Introduction

The late glacial, the transition from the last glacial maximum (LGM) to the Holocene, is a key period for understanding the mechanisms of abrupt climate change. The late glacial is characterized by three short-term oscillations, called the Oldest Dryas, Older Dryas, and Younger Dryas, respectively. They were defined at the beginning of the century in Denmark at a time when no radiometric dating was available. The distinc tion between the Older Dryas and the Younger Dryas interval was first established at the Allerad site in Denmark (Hartz and Milthers, 1901) based on sediment stratigraphy. Both phases, separated by a layer of gytt-ja that corresponds to the Allerad interstadial, are characterized by the presence of macrofossils of subarctic / alpine flora, including Dryas octopetala, and the absence of significant amounts of tree birch (Betula sp.). The B0lling interstadial in turn was identified at B0lling S0, Denmark, from a layer of sand and silty gyttja containing a high percentage of tree birch pollen (Iversen, 1942, 1954). The underlying inorganic layer was named the Oldest Dryas interval (in Sanchez-Goni, 1995). In subsequent years, these intervals were dated in many sites in Europe, yielding the following chronology: (1) Oldest Dryas, 15,500-14,500 cal. B.P.; (2) B0lling/Aller0d, 14,500-12,500 cal. B.P., interrupted by three-decade- to century-long events of the Inter-B0lling Cold Period (IBCP), the Older Dryas, and the Inter-Allerod Cold Period (IACP); and (3) the Younger Dryas, 12,500-11,000 cal. B.P. (Fig. 1).

Recent studies on marine cores from the North Atlantic region (Bard et al., 1987, 1990, 1997; Bond et al. 1993; Bond and Lotti, 1995) and on ice cores from Greenland (Jouzel et al., 1987; Johnsen et al., 1992; Dansgaard et al., 1993) have highlighted drastic climat

Dansgaard 1993

FIGURE 1 Major S18O variations and European pollen zone boundaries during late glacial times. (From Stuiver et al., 1995. With permission.)

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FIGURE 1 Major S18O variations and European pollen zone boundaries during late glacial times. (From Stuiver et al., 1995. With permission.)

ic changes of very short duration during the late glacial (Fig. 1). It appears that the last deglaciation was a two-step process with large and rapid fluctuations between warmer and colder conditions. Mounting evidence from marine records elsewhere has documented that these abrupt changes were not restricted to the North Atlantic region. In particular, studies from the Cariaco basin (Venezuela), close to the equator, have revealed the same sequence of short-term warmer and cooler periods identified in terrestrial records (Hughen et al., 1996). For the Arabian Sea, Schulz et al. (1998) and Overpeck et al. (1996) showed that following the last glaciation, the monsoon pattern strengthened in a series of steps coinciding with major climate shifts in the polar regions of the Northern Hemisphere. The similarities of events in records from the Cariaco basin, the Arabian Sea, and the high-latitude North Atlantic suggest a common forcing mechanism. For many years, it has been demonstrated that insolation, especially precession, was a dominant factor regulating tropical climate changes during the Pleistocene (Street and Grove, 1979; Prell and van Campo, 1986; Prell and Kutzbach, 1987; Partridge et al., 1997). However, the abrupt climatic changes evidenced during the last deglaciation cannot be accounted for by orbital forcing. Instead, these changes have been linked to rapid reorganizations of the North Atlantic thermohaline circulation (Broecker et al., 1985, 1988; Lehman and Keigwin, 1992; Broecker 1994; Zaucker et al., 1994).

Although in Antarctica the glacial-interglacial warming also occurred in two steps, interrupted by a slight cooling period, the Antarctic Cold Reversal (Jouzel et al., 1995; Sowers and Bender, 1995), it is out of phase with the sequence of climate shifts in the North Atlantic (Broecker, 1998; Blunier et al., 1998).

There is considerable uncertainty in terrestrial records about the exact timing, rate, and magnitude of these changes. The response of the land-atmosphere system to these climatic oscillations is apparently not regionally uniform. Therefore, the global extent of the changes remains in debate.

When the sequences of late glacial fluctuations in pa-lynological continental records along the Pole-Equator-Pole: Americas (PEP 1) transect are compared, the following questions arise:

1. What is the timing of the onset of the late glacial in different types of environments? Given differences in basal ages of the records, this question is often problematic. Records located in the tropical lowlands, for example, start right after the LGM; in high-latitude, high-altitude, and midlatitude records, sedimentation in lakes or bogs generally started only after the glaciers had retreated locally at ca. 15,000 cal. B.P.

2. How are late glacial changes expressed in different vegetation types?

3. Is the sequence of cold (stadial) and warm (interstadial) fluctuations during the time interval between 17,000 and 11,000 cal. B.P. consistent and synchronous over large distances?

In the following, we will discuss three time intervals corresponding to the European late glacial stratigraphy; these intervals are the Oldest Dryas, 15,50014,500 cal. B.P.; the B0lling-Aller0d, 14,500-12,700 cal. B.P.; and the Younger Dryas, 12,700-11,000 cal. B.P.

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