Odekon M. Encyclopedia of paleoclimatology and ancient environments

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by international accord is 10,000 radiocarbon yBP or about

11,600 calendar y

BP. However, the Wisconsinan Glaciation in

the form of the LIS still covered vast areas of northern North

America at this boundary time, and in fact, glacier ice associated

with the geologic-climate unit still exists in the form of the

Barnes Ice Cap on central Baffin Island, Nunuvit, Canada. The

boundary for the transition between the last interglacial stage

and the Wisconsinan Stage (Figure W1) can either be set to

include the need for ice growth during MIS 5d and 5b, or the

sustained interval of global ice expansion during MIS 4.

The Wisconsinan Glaciation is, however, of global impor-

tance because the LIS at its maximum at around 22,000 cal

BP (the Last Glacial Maximum or LGM, Figures W1 and W2)

(see Last Glacial Maximum) was similar to the present Antarctic

Ice Sheet in terms of area and volume. Hence, of the approxi-

mate 120 m of equivalent fall in global sea level (Figure W1),

some 70 m is attributed to the LIS (Denton and Hughes,

1981; Lambeck and Chappell, 2001). Actually, the terms Wis-

consinan Glaciation or Weichselian Glaciation (NW Europe)

are frequently used as general terms for the last glacial cycle sen-

sus lato. However, this statement begs the question of when did

the “Wisconsinan” Glaciation of North America really com-

mence? The term “Wisconsinan” above is in quotations because

stratigraphic practice in North America variously defines the

“Wisconsinan” Glaciation as beginning about 100 cal ka year

ago or at the MIS 5/4 boundary ca. 75 cal ka (Figure W1). These

temporal boundaries are not numerically dated but inferred

from the stable oxygen isotope records preserved in deep sea

sediments.

The Wisconsin an Glaciation and the marine isotope

and sea level records

The overall structure of the last glaciation at a global scale is

most frequently determined from a combination of studies of

deep sea benthic foraminifera d

O variations (Shackleton, 1987),

and evidence for past sea levels on tectonically active coast-

lines (Chappell et al., 1996). The chronology of these events

is based on a combination of radiocarbon and U-series dating

(see Radiocarbon dating), and fitting (tuning) the data to

variations in orbital insolation (see Astronomical theory of

climate change). Both methods explicitly track the global hydro-

logical balance and seek to determine the volume (global) of ice

located on land areas. Figure W1 illustrates the history of global

ice sheet growth and melt from marine isotope stage (MIS) 5

to 1 (130–10 cal ka).

Figure W1 Estimated global changes in sea level over the last 140,000 years associated with the growth and retreat of the integrated global

ice volume. The two columns to the right of the figure show alternative scenarios for Wisconsinan Glaciation and the boundary between the

Pleistocene and Holocene. E, M, and L standard for early, middle and late Wisconsinan. Sang. is the Sangamonian Stage (Interglaciation), and Ill.

is the penultimate Illinoian Stage (Glaciation). LGM is the late glacial (global) maximum. The final column shows the interval over which

radiocarbon dating can be employed.

WISCONSINAN (WEICHSELIAN, WU

RM) GLACIATION 987

Both approaches indicate that sea level was 5 m above

present during the last interglacial around 125 cal ka (MIS 5e)

but then fell dramatically by 115 cal ka (MIS 5d), indicating

that significant ice growth occurred somewhere on the planet.

In terms of volume, this increase in glacier ice amounted to

the equivalent of several present-day Greenland Ice Sheets.

A Greenland Ice Sheet Unit (GISU) amounts to 5 m of global

sea level and the change in MIS 5d is in the range 4–10 GISU.

Global sea level rose and hence ice sheets melted during

MIS 5c and 5a, with an interval of growth during MIS 5b. It

thus appears certain that significant ice sheets were present on

the Earth’s surface during MIS 5, but the exact location and

individual sizes of these ice masses is uncertain at best and

basically unknown.

Onset of glaciation in North America

The question of when the glaciation began is, at present, very

difficult to answer because deposits associated with the early

Wisconsinan Glaciation lie beyond the limits of radiocarbon

dating, i.e., >45 ka

BP (Figure W1). As noted above, a common

approach is to examine the history of changes in global sea

level since the last interglacial, the Sangamonian Interglaciation

in North America. Because the LIS was the dominant factor in

changes of sea level on glacial timescales, the history of global

sea level must therefore represent significant aspects of the

volume history of the ice sheet, but it is probably not a 1:1 cor-

respondence. It is possible that different sectors of this vast ice

sheet experienced different chronologies because of the enor-

mous range of glacial-climatic conditions that must have pre-

vailed over such a massive ice sheet.

At present, the presumed centers of ice sheet growth

and dispersal of the LIS are located in the vast uplands of

the Eastern Canadian Arctic and Keewatin (Williams, 1979),

which today have mean annual temperatures of between –13

and –5



C, with winter snowfalls ranging between 20 and

50 cm water equivalent. On Baffin Island, large areas of

lichen/vegetation-free surfaces, caused by snow-kill, suggest

that permanent snow/ice covered a large fraction on the region

even as recently as the Little Ice Age, approximately 300–400

year ago. Significantly, the marine isotope data (Figure W1)

indicate a very rapid increase in ice volume from MIS 5e to

5d. Such a rapid growth of ice sheets is incompatible with

the slow growth associated with decreased temperatures (which

would also decrease the precipitation) and therefore several

authors have linked the onset of LIS glaciation with increased

winter snowfall (Johnson and Andrews, 1979). These models

Figure W2 Extent of the ice sheets over North America and Greenland around 18,000 years ago. The regional glacial maxima vary around the

ice sheet (see text). The solid black area represents the mapped extension of the Laurentide Ice Sheet during the late Wisconsinan Glaciation

along the eastern margin of the ice sheet based on the 1987 compilations of Dyke and Prest (see Dyke et al., 2002).

988 WISCONSINAN (WEICHSELIAN, WU

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call for increased cyclonic activity in Baffin Bay, probably as

the result of increased advection of warm modified Atlantic

Water via the West Greenland Current. According to such a

scenario, the development of ice over the islands and channels

of the High Canadian Arctic would cut off the outflow of cold,

“fresh” Polar Waters into Baffin Bay, resulting in a reduction in

the extent and severity of sea-ice, in turn leading to increased

fall precipitation over the uplands of Baffin Island and

Labrador-Newfoundland. Such a paleoceanographic/climatic

scenario is supported by the stratigraphy and molluskan faunas

for glacial events along the eastern seaboard of Baffin Island

that now lie beyond the limits of glaciation. Based on amino acid

racemization constraints, these events are probably >75 cal ka

and might be correlated with intervals of ice growth in MIS 5b

and 5d (Figure W1).

However, even with such a “snowfall” model, Andrews and

Mahaffy (1976) had great difficulty in driving a 3-D glaciologi-

cal model to accommodate around 20 m of global sea level

lowering in the 10,000 year required by the oxygen isotope

data. However, this and later models all show the nucleation

of ice sheets over the Canadian Arctic eastern uplands and

growth across Hudson Strait. This must have resulted in the

ponding of a large lake within Hudson Bay and the possibility

of large outburst floods or subglacial release of water resulting

from such a configuration.

The center of the Wisconsinan Ice Sheet – the record

from Hudson Bay

The geographic center of the LIS during the Wisconsinan Stage

was the area of Hudson Bay ( Figure W2). Hence, the LIS is a

marine-based ice sheet, defined as an ice sheet with a substan-

tial fraction of its area located below sea level, with a major

drainage directed through Hudson Strait into the Labrador Sea.

Because of the density difference between ice (ca 900 kg m

–3

)

versus marine water (ca 1,028 kg m

–3

), ice margins ending in

the sea, or a lake, have the potential to “lift off,” resulting in a

rapid increase in ice flow and the collapse of ice centers. The

LIS thus had the potential to experience extensive deglaciation

due to massive calving and collapse – this certainly occurred

at the end of the glaciation some 8–9 cal ka but did it occur

earlier in the last glacial cycle?

The late Quaternary stratigraphy of the Hudson Bay Lowlands

(SW area of Hudson Bay) is beautifully preserved along hun-

dreds of km of river section (Andrews et al., 1983;Thorleifson

et al., 1992)(Figure W3). Although well preserved, the gla-

ciological interpretation of the sequence is complicated by:

(a) problems of assigning ages to materials at or beyond

the limits of radiocarbon dating, and (b) differing interpreta-

tions of the genesis of rippled sands and laminated silts.

Amino acid racemization data on marine mollusks, which

are a product of both time and temperature, show that the

presumed last interglacial marine sediments of the Bell Sea

have ratios of around 0.24 compared with ratios of 0.03

for the early Holocene Tyrrell Sea. Intermediate ratios of

0.12–0.14 were obtained from shells in sediments associated

with the Prest Sea. Radiocarbon dates on the Prest Sea shells

give finite dates between 35 and 40 ka

BP and an age of ca.

40 cal ka is compatible with a reasonable thermal history for

the site. It is possible that partial deglaciation occurred after

Heinrich event 4 at around 38 cal ka (see below).

The glacial chronology of the Hudson Bay Lowlands is a

critical but disputed element in our knowledge of the dynamics

of the LIS.

Wisconsinan Ice Sheet instability – Heinrich events

Our understanding of ice sheet behavior during the Wisconsinan

Glaciation has undergone substantial changes over the last few

decades. In the 1970s and 1980s the focus was on the relation-

ships between ice sheet mass balance, climate, and the Milanko-

vitch orbital variations as the “pace maker” of the ice ages (Hays

et al., 1976). Thus, thinking was restricted to cycles of ice sheet

growth and retraction at 20, 40, or 100 ka insolation cycles.

However, parallel with this paradigm was the realization (largely

from theory) that marine-based ice sheets should be, or might be,

unstable. This thinking was largely directed to the present and

future history of the West Antarctic Ice Sheet. However, Hein-

rich (1988) presented a coherent argument for significant, but

short-lived, ice-rafted detritus (IRD) events in the North Atlantic.

This initial paper was followed by a deluge of studies that docu-

mented and tried to explain what are now universally referred to

as Heinrich or H-events (Broecker et al., 1992; Andrews, 1998)

(see Heinrich events).

Heinrich events represent 500–1,000-year intervals when

the normal influx of IRD (see ice-rafted debris) into the North

Atlantic was dramatically elevated. A key component of the

Heinrich events is detrital carbonate. Using this as a tracer,

plus studies of the isotopic composition and ages of the non-

carbonate fraction, indicates that the primary source of the

IRD in Heinrich-sediments is from Hudson Strait, demonstrat-

ing the inherent instability of marine-based ice sheets.

The elusive Middle Wisconsinan

The Middle Wisconsinan generally encompasses MIS3 or

between 55 and 25 cal ka. Global sea level was higher than

during the main global glacial intervals of MIS4 and 2, but still

well below the interglacial levels of MIS5e and 1 (Figure W1).

Ice sheet instability characterized some of this time (see above).

A major problem is that these ages lie beyond 5 and 9 half-

lives of the most-commonly used dating method,

C dating

of wood, peats, mollusks, foraminifera, etc. (Figure W1). Con-

tamination can frequently result in finite dates on material that

show infinite ages, and temporal correlations are often unreli-

able as the stratigraphy of terrestrial deposits is difficult to trace

unambiguously from outcrop to outcrop, often compounded by

the absence of marker beds (such as tephra beds) around the

margins of the LIS.

A survey of the glacial and non-glacial stratigraphy for north-

ern North America (Andrews, 1987) indicated that middle

Wisconsinan Glaciation (MIS3) is absent from the mountains

of Alberta and British Columbia (there appears to have been

no middle Wisconsinan Cordilleran Ice Sheet), and moreover

there is no evidence that the Prairie provinces were covered by

an ice sheet (Figure W3). Similarly, in the Yukon Territory and

the mountains of Alaska, the Boutellier non-glacial interval

represents a long, middle Wisconsinan interval with reduced

glacial cover when compared with the expansion of glaciers

and ice sheets during MIS2. However, in the area of the Great

Lakes, glacial and interstadial deposits exist, dated from >45

to ca. 25 cal ka. The Plum Point interstadial was followed

by the LGM.

During MIS3, the LIS was restricted to the hard bed (glacio-

logically speaking) of the Canadian Shield, with limited

advances and retreats along the southern margin (Figure W3).

The location of the ice margin around two third of its perimeter

during MIS3 is difficult to define at present. Cosmogenic dat-

ing and deep-sea sediment records suggest a series of rapid

WISCONSINAN (WEICHSELIAN, WU

RM) GLACIATION 989

expansions and contractions especially as seen in Heinrich

events 4 and 3, which occur within MIS3. In Baffin Bay,

equivalent but possibly lagging IRD-rich units also are found,

suggesting that the northern channels and ice streams in the

Canadian Arctic Archipelago displayed a glaciological beha-

vior similar to that of the Hudson Strait ice stream.

Big ice versus small ice at the LGM

There is a long-running dispute over the area and volume of the

ice sheets during the LGM phase of the Wisconsinan Glaciation.

This was detailed in a global analysis by Denton and Hughes

(1981), who highlighted the difference between a “big” ice sheet

model, which placed large ice sheets and connecting ice shelves

over vast areas of the Arctic and the Northern hemisphere, and

a “small” ice world, where ice shelves were restricted and

continental margins frequently terminated on land. In the last

two decades, new tools, new insights into ice sheet behavior

and basal conditions, and the exploration of new areas, have

delineated ice margins that show a mosaic of “big ice” in some

areas, whereas “small ice” appears necessary in other areas.

Few papers deal with evidence for massive ice shelves across

the Arctic Ocean and marginal seas, and the implications for

deep-water iceberg scouring at some sites is frequently ambigu-

ous because of the inability to date the iceberg plough marks.

The ice sheet margin at the LGM has been extended sea-

ward along the Labrador and Baffin Island coastline, largely

on the basis of marine studies and cosmogenic dating (Steig

et al., 1998). The consensus now maintains that an ice sheet

extended across and drained through the Canadian High Arctic

Channels and was contiguous with both the Greenland Ice

Sheet along its eastern margin and the LIS along its southern

limit (Dyke et al., 2002). A large ice sheet occupied the Barents

Sea but the extent of LGM ice sheets further east along the

arctic Eurasian margin is now restricted, and there is no evi-

dence for such an ice sheet on the east Siberian shelf (Svendsen

et al., 1999).

The concept of the Late Glacial Maximum (LGM) – is

it useful?

The literature is now replete with the term “Late Glacial Max-

imum” or LGM (Figure W1). In a global sense, the term

applies to the interval within MIS2 when ice volumes reached

Figure W3 Extent of ice over northern North America and Greenland during the mid-Wisconsinan (Marine Isotope Stage 3) interval. Note the

absence of ice over the Cordilleran region and over the western sector of the Laurentide Ice Sheet. The numerous question marks (?)

indicate the major uncertainties in terms of timing and ice extent.

990 WISCONSINAN (WEICHSELIAN, WU

RM) GLACIATION

a maximum and local sea levels at sites far removed from the

ice sheet were at their lowest. However, the term should not

be taken to imply that all ice sheets and glaciers experienced

a synchronous interval of maximum extent around 22 cal ka.

It is well known that various margins of the North American

ice sheets reached their late glacial maxima at different times.

As an example, the Puget Lobe of the Cordilleran Ice Sheet

advanced south into the Puget Lowland at around 16 cal ka,

which is also the date of the main advance across Des Moines,

Iowa (Denton and Hughes, 1981). Hence the term “LGM”

should not be used indiscriminantly and it should be made clear

whether the term is being used in the global, integrated ice

volume concept (Figure W1), or whether it is being used to

refer to a regional glacial maximum limit that took place at a

specific time within MIS2.

Pattern of deglaciation

The advent of radiocarbon dating in the late 1950s and 19 60s

rapidly led to the acquisition of data on the retreat of the Wis-

consinan Ice Sheet across the northern USA states and most

of Canada (see Laurentide Ice Sheet). The first maps to show

deglaciation isochrones were similar (Bryson et al., 1969).

Subsequent maps show changes in important details, but not

in the overall pattern and timing of retreat (Dyke and Prest,

1987). The isochrones are in 1,000s of radiocarbon years

and hence underestimate the true ages by a few hundred to

1,000–2,000 years. Several authors used the maps to deter-

mine changes in area, and to derive estimates of volume

changes over time.

The main pattern of retreat for the LIS can be summarized

as follows:

1. Rapid deglaciation of the SW and W margins of the ice

sheet. These margins lay on soft, deformable sediments.

2. During the time of rapid retreat of the SW and W margins,

there was little change in the eastern margin facing Labrador

and Baffin Bay.

3. Starting ca. 14 cal ka, the retreating S, SW, W, and

NW margins fre quently wer e adjac ent to massive progla-

cial lakes which drained via several complex routes into

the Gulf of M exico, the shelf off the Hudson River,

the St. Lawrence River, Hudson Strait, or the Mckenzie

RiverinArcticCanada(DykeandPrest,1987;Clark

et al., 2001).

4. By 12 cal ka, the ice sheet had retreated across the St

Lawrence Estuary and lay along the north shore of the

Great Lakes.

5. Around 8–8.5 cal ka, a massive outburst flood linked the

proglacial lakes in southern Hudson Bay and northern

Quebec/Ontario with the encroaching sea in Hudson Strait.

Earlier outburst floods had been directed via the other routes

noted above.

6. By 7.5 cal ka, the last remnants of the LIS and the

Wisconsinan Glaciation were restricted to Baffin Island

and other sites in eastern and northern Canada where mod-

eling implied that the ice sheet had started its growth.

7. Technically speaking, t he geologic-cl imate e vent of

the Wisconsinan Glaciation still persists in terms

of deposits and climatic inferences that pertain to the

Barnes Ice Cap, Nunavit, Canada, hence the obvious

time-transgressive nature of this stratigraphic concept ver-

sus the chronostra tigraphy emb edded in the Wisconsi nan

Stage and the Holocene.

The terms Wisconsinan, Weichselian, and Wu

glaciations

The application of the stratigraphic code (Hedberg, 1976)to

issues of describing, naming, and correlating units within the

Quaternary Era was a matter of considerable concern to terres-

trial glacial stratigraphers from NW Europe, the northern USA,

and Canadian researchers in the 1970s and early 1980s. How-

ever, the 1990s have seen the focus shift to the glacial impact

on marine records, and the climatic significance and regional

and global importance of abrupt climate “events,” such as the

sequence of Heinrich IRD-rich sediment layers, the Younger

Dryas cold event, and the impact of large freshwater outbursts

on the regional to hemispheric climate. As a result, little atten-

tion has been paid to formal definitions and these terms are

being largely used in an ad hoc manner.

Summary

The Wisconsinan Stage and the Wisconsinan Glaciation repre-

sent two different stratigraphic codes, although for many the

differences are unclear. Radiocarbon dating of terrestrial and

marine materials serves to define events and effect correlations

over the last 30–40 ka (Figure W1), but correlations and events

beyond the Middle Wisconsinan lie beyond the limits of radio-

carbon dating and “dating” is frequently based on correlations

to the marine isotope record (Figure W1) or relies on a range

of new dating methods, including cosmogenic exposure ages,

various luminescence methods, and amino acid racemization.

Consequently, as yet there is little absolute knowledge about

the timing of the onset of the last glaciation on North America

or its history of expansion and retraction during MIS5, 4 and

early stage 3 (Figure W1). However, it is known that the LIS

suffered major catastrophic loss of mass during Heinrich events

4 through 1 (ca. 38 cal ka to 16 cal ka) as massive iceberg

armadas carried sediments with provenance tracers indicative

of sources in Hudson Strait and Hudson Bay into the North

Atlantic to sites just west of Portugal and Ireland. How these

events were manifest around the other sectors of the ice sheet

is unclear, especially given the evidence for relative little ice

in the Canadian Prairie provinces and even over the Cordilleran

mountains during mid-Wisconsinan time (Figure W3).

John T. Andrews

Bibliography

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Cross-references

Astronomical Theory of Climate Change

Cordilleran Ice Sheet

Glaciations, Quaternary

Heinrich Events

Ice-Rafted Debris (IRD)

Last Glacial Maximum

Last Glacial Termination

Laurentide Ice Sheet

Late Quaternary Megafloods

Oxygen Isotopes

Proglacial Lacustrine Sediments

Radiocarbon Dating

Scandinavian Ice Sheet

SPECMAP

Sea Level Change, Post-Glacial

Sea Level Change, Quaternary

992 WISCONSINAN (WEICHSELIAN, WU

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YOUNGER DRYAS

Introduction

Abrupt climate changes are events that have affected the Earth

many times; it is not only possible but also likely that these will

affect us in the future (Alley et al., 2002). While gradual climate

changes allow time for those affected to adjust, rapid events

often cause greater impact simply because of their rapidity. While

we do not yet understand past rapid climate events enough to

predict them, we can examine past data and use models to test

our ideas about why they occur. One of the most useful ways

to approach the data is to examine the occurrence, magnitude,

timing, and distribution of these events, fitting together the

pieces of a global puzzle.

Younger Dryas – prime example of abrupt climate

change

The Younger Dryas (YD) is known as the best example of

an abrupt climatic event recorded on land, in the oceans, and in

ice cores (Alley et al., 2002, 2003). It was a cold event that

occurred between 11,000 and 10,000

C years ago, (appro-

ximately 12,800 and 11,600 calendar years ago), and it lasted

about 1,200 years. About 15,000 years ago, as Northern Europe

ice sheets were retreating and trees were beginning to move

northward replacing tundra, several brief returns to cold conditions

suddenly took place. The most pronounced of these reversals,

which could be correlated from place to place in Europe, was

called the Younger Dryas. The name “Younger” was used

because it was the last of these rapid cold events, while earlier

events were called the Oldest Dryas and the Older Dryas (Jansen,

1938). The name Dryas was used because the layer contained

fossil leaves and fruits of the Dryas octopetala plant, a member

of the rose family (Rosaceae) that colonizes arctic-alpine

regions. The layer just below the Younger Dryas that contained

tree fossils was named Allerød for the Scandinavian locale.

After this striking reversal in macrofossils was noted by paleo-

ecologists in Northern Europe in the late nineteenth and early part

of the twentieth century, it was correlated from place to place using

palynology, the study of pollen and spores. Hundreds of pollen

diagrams throughout Europe now show the effect that this event

had on ecosystems and landscapes, causing the demise of forests

and the increase of erosion and expansion of tundra (Watts,

1980) during this cold interval. Radiocarbon dating beginning

in the 1950s secured the correlation of the reversal in Europe.

However, more recent comparison of calendar ages to the radio-

carbon scale indicates that plateaus exist in the radiocarbon scale,

complicating the correlations somewhat.

Over the last twenty years, many scientists have studied

the Younger Dryas in order to better understand rapid climatic

events. The geographic range (Figure Y1), timing, and magni-

tude of this event are crucial to understanding its origin.

Europe

While the timing of the event in Europe was synchronous from

place to place, each locale varied in the specific vegetational

changes that took place. For example, while tundra herbs indi-

cative of open, disturbed landscapes such as Artemisia (worm-

wood) and Chenopodiaceae replaced birch trees and conifers in

the British Isles (Walker, 1994), in Scandinavia the change was

more subtle, with colder tundra herbs replacing some that inhab-

ited warmer environments. Some of the strongest effects of the

Younger Dryas are noted in southern Europe, where summer tem-

peratures may have been 8



C lower than today. In Norway, Birks

andAmmann(2000) estimated mean July temperature to be

about 7–9



C lower than today, and about 2–4



C lower than the

preceding Allerød warm interval. Substantial changes in vegeta-

tion characterized both coastal and inland areas all the way from

northern Norway to the Mediterranean (Walker, 1994).

North America

While originally the Younger Dryas was thought to be solely a

European phenomenon, the advent of accelerator mass spectro-

metry resulted in improved radiocarbon dating and the identifi-

cation of plant macrofossils permitting identification of plants

to the species level. Along with palynology, these vegetational

changes resulted in proof of cooling that was correlative with

the European YD. These results demonstrated the presence of

a Younger Dryas event in the maritime provinces of Canada

(Mayle et al., 1993) and the northeastern coast of the United

States (Peteet et al., 1993). Estimates of the cooling range from

3–4



C in southern New England to 6–7



C in eastern Canada.

In the Maritime Provinces of Canada, shrub tundra was

replaced by herb tundra in some places, while in others shrub

tundra replaced spruce forest. In the mid Atlantic and southern

New England regions, a mixed boreal and deciduous forest was

replaced by boreal forest (spruce, fir, larch, alder, paper birch).

Further investigations continue to document the YD event in

the mid-Appalachian region, Midwestern USA, coastal British

Columbia and Alaska, and Arctic Alaska. Documentation of

the YD event in terrestrial locations as far west as Alaska

revealed that the geographic extent of this cooling was not

limited to the North Atlantic region, and that the forcing was

possibly not limited to the North Atlantic Ocean ( Figure Y1).

South America

Evidence for the YD event in South America is mixed.

Evidence is most convincing in Colombia, where many sites

show a late-glacial palynological oscillation but precise dating

is lacking. In southern South America, while some data shows

a prolonged YD which begins earlier than the North American

event, some data suggests a correlation and some does not.

In Peru, glacial advances do indicate a YD age advance.

Asia

North Pacific investigations now show that the YD was expressed

in the Japan Sea as well as off coastal California (Behl and

Kennett, 1996). While a cold interval on land in Japan has been

linked to the YD, with indications of colder temperatures and

declines in moisture, recent evidence from other annually lami-

nated lakes suggests that the cooling event of about 3



C lagged

the European event by a few centuries (Nakagawa et al., 2003).

The high-resolution speleothem monsoon record from coastal

mid-latitude Hula Cave, China, also clearly shows a YD event

(Wang et al., 2001), as do loess records in the region.

Marine evidence

In the 1980s, North Atlantic faunal changes concurrent with

the terrestrial European signals led ocean paleoecologists to

the conclusion that the North Atlantic polar front re-advanced

to its glacial position at that time (Ruddiman and McIntyre,

1981). Subsequent geochemical analyses indicate that the pro-

duction of North Atlantic Deep Water (NADW), which today

is responsible for bringing heat to Europe, was greatly reduced

at this time, as it was during the ice age (Broecker et al., 1985).

The cause for the presence of this deep-water reduction is still

in dispute, but it has been linked to diversions of Laurentide

meltwater to specific locations in the North Atlantic. Convincing

Figure Y1 Global map of distribution of palynological evidence for the Younger Dryas cooling 11–10

CkyBP (12,800–11,600 calendar years).

Cooling is strongest in regions surrounding the N. Atlantic, including Greenland, Europe, and eastern North America. Cooling has also been

identified in many places far from the N. Atlantic, such as N. Pacific Canada and Alaska, and East Asia. Tropical cooling has also been noted.

Many sites are still controversial, where some evidence points to cooling and other evidence does not show the oscillation (i.e., New Zealand, Chile;

after Peteet, 1995; Peteet et al., 1993).

994 YOUNGER DRYAS

marine evidence from the Atlantic for the YD shift comes from

the Cariaco Basin, where sediment color and other data is inter-

preted as an increase in wind-induced upwelling off coastal

Venezuela (Hughen et al., 1996) or decreased riverine runoff

from adjacent countries.

North Pacific evidence for a YD event comes from the

Santa Barbara Basin (Kennett and Ingram, 1995), the Gulf of

California, the northwestern North Pacific, and even the eastern

equatorial Pacific. An arid interval is correlated with the YD in

the tropical Atlantic (Maslin and Burns, 2000) and off Africa.

Ice cores

The advent of ice core research was very important in demon-

strating the importance of the Younger Dryas in the history of

climate change. Camp Century, Dye 3, and Renland first illu-

strated this change, which was then confirmed in the GRIP and

GISP-2 deep cores (Dansgaard et al., 1993). The signal is observed

as a several per-mil shift in oxygen isotope ratios, corresponding

to about 15



C colder than today (Cuffey et al., 1995). What is

unusual about the ice core shift as compared to pollen records is

that the magnitude of the shift in the ice cores is almost equal to

that of full-glacial conditions, while in pollen records throughout

Europe, this shift is probably at most equal to half the shift in

temperature represented by full-glacial conditions. Thus, the iso-

topes are complicated to interpret, either suggesting locally colder

conditions than on continents adjacent to the North Atlantic, or a

transport or source problem. In the ice cores, the YD is also

marked by much higher dust levels and chemical concentrations

than the previous Allerød or the subsequent Holocene. While the

onset of the YD in the ice cores is gradual, the termination is

very abrupt, occurring in as little as 1–3 years in some indicators.

Snow accumulation rates during the YD were about half that of

the Preboreal, and about one third of modern rates. The changes

at the end of the YD were completed in three 5-yr steps spread

over about 40 years (Taylor et al., 1993), and snow accumulation

changed by about 90% in one year (Alley et al., 1993). The discov-

ery of a YD methane signal in Antarctic ice was the first indication

that this event made a global imprint. However, more recent ice

core evidence for the YD event correlative with the European

and Greenland ice cores is conflicting, suggesting that in some

locations a YD is correlative but in others it precedes the northern

hemisphere change. Ice cores from more temperate locations in

Peru and Bolivia also show evidence of a late-glacial reversal,

which probably is the YD but age control is problematic.

Mechanisms

The decreased NADW production during the YD is the chief

hypothesis among those proposed as the reason for the cooling

throughout the Northern Hemisphere (Broecker et al., 1985). Cold

meltwater from the late-glacial release of Lake Agassiz to the

North Atlantic would have been located in such a position that

it could have slowed down the thermohaline circulation (THC).

The 1,500-yr oscillations in marine, terrestrial and ice cores

throughout the Holocene suggests that YD-type signals may be a

complex recurring climate oscillation that was magnified during

the YD by meltwater from receding ice sheets (Bond and Lotti,

1995). One potential problem with the THC hypothesis for

YD and other similar events is that most atmospheric models

show a restricted circum-North Atlantic response to the YD when

a shutdown is forced (Rind et al., 1986; Renssen, 1997), although

others show a muted cooling in the North Pacific (Mikolajewicz

et al., 1997). Understanding how this mechanism could cause a

significant YD signal around the globe is a major challenge

to modelers. Some recent research has focused on the tropics as

a possible amplifier of abrupt climate change caused by shifts in

the intertropical convergence zone (ITCZ) (Clement et al., 2001).

The impacts of the YD on human populations was probably

extreme, and may have contributed to the beginnings of agri-

culture around the fertile crescent (Moore and Hillman, 1992).

On the North Slope, Alaska, changes in human migration (Kunz

and Reanier, 1996) during the YD are linked to a shift to major

changes in moisture (Mann et al., 2002). Added to the human

factor in North American mammal extinctions, it is possible that

the YD pushed populations over the brink (Peteet et al., 1990).

Abrupt climate changes of lower magnitude than the YD occurred

throughout the Holocene; these have been linked to reduced

solar radiance (Bond et al., 2001). Much is still to be learned

regarding the Younger Dryas event.

Dorothy M. Peteet

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YOUNGER DRYAS 995

Author Index

Abbott, Mark B. 489

Almasi, Peter F. 929

Alverson, Keith 675

Anderson, Lesleigh 489

Andersson, Carin 609

Andrews, John T. 986

Archer, Allen W. 226

Archer, David 478, 533

Arnaud, Emmanuelle 384

Ashley, Gail M. 446

Asmerom, Yemane 916, 969

Avis, Chris 690

Baker, Paul A. 746

Baker, Victor R. 380, 504

Barker, Philip 738

Bassinot, Franck C. 911

Beer, Jürg 92, 211

Beltrami, Hugo 103

Benn, Douglas I. 382, 394, 948

Bennett, Matthew R. 768, 783, 873

Berger, André 51

Berger, Glenn W. 249

Berger, Wolfgang H. 525

Bickert, Torsten 136

Bjørklund, Kjell R. 869

Bond, Gerard 568, 929

Bosence, Dan 143

Boucot, Arthur J. 291

Bowen, David Q. 493, 798

Buckley, Brendan M. 239, 269

Burbank, Douglas W. 596

Catling, David C. 66

Cecil, C. Blaine 181

Chandler, Mark A. 566

Chappell, John M. 893

Choi, Kyungsik 183

Christensen, Phillip R. 540

Clague, John J. 206

Clarke, Jonathan 138

Colgan, Patrick M. 354

Coope, G. Russell 90

Cronin, Thomas M. 663, 757

Crosta, Xavier 31

Crowley, Thomas J. 164

Cullen, Heidi M. 619

Dalrymple, Robert W. 183

de Vernal, Anne 38, 280

DeConto, Robert M. 784

Delaygue, Gilles 278, 666

Derry, Louis A. 981

Dickens, Gerald R. 560

Dilcher, David 679

Dlussky, Konstantin G. 522

Donnelly, Jeffrey P. 763

Dorn, Ronald I. 275

Dowdeswell, Julian A. 471

Dowsett, Harry J. 338

Driese, Steven G. 748

Droser, Mary L. 10

Dugmore, Andrew J. 937

Dyke, Arthur S. 517

Emeis, Kay-Christian 875

Enters, Dirk 485

Fairbridge, Rhodes W. 294, 414, 428, 451, 551

Farquhar, James 61, 333

Fegley, Bruce, Jr. 75

Filippelli, Gabriel M. 780

Forswall, Clayton 560

Fowell, Sarah 766

Francois, Roger 634

Gagan, Michael K. 721

Gaines, Robert R. 10

Garver, John I. 247

Gasse, Françoise 733

Goodbred, Steven L. 265

Goodwin, Ian D. 89

Gornitz, Vivien 6, 294, 428, 573, 716, 813, 887

Goudie, Andrew 45

Gröcke, Darren R. 397

Gupta, Anil K. 589

Hall, Chris M. 255, 823

Harper, David A. T. 325

Hay, William W. 260

Hays, James D. 158

Haywood, Alan M. 804

Hemming, Sidney 409

Hendrix, Marc S. 192

Higginson, Matthew J. 341

Hoek, Wim Z. 100, 477

Hoffmann, Georg 282

Huber, Matthew 683

Huybrechts, Philippe 221

Jacobs, Stan 16

Johnson, Claudia C. 213

Juggins, Stephen 959

Kaplan, Jed O. 507

Karlén, Wibjörn 835

Kauffman, Erle G. 213

Kaufman, Alan J. 85

Kearney, Michael S. 899

Kneller, Margaret 766, 815

Knight, Peter G. 89, 155, 284, 320, 483,

594, 665

Kohfeld, Karen E. 286

Kolodny, Yehoshua 775

Kowalski, Elizabeth A. 679

Krijgsman, Wout 252

Kull, Christoph 675

Lacis, Andrew A. 2, 174

Lambeck, Kurt 374

Langereis, C. G. 252