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The timing and course of the last deglaciation is generally considered to be an essential component for understanding the dynamics of large ice sheets (Lindstrom and MacAyeal, 1993) and their effects on Earth’s isostasy (Lambeck, 1993; Peltier, 1994). Moreover, the disappearance of glacial ice sheets was responsible for dramatic changes in the freshwater fluxes to the oceans, which disturbed the general thermohaline circulation and, hence, global climate (e.g., Stocker and Wright, 1991). Coral reefs are excellent sea level indicators, and their accurate dating by mass spectrometry is of prime importance for determination of the timing of deglaciation events and thus for understanding of the mechanisms driving glacial–interglacial cycles. Furthermore, scleractinian coral colonies can monitor sea-surface temperatures (SSTs) and record past SSTs.

Sea Level Changes as a Global Climate Indicator

Only three deglaciation curves based on coral reef records have been accurately dated for times reaching the Pleistocene/Holocene boundary: in Barbados (Fairbanks, 1989; Bard et al., 1990a, 1990b) between 19.00 and 8.00 ka, in New Guinea between 13.00 and 6.00 ka (Chappell and Polach, 1991; Edwards et al., 1993), and in Tahiti between 13.75 ka and 2380 14C y before present (BP) (Bard et al., 1996) (Fig. F1). So far, the Barbados curve is the only one to encompass the whole deglaciation because it is based on offshore drilling. However, this site, like New Guinea, is located in an active subduction zone where tectonic movements can be large and discontinuous, so that the apparent sea level records may be biased by variations in the rates of tectonic uplift. Hence, there is a clear need to study past sea level changes in tectonically stable regions or in areas where the vertical movements are slow and/or regular.

The Barbados record suggested that the last deglaciation was characterized by two brief periods of accelerated melting superimposed on a smooth and continuous rise of sea level with no reversals (Fig. F1). These so-called meltwater pulse (MWP)-1A and MWP-1B events (~13.80 and 11.30 ka, respectively) are thought to correspond to massive inputs of continental ice (i.e., ~50–40 mm/y, roughly equivalent to annual discharge rates of 16,000 km3 for MWP-1A). MWP-1A corresponds to a short and intense cooling period between 14.10 and 13.90 ka in the Greenland records (Johnsen et al., 1992; Grootes et al., 1993) and therefore postdates the initiation of the Bölling-Alleröd warm period at ~14.90–14.70 ka (Broecker, 1992). The sea level jump evidenced in New Guinea at 11.00 ka (Edwards et al., 1993) is delayed by a few centuries when compared to that observed at Barbados. These two meltwater pulses are thought to have induced reef-drowning events (Blanchon and Shaw, 1995). Two “give-up” reef levels have been reported at 90–100 and 55–65 m water depth on the Mayotte foreslopes (Comoro Islands) and have been related to the Bölling and the post-Younger Dryas meltwater pulses (Dullo et al., 1998); similar features are recorded in the southern Great Barrier Reef (GBR) (Troedson and Davies, 2001) and in the Caribbean (MacIntyre et al., 1991; Grammer and Ginsburg, 1992). A third Acropora reef-drowning event at ~7.60 ka has been assumed by Blanchon and Shaw (1995).

However, there are still some doubts concerning the general pattern of sea level rise during the last deglaciation events, including the amplitude of the maximum lowstand during the Last Glacial Maximum (LGM) and the occurrence of increased glacial meltwater with resultant accelerated sea level rise (Broecker, 1990). Furthermore, sawtooth sea level fluctuations between 19.00 and 15.28 ka (Locker et al., 1996) and a sea level fall coeval with climatic changes at ~11 ka are still controversial topics.

Worldwide sea level compilations indicate that local sea level histories varied considerably around the world in relation to both postglacial redistribution of water masses and a combination of local processes (Lambeck, 1993; Peltier, 1994), although significant deviations between model predictions and field data have been noted in several regions (Camoin et al., 1997). The post–last glacial sea level changes at sites located far away from glaciated regions (“far field”) provide basic information regarding the melting history of continental ice sheets and the rheological structure of the Earth. The effect of hydroisostasy depends on the size of the islands: for very small islands, the addition of meltwater produces a small differential response between the island and the seafloor, whereas the meltwater load produces significant differential vertical movement between larger islands or continental margins and the seafloor (Lambeck, 1993). There is therefore a need to establish the validity of such effects at two ideal sites located at a considerable distance from the major former ice sheets: (1) on a small oceanic island and (2) on a continental margin. In both cases, it is essential for the sites chosen that the tectonic signal is small or regular within the short time period proposed for investigation so that rigorous tests of proposed northern and southern hemisphere deglaciation curves from Barbados and New Guinea can be made. Two such places are proposed: Tahiti and the GBR. This mission-specific platform (MSP) expedition will conduct investigations at the Tahiti sites only.

Climatic and Oceanographic Changes during Last Deglaciation Events

During latest Pleistocene and early Holocene, climatic variability was primarily related to the effects of seasonality and solar radiation. The results of the Long-Range Investigation, Mapping, and Prediction (CLIMAP) program suggested that LGM tropical SSTs were similar to modern ones. However, this interpretation is not consistent with snowline reconstructions and paleobotanic data (Rind and Peteet, 1985; Anderson and Web, 1994).

The available Sr/Ca and U/Ca data from coral reef areas report SSTs 5°C colder than today during the LGM and 2°C lower at ~10–9 ka at Barbados (Guilderson et al., 1994), whereas studies in the west Pacific indicate that the full amplitude of the glacial Holocene temperature change may have ranged between 3° and 6°C (McCulloch et al., 1996; Beck et al., 1997; Gagan et al., 1998) (Fig. F1). Troedson and Davies (2001) define SSTs immediately south of the GBR some 4.5°C colder during the LGM and 1°C colder at 10 ka. This casts doubt upon the phase shift of 3000 y for climate changes between the two hemispheres that was assumed by Beck et al. (1997), in clear distinction to the apparent synchronism of the last deglaciation, inferred from various sources (i.e., coral records, ice cores, snowline reconstructions, vegetation records, and alkenone palaeothermometry) (Bard et al., 1997).

Recent studies have documented Holocene climatic variations. SSTs warmer by 1°C, monsoonal rainfall, and possibly weaker El Niño–Southern Oscillation (ENSO) at ~5.80 ka in Eastern Australia have been deduced from isotopic and Sr/Ca high-resolution measurements on corals from the central GBR (Gagan et al., 1998). An ENSO-like cyclic climatic variation with a return period of 3–5 y has been evidenced in a 4150 y old coral from the Seychelles, although the intensity of the annual decrease in SST caused by monsoonal cooling was lower than that of today (Zinke et al., 2005).

Additional information is required for better knowledge of climatic conditions in tropical regions during the last deglaciation. In these areas, the most debated points are twofold: (1) the quantification of SSTs and the identification of related climatic variations during the last deglaciation events and (2) the timing of the relevant postglacial warming in the two hemispheres.

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