radiocarbon ages can be measured by A. M. S. on monospecific samples of 1,000-2,000 foraminifera picked by hand. This technique offers several advantages: First, the foraminiferal samples can be much more pure than those used for the classical method of 8-counting, which required samples so large that usually the carbonate fraction larger than 60 µ was analyzed. The presence within the North Atlantic sediment of ice-rafted carbonate, or of wind-, river- and current-transported carbonates (derived from old continental rocks), biased classical 14C ages towards older values. Second, as stable isotopes are also measured on monospecific samples, this enables an age to be assigned to the same sample which is to be used in the a180 measurement (4). A simple deconvolution model can then be used to date precisely the internal stages of the deglaciation and to measure the volocity of the climatic changes in the North Atlantic Ocean (5). Fig. 1 displays the oxygen isotope record measured in the foraminifer Globigerina bulloides, the 14 C ages measured in the same species and the sea surface temperature (S.S.T.) estimates in core SU 81-18 (37°46' N, 10°11' W). The sedimentation rate during the deglaciation is so high in this core (35 cm/kyr), that the bioturbation smoothing is negligible. The deglaciation (defined on the isotopic record) began about 14,500 years ago. The melting of continental ice first resulted in aS.S.T. drop of 6°C for both summer and winter S.S.T. The end of the first phase of the deglaciation, after 12,500 ± 150 B.P., was marked by a sharp temperature rise with a mean rate of roughly 4°C per century. This warm phase was followed by a dramatic cooling, well-known as Younger Dryas cold event (from 11,000 ± 170 B.P. to 10,400 ± 130 B.P.). The readvance of the cold surface water corresponds to a temperature drop of 0,5 to 1°C per century. The final S.S.T. warming, which led to modern conditions, was less abrupt than the first one with a mean rate of temperature rise close to 1°C per century until 9,360 ± 130 B.P.. Fig 2A displays the oxygen isotope records of both G. bulloides and N. pachyderma (left coiling) in core CH 73-139 C (54°38' N, 16°21' W). The interpretation of the isotopic date is more difficult than in core SU 81-18, because the bioturbational disturbance is not negligible. For example, the D180 values of G.bulloides exhibit

DOI link for DEGLACIAL WARMING OF THE NORTH ATLANTIC Over the last million years, the Earth's climate experienced major changes, with alternation of glacial and interglacial periods. These climatic variations are recorded in the variations of the fossil planktonic foraminiferal shells, which accumulate on the sea floor : On the one hand, as the composition of the fauna in ocean surface water depends mainly on the temperature, the composition of the fossil fauna in a deep sea core may be used to estimate the past sea surface temperature (2). On the other hand, continental ice-volume changes may be reconstructed by oxygen isotopic analysis of the same foraminiferal shells (3). The timing of the deglaciation has important consequences relating to mechanisms of climate change. However, details of the climatic evolution of the ocean and of the continental ice-volume during the last glacial to interglacial transition are not recorded in most sediment cores because of the short duration of this event. Althrough the sediment material is deposited at a rate on the order of a few centimeters per thousand years, a major problem encountered when attempting to record such a rapid climatic change in deep sea records is created by the natural process of bioturbation: The activity of benthic organisms disturbs the original stratification of the deposited shells and mixes the upper few centimeters of sediment. As a result, the variations of the paleontological and isotopic compositions in the sedimentological record are both smoothed and disturbed. Radiocarbon ages can be measured by A. M. S. on monospecific samples of 1,000-2,000 foraminifera picked by hand. This technique offers several advantages: First, the foraminiferal samples can be much more pure than those used for the classical method of 8-counting, which required samples so large that usually the carbonate fraction larger than 60 µ was analyzed. The presence within the North Atlantic sediment of ice-rafted carbonate, or of wind-, river- and current-transported carbonates (derived from old continental rocks), biased classical 14C ages towards older values. Second, as stable isotopes are also measured on monospecific samples, this enables an age to be assigned to the same sample which is to be used in the a180 measurement (4). A simple deconvolution model can then be used to date precisely the internal stages of the deglaciation and to measure the volocity of the climatic changes in the North Atlantic Ocean (5). Fig. 1 displays the oxygen isotope record measured in the foraminifer Globigerina bulloides, the 14 C ages measured in the same species and the sea surface temperature (S.S.T.) estimates in core SU 81-18 (37°46' N, 10°11' W). The sedimentation rate during the deglaciation is so high in this core (35 cm/kyr), that the bioturbation smoothing is negligible. The deglaciation (defined on the isotopic record) began about 14,500 years ago. The melting of continental ice first resulted in aS.S.T. drop of 6°C for both summer and winter S.S.T. The end of the first phase of the deglaciation, after 12,500 ± 150 B.P., was marked by a sharp temperature rise with a mean rate of roughly 4°C per century. This warm phase was followed by a dramatic cooling, well-known as Younger Dryas cold event (from 11,000 ± 170 B.P. to 10,400 ± 130 B.P.). The readvance of the cold surface water corresponds to a temperature drop of 0,5 to 1°C per century. The final S.S.T. warming, which led to modern conditions, was less abrupt than the first one with a mean rate of temperature rise close to 1°C per century until 9,360 ± 130 B.P.. Fig 2A displays the oxygen isotope records of both G. bulloides and N. pachyderma (left coiling) in core CH 73-139 C (54°38' N, 16°21' W). The interpretation of the isotopic date is more difficult than in core SU 81-18, because the bioturbational disturbance is not negligible. For example, the D180 values of G.bulloides exhibit

DEGLACIAL WARMING OF THE NORTH ATLANTIC Over the last million years, the Earth's climate experienced major changes, with alternation of glacial and interglacial periods. These climatic variations are recorded in the variations of the fossil planktonic foraminiferal shells, which accumulate on the sea floor : On the one hand, as the composition of the fauna in ocean surface water depends mainly on the temperature, the composition of the fossil fauna in a deep sea core may be used to estimate the past sea surface temperature (2). On the other hand, continental ice-volume changes may be reconstructed by oxygen isotopic analysis of the same foraminiferal shells (3). The timing of the deglaciation has important consequences relating to mechanisms of climate change. However, details of the climatic evolution of the ocean and of the continental ice-volume during the last glacial to interglacial transition are not recorded in most sediment cores because of the short duration of this event. Althrough the sediment material is deposited at a rate on the order of a few centimeters per thousand years, a major problem encountered when attempting to record such a rapid climatic change in deep sea records is created by the natural process of bioturbation: The activity of benthic organisms disturbs the original stratification of the deposited shells and mixes the upper few centimeters of sediment. As a result, the variations of the paleontological and isotopic compositions in the sedimentological record are both smoothed and disturbed. Radiocarbon ages can be measured by A. M. S. on monospecific samples of 1,000-2,000 foraminifera picked by hand. This technique offers several advantages: First, the foraminiferal samples can be much more pure than those used for the classical method of 8-counting, which required samples so large that usually the carbonate fraction larger than 60 µ was analyzed. The presence within the North Atlantic sediment of ice-rafted carbonate, or of wind-, river- and current-transported carbonates (derived from old continental rocks), biased classical 14C ages towards older values. Second, as stable isotopes are also measured on monospecific samples, this enables an age to be assigned to the same sample which is to be used in the a180 measurement (4). A simple deconvolution model can then be used to date precisely the internal stages of the deglaciation and to measure the volocity of the climatic changes in the North Atlantic Ocean (5). Fig. 1 displays the oxygen isotope record measured in the foraminifer Globigerina bulloides, the 14 C ages measured in the same species and the sea surface temperature (S.S.T.) estimates in core SU 81-18 (37°46' N, 10°11' W). The sedimentation rate during the deglaciation is so high in this core (35 cm/kyr), that the bioturbation smoothing is negligible. The deglaciation (defined on the isotopic record) began about 14,500 years ago. The melting of continental ice first resulted in aS.S.T. drop of 6°C for both summer and winter S.S.T. The end of the first phase of the deglaciation, after 12,500 ± 150 B.P., was marked by a sharp temperature rise with a mean rate of roughly 4°C per century. This warm phase was followed by a dramatic cooling, well-known as Younger Dryas cold event (from 11,000 ± 170 B.P. to 10,400 ± 130 B.P.). The readvance of the cold surface water corresponds to a temperature drop of 0,5 to 1°C per century. The final S.S.T. warming, which led to modern conditions, was less abrupt than the first one with a mean rate of temperature rise close to 1°C per century until 9,360 ± 130 B.P.. Fig 2A displays the oxygen isotope records of both G. bulloides and N. pachyderma (left coiling) in core CH 73-139 C (54°38' N, 16°21' W). The interpretation of the isotopic date is more difficult than in core SU 81-18, because the bioturbational disturbance is not negligible. For example, the D180 values of G.bulloides exhibit