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Millennial-scale climate variability and vegetation changes during the Last Glacial: Concepts and terminology
Quaternary Science Reviews 29 (2010) 2823–2827
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Introduction
Millennial-scale climate variability and vegetation changes during the Last Glacial: Concepts and terminology
The Late Quaternary, and more particularly the Last Glacial (Marine Isotope Stages 4, 3 and 2, defined here as 73,500–14,700 calendar years, 73.5–14.7 ka; ka is used solely for calendar years and kyr BP refers to 14C ages), was characterised by millennial-scale climate oscillations of irregular periodicity. The onset of these oscillations was abrupt, with most of the change in climate accomplished within 10–200 years (Steffensen et al., 2008) and the magnitude of the change was large (of the order of 8–15 C in Greenland) (Huber et al., 2006). Two types of rapid climate changes have been described: Dansgaard–Oeschger (D-O) cycles (Dansgaard et al.,1984), associated with abrupt warming and subsequent cooling in Greenland, and cold phases associated with the formation of ice-rafted debris (IRD) deposits (Heinrich layers) (Heinrich, 1988) in the North Atlantic. Dansgaard–Oeschger cycles are clearly registered in the Greenland ice-core record (e.g. Johnsen et al., 1992; North GRIP Members, 2004) and the traces of both of these climate oscillations are recorded in a variety of marine and terrestrial records worldwide (e.g. Bond et al., 1993; Allen et al., 1999; Wang et al., 2001; Gonzalez et al., 2008). The geographic pattern of registration of individual oscillations ˜ i et al., 2008), and the magnitude, nature and length of (Sanchez Gon the component phases of each recorded oscillation, appear to vary (Johnsen et al., 1992). However, documentation of regional changes has been hampered by problems of the synchronisation of individual chronologies, and our understanding of the mechanisms underlying these climate oscillations is still far from complete. Analysis of the mechanisms and impacts of large and rapid climate changes in the past is given additional impetus by the possibility that such events might occur in the future. Although the precise causes are different, investigation of the impact of the warming events at the beginning of D-O cycles or of iceberg melting during Heinrich intervals on the Meridional Overturning Circulation (MOC) in the North Atlantic speaks directly to the impact of future changes in the MOC on regional climate. Coupled ocean– atmosphere model simulations show a reduction of the MOC during the 21st century, in some cases by up to 50%, as a consequence of greenhouse-gas-induced polar warming (Gregory et al., 2005). Simulations with simpler climate models have shown that complete shutdown of the MOC can occur if the slowdown reaches a crucial threshold (which differs in different models) (Stocker and Schmittner, 1997; Stouffer and Manabe, 2003; Stouffer et al., 2006). Clearly, the coupled models do not reach the critical threshold as a result of the gradual change in greenhouse-gas forcing during the 21st century but might do so if the additional forcing due to even partial melting of the Greenland ice sheet were taken into account. 0277-3791/$ – see front matter Ó 2009 Elsevier Ltd. All rights reserved. doi:10.1016/j.quascirev.2009.11.014
There have been several recent attempts to synthesise millennial-scale climate change during the glacial (e.g. Broecker and Hemming, 2001; Alley et al., 2002; van Andel, 2002; Voelker, 2002), but most focus on marine and ice-core records. In this issue, we focus on documenting regional changes in vegetation indicated by pollen records from both marine and terrestrial cores (Fletcher et al., 2010; Takahara et al., 2010; Jime´nez-Moreno et al., 2010; Heßler et al., 2010). There are multiple reasons why this is timely. First, there has been a very rapid increase during the last decade in the number of pollen records with high temporal resolution. However, there is no global compilation of the pollen data, nor a synthesis of vegetation reconstructions based on these data. Second, the development of a new and coherent ice-core reference chronology (GICC05: Svensson et al., 2006, 2008; see also Wolff et al., 2010) makes it possible to achieve a better synchronisation between documented changes in Greenland and the pollen records. Third, a regional synthesis of charcoal records from North America (Marlon et al., 2009) shows that fire regimes respond to abrupt climate changes during the last deglaciation – but little is known about the response of fire globally to millennial-scale climate variability and associated vegetation changes during earlier intervals (Daniau et al., 2010). Fourth, investigation of the impact of changes in vegetation cover, including wetland extent, and in fire regimes is important for understanding the rapid and extremely large (up to ca 200 ppb) changes in methane during D-O cycles (Blunier and Brook, 2001; Flu¨ckiger et al., 2004) and the potential climate feedback. Finally, modelling groups have recently begun to explore the impact of, e.g., changes in freshwater forcing under glacial conditions on regional climates (e.g. Crucifix et al., 2001; Ganopolski and Rahmstorf, 2001; Claussen et al., 2003; Knutti et al., 2004; Flu¨ckiger et al., 2006, 2008; see also Kageyama et al., 2010) but more detailed documentation of observed changes is required for the evaluation of these experiments. A multiplicity of terms is used in discussing rapid climate changes and millennial-scale climate variability during the glacial, and this has led to some confusion particularly in relating records from different regions. Alley et al. (2002), in a definition adopted by the IPCC (Meehl et al., 2007), define abrupt climate change as one that takes place more rapidly than the underlying forcing, pointing out that this kind of behaviour can only occur when the climate system crosses a critical threshold defining the limit between two different climate states. This definition provides a theoretical basis for understanding abrupt climate changes (e.g. Kageyama et al., 2010) but the rapidity of the climate change is
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a function of the temporal scale characteristic of the specific forcing involved and the definition may be difficult to apply in the case of geological records where the nature of the forcing is a priori unknown. For practical reasons, many authors therefore identify abrupt climate changes in geological records in terms of some combination of magnitude of the change and the rapidity with which it is accomplished (e.g. McManus et al., 1999; Martrat et al., 2004). Here, we continue to use the term abrupt climate change in the sense of Alley et al. (2002), but refer to changes in climate that take place in <200 years and, in magnitude, exceed the decadal variability typical of the interval in which they occur as a rapid climate change. Dansgaard–Oeschger (D-O) cycles are an example of a rapid climate change, characterised in Greenland by a marked warming followed by a cooling. In most cases, the cooling occurs in three phases: initially relatively slow, then more precipitous, and a final slow phase. Here we refer to the initial rapid warming as the D-O event (Fig. 1). This event and the first slow phase of the subsequent cooling is referred to as the Greenland Interstadial (GI: Svensson et al., 2006; note this was originally abbreviated to GIS, see North GRIP Members, 2004); the final phase of cooling is referred to as the Greenland Stadial (GS: Svensson et al., 2006); the transition between the GI and GS is placed during the rapid phase of cooling (Fig. 1). Generally, the warming takes place in ca 60 years but the cooling can take more than 2000 years (e.g. D-O 12). However, there are D-O cycles in which initial warming persists for ca 100 years and the cooling occurs almost as quickly as the initial warming (e.g. D-O 6). Where there is a marked phase of rapid cooling, the limit between the GI and GS is relatively easy to define, but there are many D-O cycles where the location of transition between GI and GS is more subjective and can only be defined in terms of intervals when temperatures are equivalent to the general ‘‘glacial’’ baseline level (Wolff et al., 2010). The term ‘‘stadial’’ corresponds to cold intervals marked not by global but by local ice readvances (e.g. Lowe and Walker, 1984).
Fig. 1. Structure of exemplar D-O cycles, showing the rapid warming, and the intervals corresponding to the Greenland Interstadial (GI) and the Greenland Stadial (GS).
However, the term ‘‘interstadial’’ was originally defined in terms of vegetation development in Europe, and specifically to denote periods of warming that were either too short or of insufficient magnitude to lead to the development of deciduous forest in northern Europe (Fletcher et al., 2010). The term ‘‘interstadial’’ and, by extension ‘‘stadial’’ have subsequently been applied in other regions to indicate climate conditions associated with different types of vegetation development (e.g. Tzedakis et al., 2002). This climate-related terminology continues to be used in the palynological literature, and in cases where the pollen records are not precisely dated this can lead to confusion with the chronostratigraphic terms GI and GS. Here, and we would suggest in future discussions of pollen records, we confine the use of ‘‘interstadial’’ and ‘‘stadial’’ to those temporal intervals defined by the GI and GS – whatever the nature of the accompanying vegetation changes or indeed whether these chronostratigraphic intervals are accompanied by changes in vegetation or not. This is particularly important for documentation of millennial-scale variability in the southern hemisphere (Heßler et al., 2010). As has been clearly demonstrated, Antarctic temperatures cool during warm intervals in Greenland and warm during cooling phases in Greenland. The warm phase in the Antarctic leads D-O warming events by ca 1–2.5 ka. The warm phases were originally referred to as Antarctic warmings (Blunier et al., 1998) but the term Antarctic Isotopic Maxima (AIM: EPICA community members, 2006; Wolff et al., 2010) is preferable to avoid any possible implication that these are synchronous or necessarily mechanistically linked with the GI. The cold intervals associated with Heinrich layers in the North Atlantic (Heinrich, 1988) are a second example of rapid climate change. Again, there is some confusion in the literature between terms that are related to the observations and those that describe the climate changes associated with these observations. Thus, the term Heinrich Event (HE) was originally defined as corresponding to the period of time synchronous with the deposition of the Heinrich layer (e.g. Bond and Lotti, 1995), but subsequently applied to cover the whole of the associated cold interval (e.g. Bard et al., 2000). In some regions of the North Atlantic, the cold interval is indeed coincident with continuous deposition of IRD. However, further south, the deposition of IRD occupies a short interval of time and occurs at the end of the cold phase (e.g. off the SW Iberian Peninsula: Naughton et al., 2009) or there are low levels of IRD throughout but with a pronounced maximum of deposition towards the end of the cold phase (e.g. off the NW Iberian Peninsula, offshore from the SE USA: Naughton et al., 2009). Here (Fig. 2), we use the term HE to refer to the interval of the formation of the Heinrich layer in a given region and the cold interval associated with an HE as the Heinrich Stadial (HS: Barker et al., 2009). This usage is consistent with the terminology employed by, e.g., Elliot et al. (2002) and Hemming (2004), who referred to the whole of the cold period as a Heinrich Event because of their focus on the northern North Atlantic where the deposition of IRD occurs throughout the cold phase (Fig. 2). We use the term HS as a chronostratigraphic term, defined as the maximum length of the interval of IRD deposition based on a composite of marine records from the North Atlantic (Elliot et al., 1998, 2001, 2002; Hemming, 2004; Table 1). In some regions, e.g. the tropics (Heßler et al., 2010), the HS is registered as an increase in plant-available moisture rather than a cooling. Indeed, in some regions of the Southern Hemisphere, the chronostratigraphic interval corresponding to HS is not marked by strong vegetation changes (Kershaw personal communication). The deposition of Heinrich layers is associated with certain of the D-O cycles (e.g. HS 5 with the cold phase of D-O 13, i.e. GS 12/13) (Table 1). In most cases, however, the HS (as defined by
Introduction / Quaternary Science Reviews 29 (2010) 2823–2827
Fig. 2. Structure of an exemplar Heinrich Stadial, showing the relationship between sea-surface temperature and deposition of Heinrich layers, and hence the coincidence or otherwise between Heinrich Events (HE) and Heinrich Stadials (HS), at a typical northern and a typical southern North Atlantic site.
cold sea-surface temperatures in the North Atlantic) is shorter than the cold interval in Greenland (as defined by the isotopic record). For example, the estimated length of HS2 is 26.5–24.3 (Elliot et al., 2002) but GS2/3 lasts from 27.48 to 23.34. In the case of HS4, however, the two cold phases appear to be of approximately the same length. Not all D-O cycles are associated with HE. The rapid climate changes associated with D-O cycles or HS are superimposed on the longer-term climate trends within the glacial defined by isotope records from marine cores. These Marine Isotope Stages (MIS) correspond to intervals of more (MIS 4, MIS 2) or less (MIS 3) ice cover. Conventionally, the beginning of MIS 4 corre˜ i, 2007) and the beginning of MIS 3 sponds to D-O 19 (Sa´nchez Gon corresponds to the base of D-O 17 (Voelker, 2002; Shackleton et al., 2004). The definition of the beginning of MIS 2 is more problematic: Voelker (2002) place this boundary before D-O 4, Shackleton et al. (2004) indicate that the boundary occurs between D-O 4 and D-O 3 (which is consistent with the boundary implied by Martinson et al., 1987), while Andersen et al. (2006) place the
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boundary during/at the end of D-O 3. From a mechanistic point of view, and given the similarity of the two D-O cycles, it makes very little sense to place the division between them, and we suggest this issue needs to be re-examined carefully. Here, however, we adopt the consensus view that the boundary occurs between D-O 4 and D-O 3. There are several versions of the chronological limits of the MIS in the literature. In early versions (e.g. Duplessy et al., 1981), the chronology was expressed in 14C ages. Although this source of confusion is removed in later chronologies, discrepancies remain (Table 2) because the isotopic record from marine environments is influenced by factors other than global ice volume, including, e.g., salinity and water temperature (Shackleton and Opdyke, 1973); the influence of these other factors is more marked when the isotopic records are derived from planktonic organisms. The use of stacked isotopic records goes some way to minimise the influence of salinity and water temperature (Martinson et al., 1987), especially if the stacked curves are derived from benthic organisms. Complications also arise because of dating uncertainties, different assumptions about the size of the marine carbon reservoir (Waelbroeck et al., 2001) and the synchronicity of registration of events in different environments, and the application of different calibration techniques. Here we use the GICC05 chronology to define the limit of MIS4/3 at 59.4 ka and MIS3/2 at 27.8 ka (Table 2), consistent with the timings for D-O cycles given in Wolff et al., (2010). Table 2 also shows the limits of the Last Glacial Maximum (LGM), consistent with the definitions used for other palaeoclimatic and palaeoenvironmental syntheses (e.g. MARGO, 2009). As with other chronostratigraphic terms defined here, the global ice maximum does not necessarily coincide with regional glacial maxima or the coldest phases of the last climatic cycle (e.g. Turon et al., 2003; Eynaud et al., 2009). The papers in this special issue document the regional changes in vegetation (Fletcher et al., 2010; Takahara et al., 2010; Jime´nezMoreno et al., 2010; Heßler et al., 2010) and fire regimes (Daniau et al., 2010) associated with the chronostratigraphic intervals of the GI, GS and HS during the Last Glacial. This is the first global synthesis of terrestrial vegetation records during the Last Glacial. Despite all the uncertainties associated with site type, pollen source regions, taphonomy, the construction of age models, or the interpretation of individual records, it is clear that vegetation and fire regimes varied on millennial timescales through the glacial. The vegetation records show that events seen in Greenland or in the North Atlantic are not necessarily registered everywhere in the world. Even when a particular event is registered globally, the climate change is expressed differently in different regions: some events are registered as, e.g., a cooling in one region and an increase
Table 1 Definition of the boundaries of Heinrich Stadials (HS).
HS1 HS2 HS3 HS4 HS5 HS6
Top Base Top Base Top Base Top Base Top Base Top Base
Elliot et al. (2001, 2002)
˜i Sa´nchez Gon et al. (2008)
This paper (GICC05 chronology)
Age (14C yr BP)
Age (ka)
Age (ka)
Duration (ka)
13.4 15.1 20.4 22.1 26.1 27.4 33.9 34.9 n/a n/a n/a n/a
n/a n/a n/a n/a n/a n/a n/a n/a 47.0 50.0 60.1 63.2
15.6 18.0 24.3 26.5 31.3 32.7 38.3 40.2 47.0 50.0 60.1 63.2
Relationship with GI/GS (see Wolff et al., 2010)
Length of HS compared to GS
2.4
After GI 2; coincident with GS 1/2
Shorter
2.2
After GI 3; coincident with GS 2/3
Shorter
1.4
After GI 5; coincident with GS 4/5
Shorter
1.9
After GI 9; coincident with GS 8/9
z same length
3.0
After GI 13; coincident with GS 12/13
Unclear
3.1
After GI 18; coincident with GS 17/18
Unclear
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Table 2 Stage boundaries as defined by various authors. Limits originally defined in 14C years (yr BP) have been rendered in calendar years (ka) for comparison, after taking the marine reservoir age into account. GICC05; Svensson et al., 2006, 2008
Duplessy et al., 1981
Martinson et al., 1987
Voelker, Shackleton 2002 et al., 2004
Lisiecki and Raymo, 2005
Peltier and Fairbanks, 2006
Stage1/2
14.7 ka
n/a
14.0 ka
n/a
23.5 ka 27.2 ka 27.8 ka
n/a n/a 29.0 ka
n/a n/a 29.0 ka
n/a n/a n/a
19.0 ka 26.0 ka n/a
Stage 3/4 Stage 4/5 Stage 5/6
59.4 ka n/a n/a
n/a n/a n/a
12.05 kyr BP 13.85 ka n/a n/a 24.11 kyr BP 28.86 ka 58.96 ka 73.91 ka 129.84 ka
n/a
End LGM Onset LGM Stage 2/3
13.0 kyr BP 14.6 ka n/a n/a n/a
59.0 ka n/a n/a
59.0 ka n/a n/a
n/a n/a 130.0 ka
n/a n/a n/a
in plant-available moisture in another. The geographic patterns of vegetation and climate change provide indications of potential mechanisms for the generation and transmission of millennialscale variability. The global maps of vegetation and inferred climate changes presented here (Harrison et al., 2010) will provide a wealth of targets for future model exploration (Kageyama et al., 2010) of millennial-scale climate variability.
Acknowledgements The papers in this issue were conceived and developed at a series of workshops sponsored by the UK QUEST (Quantifying Uncertainty in the Earth System) programme of NERC in support of the QUEST project DESIRE (Dynamics of the Earth System and the Ice-core REcord) and are a contribution to the ongoing work of the QUEST Working Group on Abrupt Climate Change. We are grateful to QUEST for their long-term support of the Working Group on Abrupt Climate Change. SPH was supported during the final production of the issue by QUEST and through a Professorial Fellowship at the University of Bordeaux. We thank Pat Bartlein and Eric Wolff for their insightful comments on the original manuscript, and our colleagues in the QUEST Working Group for discussions that clarified our thinking on the terminological issues. This is Bordeaux 1 University, UMR-CNRS 5805 EPOC Contribution n 1857.
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˜ i* Maria Fernanda Sanchez Gon EPHE, UMR-CNRS 5805 EPOC, Universite´ Bordeaux 1, Avenue des Faculte´s, 33405 Talence, France * Corresponding author. E-mail address: [email protected] Sandy P. Harrison School of Geographical Sciences, University of Bristol, Bristol, BS8 1SS, UK E-mail address: [email protected] 11 November 2009
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Introduction
Millennial-scale climate variability and vegetation changes during the Last Glacial: Concepts and terminology
The Late Quaternary, and more particularly the Last Glacial (Marine Isotope Stages 4, 3 and 2, defined here as 73,500–14,700 calendar years, 73.5–14.7 ka; ka is used solely for calendar years and kyr BP refers to 14C ages), was characterised by millennial-scale climate oscillations of irregular periodicity. The onset of these oscillations was abrupt, with most of the change in climate accomplished within 10–200 years (Steffensen et al., 2008) and the magnitude of the change was large (of the order of 8–15 C in Greenland) (Huber et al., 2006). Two types of rapid climate changes have been described: Dansgaard–Oeschger (D-O) cycles (Dansgaard et al.,1984), associated with abrupt warming and subsequent cooling in Greenland, and cold phases associated with the formation of ice-rafted debris (IRD) deposits (Heinrich layers) (Heinrich, 1988) in the North Atlantic. Dansgaard–Oeschger cycles are clearly registered in the Greenland ice-core record (e.g. Johnsen et al., 1992; North GRIP Members, 2004) and the traces of both of these climate oscillations are recorded in a variety of marine and terrestrial records worldwide (e.g. Bond et al., 1993; Allen et al., 1999; Wang et al., 2001; Gonzalez et al., 2008). The geographic pattern of registration of individual oscillations ˜ i et al., 2008), and the magnitude, nature and length of (Sanchez Gon the component phases of each recorded oscillation, appear to vary (Johnsen et al., 1992). However, documentation of regional changes has been hampered by problems of the synchronisation of individual chronologies, and our understanding of the mechanisms underlying these climate oscillations is still far from complete. Analysis of the mechanisms and impacts of large and rapid climate changes in the past is given additional impetus by the possibility that such events might occur in the future. Although the precise causes are different, investigation of the impact of the warming events at the beginning of D-O cycles or of iceberg melting during Heinrich intervals on the Meridional Overturning Circulation (MOC) in the North Atlantic speaks directly to the impact of future changes in the MOC on regional climate. Coupled ocean– atmosphere model simulations show a reduction of the MOC during the 21st century, in some cases by up to 50%, as a consequence of greenhouse-gas-induced polar warming (Gregory et al., 2005). Simulations with simpler climate models have shown that complete shutdown of the MOC can occur if the slowdown reaches a crucial threshold (which differs in different models) (Stocker and Schmittner, 1997; Stouffer and Manabe, 2003; Stouffer et al., 2006). Clearly, the coupled models do not reach the critical threshold as a result of the gradual change in greenhouse-gas forcing during the 21st century but might do so if the additional forcing due to even partial melting of the Greenland ice sheet were taken into account. 0277-3791/$ – see front matter Ó 2009 Elsevier Ltd. All rights reserved. doi:10.1016/j.quascirev.2009.11.014
There have been several recent attempts to synthesise millennial-scale climate change during the glacial (e.g. Broecker and Hemming, 2001; Alley et al., 2002; van Andel, 2002; Voelker, 2002), but most focus on marine and ice-core records. In this issue, we focus on documenting regional changes in vegetation indicated by pollen records from both marine and terrestrial cores (Fletcher et al., 2010; Takahara et al., 2010; Jime´nez-Moreno et al., 2010; Heßler et al., 2010). There are multiple reasons why this is timely. First, there has been a very rapid increase during the last decade in the number of pollen records with high temporal resolution. However, there is no global compilation of the pollen data, nor a synthesis of vegetation reconstructions based on these data. Second, the development of a new and coherent ice-core reference chronology (GICC05: Svensson et al., 2006, 2008; see also Wolff et al., 2010) makes it possible to achieve a better synchronisation between documented changes in Greenland and the pollen records. Third, a regional synthesis of charcoal records from North America (Marlon et al., 2009) shows that fire regimes respond to abrupt climate changes during the last deglaciation – but little is known about the response of fire globally to millennial-scale climate variability and associated vegetation changes during earlier intervals (Daniau et al., 2010). Fourth, investigation of the impact of changes in vegetation cover, including wetland extent, and in fire regimes is important for understanding the rapid and extremely large (up to ca 200 ppb) changes in methane during D-O cycles (Blunier and Brook, 2001; Flu¨ckiger et al., 2004) and the potential climate feedback. Finally, modelling groups have recently begun to explore the impact of, e.g., changes in freshwater forcing under glacial conditions on regional climates (e.g. Crucifix et al., 2001; Ganopolski and Rahmstorf, 2001; Claussen et al., 2003; Knutti et al., 2004; Flu¨ckiger et al., 2006, 2008; see also Kageyama et al., 2010) but more detailed documentation of observed changes is required for the evaluation of these experiments. A multiplicity of terms is used in discussing rapid climate changes and millennial-scale climate variability during the glacial, and this has led to some confusion particularly in relating records from different regions. Alley et al. (2002), in a definition adopted by the IPCC (Meehl et al., 2007), define abrupt climate change as one that takes place more rapidly than the underlying forcing, pointing out that this kind of behaviour can only occur when the climate system crosses a critical threshold defining the limit between two different climate states. This definition provides a theoretical basis for understanding abrupt climate changes (e.g. Kageyama et al., 2010) but the rapidity of the climate change is
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a function of the temporal scale characteristic of the specific forcing involved and the definition may be difficult to apply in the case of geological records where the nature of the forcing is a priori unknown. For practical reasons, many authors therefore identify abrupt climate changes in geological records in terms of some combination of magnitude of the change and the rapidity with which it is accomplished (e.g. McManus et al., 1999; Martrat et al., 2004). Here, we continue to use the term abrupt climate change in the sense of Alley et al. (2002), but refer to changes in climate that take place in <200 years and, in magnitude, exceed the decadal variability typical of the interval in which they occur as a rapid climate change. Dansgaard–Oeschger (D-O) cycles are an example of a rapid climate change, characterised in Greenland by a marked warming followed by a cooling. In most cases, the cooling occurs in three phases: initially relatively slow, then more precipitous, and a final slow phase. Here we refer to the initial rapid warming as the D-O event (Fig. 1). This event and the first slow phase of the subsequent cooling is referred to as the Greenland Interstadial (GI: Svensson et al., 2006; note this was originally abbreviated to GIS, see North GRIP Members, 2004); the final phase of cooling is referred to as the Greenland Stadial (GS: Svensson et al., 2006); the transition between the GI and GS is placed during the rapid phase of cooling (Fig. 1). Generally, the warming takes place in ca 60 years but the cooling can take more than 2000 years (e.g. D-O 12). However, there are D-O cycles in which initial warming persists for ca 100 years and the cooling occurs almost as quickly as the initial warming (e.g. D-O 6). Where there is a marked phase of rapid cooling, the limit between the GI and GS is relatively easy to define, but there are many D-O cycles where the location of transition between GI and GS is more subjective and can only be defined in terms of intervals when temperatures are equivalent to the general ‘‘glacial’’ baseline level (Wolff et al., 2010). The term ‘‘stadial’’ corresponds to cold intervals marked not by global but by local ice readvances (e.g. Lowe and Walker, 1984).
Fig. 1. Structure of exemplar D-O cycles, showing the rapid warming, and the intervals corresponding to the Greenland Interstadial (GI) and the Greenland Stadial (GS).
However, the term ‘‘interstadial’’ was originally defined in terms of vegetation development in Europe, and specifically to denote periods of warming that were either too short or of insufficient magnitude to lead to the development of deciduous forest in northern Europe (Fletcher et al., 2010). The term ‘‘interstadial’’ and, by extension ‘‘stadial’’ have subsequently been applied in other regions to indicate climate conditions associated with different types of vegetation development (e.g. Tzedakis et al., 2002). This climate-related terminology continues to be used in the palynological literature, and in cases where the pollen records are not precisely dated this can lead to confusion with the chronostratigraphic terms GI and GS. Here, and we would suggest in future discussions of pollen records, we confine the use of ‘‘interstadial’’ and ‘‘stadial’’ to those temporal intervals defined by the GI and GS – whatever the nature of the accompanying vegetation changes or indeed whether these chronostratigraphic intervals are accompanied by changes in vegetation or not. This is particularly important for documentation of millennial-scale variability in the southern hemisphere (Heßler et al., 2010). As has been clearly demonstrated, Antarctic temperatures cool during warm intervals in Greenland and warm during cooling phases in Greenland. The warm phase in the Antarctic leads D-O warming events by ca 1–2.5 ka. The warm phases were originally referred to as Antarctic warmings (Blunier et al., 1998) but the term Antarctic Isotopic Maxima (AIM: EPICA community members, 2006; Wolff et al., 2010) is preferable to avoid any possible implication that these are synchronous or necessarily mechanistically linked with the GI. The cold intervals associated with Heinrich layers in the North Atlantic (Heinrich, 1988) are a second example of rapid climate change. Again, there is some confusion in the literature between terms that are related to the observations and those that describe the climate changes associated with these observations. Thus, the term Heinrich Event (HE) was originally defined as corresponding to the period of time synchronous with the deposition of the Heinrich layer (e.g. Bond and Lotti, 1995), but subsequently applied to cover the whole of the associated cold interval (e.g. Bard et al., 2000). In some regions of the North Atlantic, the cold interval is indeed coincident with continuous deposition of IRD. However, further south, the deposition of IRD occupies a short interval of time and occurs at the end of the cold phase (e.g. off the SW Iberian Peninsula: Naughton et al., 2009) or there are low levels of IRD throughout but with a pronounced maximum of deposition towards the end of the cold phase (e.g. off the NW Iberian Peninsula, offshore from the SE USA: Naughton et al., 2009). Here (Fig. 2), we use the term HE to refer to the interval of the formation of the Heinrich layer in a given region and the cold interval associated with an HE as the Heinrich Stadial (HS: Barker et al., 2009). This usage is consistent with the terminology employed by, e.g., Elliot et al. (2002) and Hemming (2004), who referred to the whole of the cold period as a Heinrich Event because of their focus on the northern North Atlantic where the deposition of IRD occurs throughout the cold phase (Fig. 2). We use the term HS as a chronostratigraphic term, defined as the maximum length of the interval of IRD deposition based on a composite of marine records from the North Atlantic (Elliot et al., 1998, 2001, 2002; Hemming, 2004; Table 1). In some regions, e.g. the tropics (Heßler et al., 2010), the HS is registered as an increase in plant-available moisture rather than a cooling. Indeed, in some regions of the Southern Hemisphere, the chronostratigraphic interval corresponding to HS is not marked by strong vegetation changes (Kershaw personal communication). The deposition of Heinrich layers is associated with certain of the D-O cycles (e.g. HS 5 with the cold phase of D-O 13, i.e. GS 12/13) (Table 1). In most cases, however, the HS (as defined by
Introduction / Quaternary Science Reviews 29 (2010) 2823–2827
Fig. 2. Structure of an exemplar Heinrich Stadial, showing the relationship between sea-surface temperature and deposition of Heinrich layers, and hence the coincidence or otherwise between Heinrich Events (HE) and Heinrich Stadials (HS), at a typical northern and a typical southern North Atlantic site.
cold sea-surface temperatures in the North Atlantic) is shorter than the cold interval in Greenland (as defined by the isotopic record). For example, the estimated length of HS2 is 26.5–24.3 (Elliot et al., 2002) but GS2/3 lasts from 27.48 to 23.34. In the case of HS4, however, the two cold phases appear to be of approximately the same length. Not all D-O cycles are associated with HE. The rapid climate changes associated with D-O cycles or HS are superimposed on the longer-term climate trends within the glacial defined by isotope records from marine cores. These Marine Isotope Stages (MIS) correspond to intervals of more (MIS 4, MIS 2) or less (MIS 3) ice cover. Conventionally, the beginning of MIS 4 corre˜ i, 2007) and the beginning of MIS 3 sponds to D-O 19 (Sa´nchez Gon corresponds to the base of D-O 17 (Voelker, 2002; Shackleton et al., 2004). The definition of the beginning of MIS 2 is more problematic: Voelker (2002) place this boundary before D-O 4, Shackleton et al. (2004) indicate that the boundary occurs between D-O 4 and D-O 3 (which is consistent with the boundary implied by Martinson et al., 1987), while Andersen et al. (2006) place the
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boundary during/at the end of D-O 3. From a mechanistic point of view, and given the similarity of the two D-O cycles, it makes very little sense to place the division between them, and we suggest this issue needs to be re-examined carefully. Here, however, we adopt the consensus view that the boundary occurs between D-O 4 and D-O 3. There are several versions of the chronological limits of the MIS in the literature. In early versions (e.g. Duplessy et al., 1981), the chronology was expressed in 14C ages. Although this source of confusion is removed in later chronologies, discrepancies remain (Table 2) because the isotopic record from marine environments is influenced by factors other than global ice volume, including, e.g., salinity and water temperature (Shackleton and Opdyke, 1973); the influence of these other factors is more marked when the isotopic records are derived from planktonic organisms. The use of stacked isotopic records goes some way to minimise the influence of salinity and water temperature (Martinson et al., 1987), especially if the stacked curves are derived from benthic organisms. Complications also arise because of dating uncertainties, different assumptions about the size of the marine carbon reservoir (Waelbroeck et al., 2001) and the synchronicity of registration of events in different environments, and the application of different calibration techniques. Here we use the GICC05 chronology to define the limit of MIS4/3 at 59.4 ka and MIS3/2 at 27.8 ka (Table 2), consistent with the timings for D-O cycles given in Wolff et al., (2010). Table 2 also shows the limits of the Last Glacial Maximum (LGM), consistent with the definitions used for other palaeoclimatic and palaeoenvironmental syntheses (e.g. MARGO, 2009). As with other chronostratigraphic terms defined here, the global ice maximum does not necessarily coincide with regional glacial maxima or the coldest phases of the last climatic cycle (e.g. Turon et al., 2003; Eynaud et al., 2009). The papers in this special issue document the regional changes in vegetation (Fletcher et al., 2010; Takahara et al., 2010; Jime´nezMoreno et al., 2010; Heßler et al., 2010) and fire regimes (Daniau et al., 2010) associated with the chronostratigraphic intervals of the GI, GS and HS during the Last Glacial. This is the first global synthesis of terrestrial vegetation records during the Last Glacial. Despite all the uncertainties associated with site type, pollen source regions, taphonomy, the construction of age models, or the interpretation of individual records, it is clear that vegetation and fire regimes varied on millennial timescales through the glacial. The vegetation records show that events seen in Greenland or in the North Atlantic are not necessarily registered everywhere in the world. Even when a particular event is registered globally, the climate change is expressed differently in different regions: some events are registered as, e.g., a cooling in one region and an increase
Table 1 Definition of the boundaries of Heinrich Stadials (HS).
HS1 HS2 HS3 HS4 HS5 HS6
Top Base Top Base Top Base Top Base Top Base Top Base
Elliot et al. (2001, 2002)
˜i Sa´nchez Gon et al. (2008)
This paper (GICC05 chronology)
Age (14C yr BP)
Age (ka)
Age (ka)
Duration (ka)
13.4 15.1 20.4 22.1 26.1 27.4 33.9 34.9 n/a n/a n/a n/a
n/a n/a n/a n/a n/a n/a n/a n/a 47.0 50.0 60.1 63.2
15.6 18.0 24.3 26.5 31.3 32.7 38.3 40.2 47.0 50.0 60.1 63.2
Relationship with GI/GS (see Wolff et al., 2010)
Length of HS compared to GS
2.4
After GI 2; coincident with GS 1/2
Shorter
2.2
After GI 3; coincident with GS 2/3
Shorter
1.4
After GI 5; coincident with GS 4/5
Shorter
1.9
After GI 9; coincident with GS 8/9
z same length
3.0
After GI 13; coincident with GS 12/13
Unclear
3.1
After GI 18; coincident with GS 17/18
Unclear
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Table 2 Stage boundaries as defined by various authors. Limits originally defined in 14C years (yr BP) have been rendered in calendar years (ka) for comparison, after taking the marine reservoir age into account. GICC05; Svensson et al., 2006, 2008
Duplessy et al., 1981
Martinson et al., 1987
Voelker, Shackleton 2002 et al., 2004
Lisiecki and Raymo, 2005
Peltier and Fairbanks, 2006
Stage1/2
14.7 ka
n/a
14.0 ka
n/a
23.5 ka 27.2 ka 27.8 ka
n/a n/a 29.0 ka
n/a n/a 29.0 ka
n/a n/a n/a
19.0 ka 26.0 ka n/a
Stage 3/4 Stage 4/5 Stage 5/6
59.4 ka n/a n/a
n/a n/a n/a
12.05 kyr BP 13.85 ka n/a n/a 24.11 kyr BP 28.86 ka 58.96 ka 73.91 ka 129.84 ka
n/a
End LGM Onset LGM Stage 2/3
13.0 kyr BP 14.6 ka n/a n/a n/a
59.0 ka n/a n/a
59.0 ka n/a n/a
n/a n/a 130.0 ka
n/a n/a n/a
in plant-available moisture in another. The geographic patterns of vegetation and climate change provide indications of potential mechanisms for the generation and transmission of millennialscale variability. The global maps of vegetation and inferred climate changes presented here (Harrison et al., 2010) will provide a wealth of targets for future model exploration (Kageyama et al., 2010) of millennial-scale climate variability.
Acknowledgements The papers in this issue were conceived and developed at a series of workshops sponsored by the UK QUEST (Quantifying Uncertainty in the Earth System) programme of NERC in support of the QUEST project DESIRE (Dynamics of the Earth System and the Ice-core REcord) and are a contribution to the ongoing work of the QUEST Working Group on Abrupt Climate Change. We are grateful to QUEST for their long-term support of the Working Group on Abrupt Climate Change. SPH was supported during the final production of the issue by QUEST and through a Professorial Fellowship at the University of Bordeaux. We thank Pat Bartlein and Eric Wolff for their insightful comments on the original manuscript, and our colleagues in the QUEST Working Group for discussions that clarified our thinking on the terminological issues. This is Bordeaux 1 University, UMR-CNRS 5805 EPOC Contribution n 1857.
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˜ i* Maria Fernanda Sanchez Gon EPHE, UMR-CNRS 5805 EPOC, Universite´ Bordeaux 1, Avenue des Faculte´s, 33405 Talence, France * Corresponding author. E-mail address: [email protected] Sandy P. Harrison School of Geographical Sciences, University of Bristol, Bristol, BS8 1SS, UK E-mail address: [email protected] 11 November 2009