Response of interior North America to abrupt climate oscillations in the North Atlantic region during the last deglaciation

Ž .Earth-Science Reviews 52 2001 333–369www.elsevier.comrlocaterearscirev

Response of interior North America to abrupt climate oscillationsin the North Atlantic region during the last deglaciation

Zicheng Yu a,), H.E. Wright Jr. b

a Canadian Forest SerÕice, 5320 122 Street, Edmonton, Alberta, Canada T6H 3S5b Limnological Research Center, UniÕersity of Minnesota, Minneapolis, MN 55455, USA

Received 10 May 2000; accepted 29 September 2000

Abstract

Several broad-scale climatic oscillations during the last deglaciation are well documented in regions around the NorthAtlantic Ocean. This paper reviews empirical evidence for these deglacial climatic oscillations from non-coastal NorthAmerica and discusses implications for testing climatic simulations and for understanding the cause and transmittalmechanisms. Paleoclimatic interpretation of oxygen-isotope records from several small sites in the eastern Great Lakesregion indicates a classic deglacial climatic sequence that is comparable with records from Europe and Greenland. The

Ž .climatic events as interpreted from Crawford Lake oxygen isotopes include the Bølling–Allerød BOA warming at;12,700 14C BP, a warm BOA at ;12,500–10,920 14C BP, an intra-Allerød cold period shortly before 11,000 14C BP, a

Ž . 14 14cold Younger Dryas YD climate reversal at 10,920–10,000 C BP, the Holocene warming at 10,000 C BP, a briefŽ . 14 14 Ž .Preboreal Oscillation PB at 9650 C BP, and an early-Holocene cold event at 7500 C BP 8200 cal BP . Some of these

events were also evident from changes in upland and aquatic vegetation and sediment lithology. The pronounced YDclimatic reversal has been documented from pollen records along the ecotones at this time and from glacier readvances in theGreat Lakes region. Along the Rocky Mountains from Alberta to Colorado, the YD event is indicated by alpine glacieradvance andror shift in timberline vegetation. In Minnesota and upper New York, early-Holocene climatic instability is alsosuggested by oxygen-isotope records andror varve thickness. The regional variations in evidence for the YD and otherevents in North America suggest that climatic oscillations may have different expressions in paleo-records, depending ongeographic location and characteristics of a particular site. The extent and magnitude of these climatic oscillations acrossNorth America suggest that these oscillations are an expression of climatic change that was probably widespread rather thanlocally induced by a nearby glacier. The location of these sites implies that climatic signals were likely carried over theNorthern Hemisphere through the atmosphere, as indicated by general circulation models. We hypothesize that the lack ofevidence for a cold YD in interior North America west of the Great Lakes region and east of the Rocky Mountains wascaused by the trapping of cold arctic air mass north of the Laurentide ice sheet and by uninterrupted northward expansion ofwarm Caribbean air; this strongly contrasted climate also produced non-analogous biological assemblages. q 2001 ElsevierScience B.V. All rights reserved.

Keywords: North America; Late glacial; Younger Dryas; deglacial climatic oscillations; multiple proxy paleorecords; air-mass distribution;Amphi-Atlantic climatic contrasts

) Corresponding author. Tel.: q1-780-435-7304; fax: q1-780-435-7359.Ž .E-mail address: [email protected] Z. Yu .

0012-8252r01r$ - see front matter q 2001 Elsevier Science B.V. All rights reserved.Ž .PII: S0012-8252 00 00032-5

( )Z. Yu, H.E. Wright Jr.rEarth-Science ReÕiews 52 2001 333–369334

1. Introduction

In the North Atlantic region, the large and abruptclimatic oscillations that characterized the last glacia-tion continued into the last deglacial period. Theseoscillations include quasi-periodic millennial-scale

Ž . ŽDansgaard–Oeschger D-O events ;1500-year.spacing and related low-frequency Heinrich–BondŽcycles Dansgaard et al., 1993; Bond et al., 1993,

1997; Bond and Lotti, 1995; Grootes and Stuiver,1997; Mayewski et al., 1997; Alley, 1998; Alley and

.Clark, 1999 . During the last glacial–interglacialtransition, the most recent manifestation of thesemillennial- and century-scale climatic events wasbroadly experienced in both Europe and North

ŽAmerica e.g., Lotter et al., 1992; Levesque et al.,1993; Bjorck et al., 1996; Yu and Eicher, 1998; Von¨

.Grafenstein et al., 1999 . These climatic events in-Žclude the Bølling warming at 14,600 cal BP 12,700

14 .C BP and the gradual cooling trend toward theŽ .Younger Dryas YD climate reversal, which can be

regarded as the last Bond cycle. During the generalŽ .cooling trend of the Bølling–Allerød BOA warm

period, several century-scale oscillations include theŽ .Older Dryas OD and the intra-Allerød cold period

ŽIACP; also called Gerzensee or Killarney Oscilla-.tion . In the early Holocene, Preboreal Oscillation

Ž .PB and 8200 cal BP cooling events have beenŽwidely recognized in paleoclimatic records Alley et

.al., 1997 . The YD event is an abrupt and verymarked millennial-scale cooling episode between ;

11,000 and 10,000 14C BP. It is regarded as the latestŽ .Heinrich event H0 and has been more extensively

studied than previous ones.

These deglacial climatic oscillations have beenwell documented around the North Atlantic Seaboard,but until recently few studies have addressed the

Žcontinental interior of North America Rind et al.,.1986; Shane, 1987; Wright, 1989; Peteet, 1995 .

Establishing the geographic extent and relative mag-nitude of these deglacial climatic oscillations is es-sential for understanding the mechanisms and causesof these abrupt climatic changes in particular and ofEarth’s climatic system in general. Here we reviewrecent multiple proxy studies from non-coastal North

Ž .America. The objectives of this paper are: 1 to putevidence for deglacial climatic oscillations from inte-rior North America in the context of other regions;Ž .2 to evaluate the transmittal mechanisms and toexplain the contrasting response of the continental

Ž .interior; and 3 to discuss strategy and future direc-tions for obtaining empirical paleoclimatic records.The paper will first provide a brief overview ofdeglacial climatic events from regions around theNorth Atlantic and from the west coast of NorthAmerica. We then focus on evidence from interiorNorth America, which has lacked examples ofdeglacial climatic oscillations. Lastly, we evaluatethe ultimate and proximate causes and mechanismsof transmittal based on a coupled ocean–atmospheregeneral circulation model and propose a hypothesisthat explains the contrasting late-glacial climates be-tween the Atlantic region and the North Americancontinental interior.

The time period covered in this paper spans thelate-glacial and early Holocene. The late-glacial hererefers to the time period of ;13,000–10,000 14CBP, which is characterized by extreme climatic insta-

Table 1Ž .Ages of deglacial climatic events in the North Atlantic region from Mangerud et al., 1974; Wohlfarth, 1996; Bjorck et al., 1998¨

Ž . Ž .Mangerud et al.’s 1974 Radiocarbon age Bjorck et al.’s 1998 Calendar age for the onsetäŽ . Ž .chronozones years BP event stratigraphy GRIP year BP

Preboreal 10,000–9000 Holocene 11,500–Younger Dryas 11,000–10,000 GS-1 12,650–Allerød 11,800–11,000 GI-1a 12,900–

Ž .GI-1b IACP 13,150–GI-1c 13,900–

Older Dryas 12,000–11,800 GI-1d 14,050–Bølling 13,000–12,000 GI-1e 14,700–Oldest Dryas )13,000 GS-2a 16,900–

aGS — Greenland Stadial; GI — Greenland Interstadial; IACP — intra-Allerød cold period.

( )Z. Yu, H.E. Wright Jr.rEarth-Science ReÕiews 52 2001 333–369 335

bility and includes classic European chronozonesŽ 14 . ŽBølling 13,000–12,000 C BP , Allerød 11,800–

14 . Ž 14 .11,000 C BP and YD 11,000–10,000 C BPŽ .Mangerud et al., 1974; Wohlfarth, 1996 . The termi-

Ž .nology of Mangerud et al. 1974 is used here forŽ .convenience. Recently, Bjorck et al. 1998 proposed¨

a new late-glacial event stratigraphy for the NorthAtlantic region based on the stratotype of the Green-

Ž .land ice-core isotope record GRIP . Their eventstratigraphy extends from the beginning of theHolocene downward, including events Greenland

Ž .Stadial 1 GS-1 for the YD, Greenland Interstadial 1Ž .GI-1 for the Allerød and Bølling, and GS-2 for theOldest Dryas. The GI-1 is further subdivided into

Žepisodes GI-1a warm period between YD and. Ž . ŽIACP , GI-1b IACP , GI-1c warm period between

. Ž . ŽIACP and OD , GI-1d OD and GI-1e Bølling.warm period . Table 1 lists radiocarbon and calendar

ages of deglacial climatic events and correlation ofŽ .the Mangerud et al. 1974 terminology and the

Ž .Bjorck et al. 1998 event stratigraphy. Both calendaränd radiocarbon ages are used in the paper, becausecalibration between the two time scales for thatperiod is not straightforward or reliable.

2. Deglacial climatic oscillations

Many sites in western and central Europe andGreenland show multiple deglacial climatic oscilla-tions. Here we select several sites from regions

Ž .around the North Atlantic Ocean Fig. 1 to illustrateŽ .these climatic events Fig. 2 . Before we discuss the

evidence from the interior of North America, webriefly review well-documented records from coastal

Ž .regions Fig. 3 .Many paleoclimatic records from the North At-

lantic region during the last deglaciation from 15,000to about 7000 cal BP show a classic deglacial cli-matic sequence from the initial BOA warming, toBOA warm period with two or more minor oscilla-tions, to the YD cold period, and to Holocene warm-ing with PB and a cooling event at 8200 cal BPŽ . Ž .8.2-ka event Fig. 2B–E . The last Bond cyclefrom the peak BOA warming at ;14,500 cal BP tothe onset of the Holocene at ;11,500 cal BP showsa progressive cooling through two or three D-Ocycles, followed by an abrupt warming. The coldest

Žpart of this Bond cycle is marked by the YD Hein-.rich layer H0; Andrews et al., 1994 . These events

ŽFig. 1. Map showing location of selected paleoclimatic sites around the North Atlantic Ocean: Crawford Lake in North America Yu and. Ž .Eicher, 1998 , Summit ice-cores from Greenland Dansgaard et al., 1993; Grootes et al., 1993; Stuiver et al., 1995 , marine core Troll3.1

Ž . Ž .from North Atlantic Lehman and Keigwin, 1992 , Gerzensee in central Europe Eicher, 1980 , and marine core PL07-56PC from CariacoŽ .basin in tropical Atlantic Hughen et al., 1996 .


Ž .can be traced from North America Fig. 2A toŽ . ŽGreenland Fig. 2B , the North Atlantic Ocean Fig.

. Ž .2C , central Europe Fig. 2D , and the tropical At-Ž .lantic Ocean Fig. 2E . Atmospheric CH concentra-4

tions show similar fluctuations during the major andŽ .some minor climatic events Fig. 2F .

2.1. Greenland and North Atlantic Ocean

Multiparameter records from Greenland ice coresprovide the clearest evidence of abrupt deglacial

Žclimatic changes in many aspects e.g., Fig. 2B and. Ž .F; Alley, 2000 . One of these ice cores GRIP was

Ž .proposed by Bjorck et al. 1998 as the stratotype for¨Ž .an event stratigraphy Table 1 . The detailed glacio-

chemical data benefit from well preserved annual icelayers that can be counted with confidence. Theoxygen isotopes from the ice reflect mostly changesin air temperatures, showing the major YD climaticreversal between the BOA warm period and the

Žwarm Holocene Fig. 2B; Johnsen et al., 1992;Grootes et al., 1993; Dansgaard et al., 1993; Stuiver

.et al., 1995; Stuiver and Grootes, 2000 . During thepre- and post-YD warm periods, several minor oscil-lations are also clearly documented. Other proxiesshow that the transitions of these abrupt climaticshifts, especially the YD, occur very rapidly within a

Žfew years to a few decades Dansgaard et al., 1989;Alley et al., 1993; Taylor et al., 1993, 1997; Sever-

.inghaus et al., 1998; Severinghaus and Brook, 1999 .A significant change in atmospheric chemical com-

Žposition occurred during the YD Mayewski et al.,.1994 .

Ž .Ruddiman and McIntyre 1981 used foraminiferalstratigraphy of ocean cores to show that the southernlimit of the polar front shifted southward during theYD climatic reversal. Deep-sea records traditionallyhave low temporal resolution due to usually very lowsediment-accumulation rates, but many marine

records with high resolution demonstrate theseŽdeglacial climatic oscillations in more detail Fig. 2C

and E; Lehman and Keigwin, 1992; Hughen et al.,.1996 . Other marine records from the North Atlantic

Žalso show these deglacial climatic events e.g., Koc.Karpuz and Jansen, 1992; Bond et al., 1993 .

2.2. Western and central Europe

The YD cooling had significant impact on Euro-pean terrestrial and aquatic ecosystems, and the mostconvincing continental evidence for the deglacialclimate events comes from western and central Eu-rope. Initially, the YD climatic reversal was recog-nized in glacier reconstructions and in pollen andmacrofossil records showing vegetation containing

Žthe tundra plant Dryas octopetala in Denmark Iver-.sen, 1954 . Over the last century, abundant evidence

is available from multidisciplinary studies, but differ-ent proxies may have responded differently and have

Žrecorded different timings for the cooling Wright,.1984; Pennington, 1986; Birks and Ammann, 2000 .

After the Scandinavian Ice Sheet retreated fromthe Last Glacial Maximum, it paused or readvancedin many places, as clearly documented in southern

Ž .Sweden Bjorck et al., 1988; Berglund et al., 1994 .¨The YD is marked in northwestern Europe by achange to a pronounced periglacial condition, withrenewed soil erosion and mineral inwash into lakesŽ .Walker et al., 1994 . Alpine glaciers also respondedto the YD cooling. In Switzerland, for example, theEgesen moraine was formed during the YD, as hasbeen recently confirmed by 10 Be, 26Al and 36Cl AMS

Ž .dating Ivy-Ochs et al., 1996 . Oxygen-isotoperecords show detailed evidence for the multiple cli-matic oscillations during the last deglaciation inwestern and central Europe. Several sites in lowlandSwitzerland show a climatic sequence of BOAwarming, BOA warm period, YD cold interval, and

Ž . Ž . 18Fig. 2. Correlation of paleorecords during the last deglaciation 15,000–7000 cal BP . A d O of lacustrine carbonates at Crawford LakeŽ . Ž . 18 Ž . Ž . ŽYu et al., 1997; Yu and Eicher, 1998 ; B d O of ice-core GISP2 Grootes et al., 1993 ; C Forams of Troll3.1 Lehman and Keigwin,

. Ž . 18 Ž . Ž .1992; age scale was tuned to Greenland ice-core record ; D d O of lacustrine carbonates at Gerzensee Eicher, 1980 ; E Grey scale ofŽ . Ž . Žcore PL07-56PC Hughen et al., 1996; revised age scale ; F Atmospheric CH concentration of ice cores GRIP circles and thin line;4

. Ž .Chappellaz et al., 1993 and GISP2 thick line; Brook et al., 2000 . Late-glacial ages for Crawford, Gerzensee and Troll3.1 cores weretentatively tuned to GISP2 ice-core time scales on the basis of similar climatic oscillations, so the similarity in timing does not meansynchrony of events at centennial scales.


warm Holocene, with two minor oscillations in BOAŽAegelseesOlder Dryas; Gerzensees intra-Allerød

. Ž . Žcold period and one PB in early Holocene Eicher

and Siegenthaler, 1976; Eicher, 1980, 1987; Siegen-thaler and Eicher, 1986; Lotter et al., 1992; Ammann

.et al., 2000 . This sequence shows remarkable simi-


Fig. 3. Location map of paleoclimatic sites in the interior of NorthŽ .America discussed in the text. Coastal regions shaded areas that

show evidence for deglacial climatic oscillations are not discussedŽ . Žin detail in the text: A Greenland e.g., Dansgaard et al., 1993;. Ž .Grootes et al., 1993 ; B Baffin Island and northern Labrador

Ž . Ž . ŽAndrews, 1994 ; C Eastern coast of North America Mott et al.,1986; Peteet et al., 1990, 1993; Mayle et al., 1993a,b; Cwynar and

. Ž . ŽLevesque, 1995 ; D Southern Alaskan coast Engstrom et al.,. Ž .1990; Peteet and Mann, 1994; Hansen and Engstrom, 1996 ; E

ŽPacific Northwest coast Mathewes et al., 1993; Mathewes, 1993;. Ž . ŽGrigg and Whitlock, 1998 and F California coast e.g., Benson

.et al., 1997 . Dots show location of sites in central North America:Ž . Ž .1–3 Crawford Lake, Twiss Marl Pond Yu and Eicher, 1998 ,

Žand Gage Street site, southern Ontario, respectively Fritz et al.,. Ž .1987; Yu, 2000 ; 4 Lake Seneca of the Finger Lakes region,

Ž . Ž .New York Anderson et al., 1997 ; 5 Nichols Brook site, NewŽ . Ž .York Fritz et al., 1987; Yu, 2000 ; 6 Pyle and Stotzel-Leis sites,Ž . Ž .Ohio Shane, 1987; Shane and Anderson, 1993 ; 7 Prince Lake,

Ž . Ž .northwestern Ontario Saarnisto, 1974 ; 8 Lake Gribben forestŽ . Ž .bed, Michigan Lowell et al., 1999 ; 9 Deep Lake, Minnesota

Ž . Ž .Hu et al., 1997, 1999 ; 10 Rattle Lake, northwestern OntarioŽ . Ž .Bjorck, 1985 ; 11 Sky Pond and Black Mountain Lake, CO¨Ž . Ž .Reasoner and Jodry, 2000 ; 12 Inner Titcomb Lakes moraine of

Ž . Ž .the Wind River Basin, WY Gosse et al., 1995 ; 13 CrowfootŽmoraine, Alberta Reasoner et al., 1994; Reasoner and Huber,

. Ž . Ž .1999 ; 14 Southern High Plains Holliday, 2000 .

larity with ice-core records from GreenlandŽ .Siegenthaler et al., 1984; Lotter et al., 1992 . Oxy-gen-isotope records that show deglacial climatic os-

Žcillations are also available from Poland Goslar et.al., 1993, 1995; Ralska-Jasiewiczowa et al., 1998 ,

Žsouthern Germany von Grafenstein et al., 1992,. Ž .1994, 1998, 1999 , Ireland Ahlberg et al., 1996 ,

Ž .and southern Sweden Hammarlund et al., 1999 .Paleobotanical records from pollen and macrofos-

sil data generally show an open vegetation during theYD cold period. For example, the YD cooling was

Žmarked by tundra taxa in Britain and Ireland Hunt-.ley and Birks, 1983; Walker, 1984; Watts, 1985 or

by the expansion of Artemisia in the grasslandŽ .Watts, 1977; Cwynar and Watts, 1989 . In theNetherlands, closed forests were replaced by eitheropen woodlands or dwarf shrub and heath vegetationŽ .Van Geel et al., 1989; Bohncke, 1993; Hoek, 1997 .

Ž .In lowland Switzerland the Swiss Plateau , the YDevent is recorded by an increase of juniper andpioneer heliophilous taxa in the pine-birch wood-

Žlands Ammann, 1989; Ammann and Lotter, 1989;Lotter et al., 1992; Ammann et al., 1994; Lotter,

.1999 . In western and central Europe, the YD isclearly shown in paleobotanical records, but the mi-

Ž .nor climatic oscillations e.g., PB, GrK, OD are notŽso obvious but see Lotter, 1999; Birks and Am-

.mann, 2000; Leroy et al., 2000 . Fossil insects fromnorthern Europe show clear evidence for deglacialclimatic oscillations, including the BOA warming,declining temperature trend during the BOA warm

Žperiod, and the YD Coope, 1977; Atkinson et al.,.1987; Coope and Lemdahl, 1995; Coope et al., 1998 .

Other paleoclimatic proxy data also show sensitiveresponse to climatic oscillations, for example, chi-

Ž .ronomids Brooks et al., 1997; Mayle et al., 1999 ,Ž .land snails Rousseau et al., 1998 , and lake levels

Ž .Magny and Ruffaldi, 1995 . Isarin and BohnckeŽ .1999 mapped the inferred July temperatures ofnorthern and central Europe during the YD interval

Žon the basis of climatic indicator species e.g., cer-. Žtain aquatic plants see also Isarin and Renssen,

.1999 .

2.3. Eastern coast of North America

Ž .Mott et al. 1986 summarized the geologic andpalynologic results from Maritime Canada thatshowed clear evidence of cooling between about11,000 and 10,000 14C BP. That paper gainedwidespread acceptance as evidence for the YD on


the other side of the Atlantic Ocean. In southernNew Brunswick and central mainland Nova Scotia,the YD is indicated by a succession from

Žwoodlandrforest to tundra Mayle et al., 1993a,b;Cwynar et al., 1994; Levesque et al., 1994; Mayle

.and Cwynar, 1995 , but in central New Brunswickand northern Nova Scotia from shrub tundra to herb

Žtundra Levesque et al., 1993; Mayle et al., 1993a,b;. ŽMott, 1994 . Relatively low organic matter as mea-

.sured by loss on ignition indicates renewed solifluc-tion, increased minerogenic inwash, or reduced lake

Žproductivity due to cooling e.g., Mott et al., 1986;.Mayle et al., 1993a; Mott and Stea, 1993 . Chirono-

mids have been successfully used to reconstructsummer surface-water temperatures for the YD and

ŽGrK events in Maritime Canada and Maine Walkeret al., 1991; Wilson et al., 1993; Levesque et al.,

.1993, 1994; Cwynar and Levesque, 1995 . AndersonŽ .and MacPherson 1994 summarized palynologic data

from Newfoundland and showed an increase in herbtundra taxa and in sediment alkali during the YD,suggesting reduced vegetation cover and acceleratedsoil erosion. They also reported evidence for minorclimatic oscillations shortly before the YD and in the

Ž .early Holocene. Miller 1997 analysed fossil insectsfrom three late-glacial sites in Nova Scotia andfound evidence for the YD, but it is not as strong asthe palynologic and lithologic records.

In southern Quebec, slight changes in pollen as-semblages within herb tundra and decreased pollenconcentration suggest the YD climatic reversalŽ .Marcoux and Richard, 1995 as well as an early-Ho-

14 Žlocene cold event at 9500 C BP Richard et al.,.1992; Richard, 1994 . In southern New England, the

YD is recorded by a decrease in the pollen percent-age of thermophilous temperate deciduous trees suchas Quercus and Fraxinus and increase of borealforest taxa such as Abies, Larix, Picea, Betula, and

ŽAlnus e.g., Peteet, 1987; Peteet et al., 1990, 1993,.1994; Maenza-Gmelch, 1997a,b . In Baffin Island

Ž .and northern Labrador, Andrews 1994 reviewedterrestrial and marine evidence and showed that a

Ž .permanent ice cover sea ice or an ice shelf existedŽ .during the YD interval. Lowe 1994 and Lowe et al.

Ž .1995 summarized evidence for climatic oscillationsduring the last deglaciation around the North At-lantic Seaboard, including the east coast of North

Ž .America Fig. 4 .

Fig. 4. Summary paleotemperature curves during the last deglacia-Žtion for the regions in east coast of North America modified from

.Lowe and NASP Members, 1995 .

2.4. West coast of North America

On Pleasant Island in the Glacier Bay area ofŽ .southeastern Alaska, Engstrom et al. 1990 and

Ž .Hansen and Engstrom 1996 described a pollensequence implying that an established lodgepole pineŽ .Pinus contorta parkland or forest reverted to shrub-

Žand herb-dominated tundra Ericaceae, Poaceae,.Cyperaceae, and Artemisia from 10,800 to 9800

14C BP. This vegetational change is matched bygeochemical evidence of reduction of organic matterfrom catchment soils and increased mineral erosionŽ .Engstrom et al., 1990 . They interpreted these vege-tational and geochemical changes as caused by apossible YD cooling event, because no known read-vance of glaciers at that time in the area couldexplain the changes. On Kodiak Island in southwest-

Ž .ern Alaska, Peteet and Mann 1994 found distinctlithologic oscillations and a dramatic reversal invegetation involving the near disappearance of Poly-

Ž . 14podiaceae AFern GapB from 10,800 to 10,000 CBP, suggesting colder and drier conditions. Theyinterpreted this reversal as a high-latitude Pacificexpression of the YD event.

Just as with the YD event around the NorthAtlantic, evidence for the climatic reversal in thePacific Northwest is best expressed in hypermaritime

Ž .and maritime climatic regions Mathewes, 1993 .The most notable evidence for the YD is the pollen

Ž .peaks of mountain hemlock Tsuga mertensiana , an


indicator of cool and moist climate, in westernŽ .Washington Heusser, 1973, 1977 , along the British

Ž .Columbia coast Mathewes, 1973 , on Queen Char-Ž .lotte Island Mathewes et al., 1993 , and on Vancou-Ž .ver Island Hebda, 1983 . At two sites in western

Ž .Oregon, Grigg and Whitlock 1998 found an in-Ž .crease in pollen of western white pine P. monticola

between 12,400 and 11,000 cal BP, indicating coolerwinters and drier summers than in earlier periods,and they suggested that this Pinus expansion maycorrelate with the YD climatic reversal. They sug-gested that the weak expression of the YD event wasdue to the relative inland location of their study sites.Farther south in the Great Basin of California, Ben-

Ž .son et al. 1997 reconstructed the hydrological bal-ance of Owens Lake from carbonate d

18 O values andfound that millennial and century-scale dry events inwestern North America during the last deglaciationoccurred at about the same time as cold eventsrecorded in Greenland ice, including the YD andpre-YD oscillations.

2.5. Other places in the world

Several recent reviews have been published thatsummarize the state of knowledge about evidence fordeglacial climatic oscillations, especially the major

ŽYD climatic reversal Rind et al., 1986; Wright,.1989; Peteet, 1995; Rutter et al., 2000 . Increasing

evidence is available for South America from paly-Žnologic data Kuhry et al., 1993; Van der Hammen

and Hooghiemstra, 1995; Isbele et al., 1995; Leyden,. Ž1995; Hansen, 1995 and the glacier record Osborn

.et al., 1995 . In New Zealand, Denton and HendyŽ .1994 correlated a glacial advance of the FranzJosef Glacier with the YD cooling. Southern Hemi-sphere ice cores also suggested climatic oscillations

Ž .during the last deglaciation Thompson et al., 1995 .Ž .An et al. 1993 suggested that a strengthened sum-

mer monsoon indicated by loess data from centralChina was correlated with the YD event. Zhou et al.Ž .1996 suggested that significant and rapid oscilla-tion of paleomonsoon proxy for monsoonal easternAsia occurred during the YD interval. Porter and AnŽ .1995 suggested that climatic events in China corre-lated with the North Atlantic region during the lastdeglaciation. In the varved sediments of Lake Van ineastern Turkey, the YD is well represented in the

sediment chemistry and the pollen profiles by indica-Žtions of a cold dry climate Lemcke and Sturm,

.1996 . In east Africa, a lower lake level was corre-Ž .lated with the YD interval Roberts et al., 1993 , asŽin other lakes in the tropics Street-Perrott and Per-

.rott, 1990 .Ž .The Antarctic Cold Reversal ACR from the

Byrd and Vostock ice cores in Antarctica appearedto precede the YD from Greenland ice cores by atleast 1000 years, as indicated by synchronization ofrecords with the isotopic ratio of molecular oxygenŽ .Bender et al., 1994; Sowers and Bender, 1995 andwith the global atmospheric concentrations of

Ž .methane Blunier et al., 1997 . Similar phase rela-tionship also existed earlier during the last deglacia-tion, showing that Greenland warming associatedwith D-O events lags the similar events in Antarctica

Ž .by a couple of thousand years Blunier et al., 1998 .These data suggest that these climate changes origi-nated in the Southern Hemisphere and that theyspread to the north. However, isotopic climate proxyrecord from Taylor Dome ice core in coastal EastAntarctica shows synchronous climatic changes with

Ž .Greenland during the YD Steig et al., 1998 . High-resolution records of isotopic and methane data fromGreenland ice cores indicate that the warmings in theNorth Atlantic during the last deglaciation occurredseveral decades before the tropical warmingsŽSeveringhaus et al., 1998; Severinghaus and Brook,

.1999 . Understanding how the Northern and South-ern Hemispheres are coupled during climatic events

Žis a central issue in climate dynamics Blunier et al.,.1998; White and Steig, 1998 . It appears that many

questions remain to be answered.

3. Evidence from the Great Lakes region

3.1. Stable isotopes

Clear evidence for deglacial climatic oscillationsin the Great Lakes region have been recently docu-mented by stable-isotope records at two small lakesalong the Niagara Escarpment in southern OntarioŽ .sites 1 and 2 in Fig. 3; Yu and Eicher, 1998 . Theoxygen-isotope records from three cores at Crawford

Ž .Lake Figs. 5 and 6 show the classic Europeanclimatic sequence. A striking feature from all three


Fig. 5. Photographs and correlation of sediment sections of cores DC, SC and BC from Crawford Lake, southern Ontario. PB, PreborealŽ .Oscillation; GrK, Gerzensee or Killarney Oscillation modified from Yu, 1997 .

cores is the d18 O minimum in the upper Picea-

dominated pollen zone, which indicates a decline intemperature. These negative shifts of d

18 O of 0.5‰,

1.5‰, and 1.2‰ for cores DC, SC, and BC, respec-tively, correspond with the beginning of the YD cold

14 Ž .event at ;11,000 C BP Yu and Eicher, 1998 .


18 Ž .Fig. 6. Event correlation of d O profiles from carbonates at Crawford Lake core SC: heavy curve; core BC: dashed curve in the GreatŽ . 14 ŽLakes region and from Greenland ice-core GISP2 left bar-curve on C and calendar-year time scales, respectively Grootes et al., 1993;

.Yu and Eicher, 1998 . At Crawford Lake, the sequence of climatic oscillations includes the BOA warming, GrK, YD, PB and HE-5w x Ž .Holocene Event 5; Bond et al., 1997; i.e., the 8200 cal BP cooling event , and possibly OD Older Dryas . All these oscillations recorded atCrawford Lake have about a third to half the amplitude of the Greenland record, and PB and HE-5 have half the amplitude of the YD and

Ž .GrKrIACP at both locations modified from Yu and Eicher, 1998 .

After this minimum, the d18 O values show an abrupt

positive shift of 1.2‰, 1.5‰, and 3‰ for cores DC,SC, and BC, respectively, indicating the onset ofHolocene warming linked with the Picea–Pinustransition at 10,000 14C BP. In the early Holocene, aminor negative d

18 O excursion of ;0.5‰ at 965014C BP occurs in the three cores, corresponding tothe PB event. The 8200 cal BP cooling event is

Ž .evident from two shallow cores SC and BC . Duringthe BOA warm period, a negative excursion of d

18 Oshortly before the YD at core SC indicates a coldinterval correlative with the intra-Allerød cold periodŽ . Ž .IACP in Greenland Lehman and Keigwin, 1992 ,

Ž . ŽGerzensee Oscillation G in central Europe Eicher,. Ž .1980 , and Killarney Oscillation K in Atlantic

Ž .Canada Levesque et al., 1993 . This climatic se-quence is remarkably similar to records from around

Žthe North Atlantic Eicher 1980; Lotter et al., 1992;Johnsen et al., 1992; Levesque et al., 1993; Dans-

gaard et al., 1993; Stuiver et al., 1995; Bjorck et al.,¨.1996; Alley et al., 1997 .

Ž 18 .At Twiss Marl Pond, oxygen-isotope d O re-sults from mollusc shells, Chara-encrustations, andbulk carbonates also show a classic climatic se-quence of a warm BOA at ;12,500–10,920 14CBP, a cold YD at 10,920–10,000 14C BP, theHolocene warming at 10,000 14C BP, a brief PB at9650 14C BP, and a possibly IACP event shortly

14 Žbefore 11,000 C BP Yu and Eicher, 1998; Yu,. 182000 . The d O records show a negative shift of

14 Ž .1.3‰ at 10,920 C BP at 490 cm and a positive14 Žshift of up to 2‰ around ;10,000 C BP at 390

. Ž .cm Fig. 7A . The more negative intervening inter-val indicates a cold period corresponding with theYD event. During the YD interval, d

18 O fluctuatedand had minimum values at first and slight increasein the second half, similar to the bipartition indicated

Žin Europe by sensitive plant fossils Isarin and


Ž 18 . Ž . Ž . Ž .Fig. 7. Correlation of oxygen isotopes d O from sites in the eastern Great Lakes region. A Twiss Marl Pond mollusc shells ; BŽ . Ž . Ž . Ž .Crawford Lake core SC Yu and Eicher, 1998 ; C Nichols Brook site; and D Gage Street site Fritz et al., 1987; Yu, 2000 . See Fritz et

Ž . Ž .al. 1987 for a detailed description of their chronology and pollen zonation and Yu 2000 for reinterpretation of chronology. All profilesŽ . Ž .are based on depth scales shown only for Twiss Marl Pond . PB, Preboreal Oscillation e.g., Lotter et al., 1992; Bjorck et al., 1996 ; GrK,¨

Ž . Ž .Gerzensee or Killarney Oscillations Eicher, 1980; Lotter et al., 1992; Levesque et al., 1993 ; OD, Older Dryas Dansgaard et al., 1993 .V-PDB, Vienna Peedee belemnite.

.Bohncke, 1999 . After the major positive shift ind

18 O at the beginning of the Holocene, a minor14 Žnegative excursion of ;0.4‰ at 9600 C BP 370–

.380 cm correlates with the PB event. This minoroscillation was also indicated by recurrence of Picea

Ž .pollen at Twiss Marl Pond Yu, 2000 and moreŽclearly at nearby Crawford Lake Fig. 8; Yu and

.Eicher, 1998 . Before the YD interval, another slightnegative excursion of d

18 O at 500–510 cm wasrecorded in mollusc shells and bulk marl, possiblycorrelated with the IACPrGrK events, for it was

Žalso indicated by peaks of Al and K Yu and Eicher,.1998 .

In the eastern Great Lakes region, several otherstudies have examined climatic and vegetationalchanges during the last glacial–interglacial transi-tion. Multiproxy data from several small lakes, how-ever, failed to show evidence for late-glacial climaticoscillations and even climatic warming at the onset

Ž . Ž .of the Holocene Fritz et al., 1987 . Yu 2000provided an alternative interpretation of stable-iso-tope records from the Gage Street and Nichols Brook

Ž . Ž .sites Fig. 7C and D of Fritz et al. 1987 . At theŽ . 18Gage Street site site 3 in Fig. 3 , the d O profile

shows fluctuations in the late-glacial and earlyHolocene. A negative excursion of ;1.2‰ occursin the upper Picea zone, especially in marl samplesŽ .Fig. 7D but also in ValÕata tricarinata shell sam-

Ž .ples Fritz et al., 1987 . The modified chronology,based on the major Picea–Pinus transition and the

Ž .herb–Picea transition Yu, 2000 , shifts this negative18 Ž .excursion of d O at 400–360 cm close to the YD

interval. The depletion trend of d13C during that

interval is also similar to the Twiss and Crawford13 Ž .d C records Yu and Eicher, 1998; Yu, 1997, 2000 .

In addition, a slight negative excursion in d18 O

shortly after the YD corresponds to the recurrence orpersistence of Picea pollen at the Gage Street siteŽ .Fritz et al., 1987 , which may correlate with the PBat 9650 BP as demonstrated at Twiss Marl Pond andCrawford Lake.

At the Nichols Brook site in northwestern NewŽ .York site 5 in Fig. 3 , on the basis of lithology,

pollen, fossil insects, and detailed stable-isotope


Ž . Ž .Fig. 8. Composite summary diagram for Crawford Lake modified from Yu, 1997 . Chronozones follow Mangerud et al. 1974 . PB,Preboreal Oscillation; YD, Younger Dryas; GrK, Gerzensee or Killarney Oscillations.

records from mollusc shells, Chara-encrustations,Ž .and bulk carbonates, Fritz et al. 1987 concluded

that there was no evidence for late-glacial climaticoscillations. However, oxygen isotopes do show sev-eral shifts. The d

18 O values from ValÕata tricarinatashells show a negative shift of ;1‰ from y9.3‰Ž . Ž .at 310–270 cm to y10.3‰ at 240–170 cm , and

Ža positive shift of ;1‰ at 170 cm Fig. 7C; Fritz et.al., 1987 . The revised chronology, based on rejec-

tion of 14C dates from the flowing-water interval assuggested by fossil insects, would place the negatived

18 O shift at 250 cm at about 11,600 BP and thepositive shift at 170 cm at 10,200 BP. With therevised chronology, the negative excursion of ;1‰in d

18 O at 250–170 cm appears to correspond withthe YD event. The PB event was also suggested at

Žthis site by a recurrence of Picea pollen Fritz et al.,. 181987 and a slight negative excursion of d O from

Ž .mollusc shells at 150 cm 9600 BP? .Two coring sites from Lake Erie documented

18 Žmajor shifts in d O Fritz et al., 1975; Lewis and.Anderson, 1992 , which were initially interpreted as

Ž .representing temperature changes Fritz et al., 1975and later as caused by changes in meltwater dynam-

Žics of the Great Lakes Fritz, 1983; Fritz et al., 1987;Lewis and Anderson, 1992; also see Rea et al.,

.1994a,b . At coring site 13194, two large increasesŽ . 14 18;5‰ at 12,700 C BP and 10,000 BP in d O

Ž .from mollusc and ostracode shells Fritz et al., 1975were caused by a reduction in glacial meltwater in

Žthe Great Lakes drainage system Fritz, 1983; Fritz.et al., 1987 . During deglaciation, the Great Lakes

Žreceived varying quantities of meltwater Rea et al.,.1994a,b . Consequently, changes in water budget

may have complicated climatic signals in these iso-topic records. Such meltwater complications mayalso apply to the isotopic record at the Long Point

Žsite in the Lake Erie basin Lewis and Anderson,. Ž .1992 , which also shows a large increase ;8‰ in

d18 O at ;10,000 BP, and for sites from Lake Huron

Ž .and Georgian Bay Rea et al., 1994a, b . In fact,these authors interpret their isotopic records as causedby meltwater dynamics and lake-level history, ratherthan climatic change directly.

At Deep Lake, southeast of proglacial Lake Agas-Ž .siz in northwestern Minnesota site 9 in Fig. 3 , Hu

Ž .et al. 1997, 1999 found two negative excursions inŽoxygen isotopes from bulk carbonates having 20–

.50% carbonates during the early Holocene. A 3‰decrease in d

18 O from 11,200 to 10,200 cal BP wasinterpreted as a result of decreased summer tempera-ture caused by the nearby presence of Lake Agassiz


Ž .Hu et al., 1997 , but the lack of pollen and vegeta-tion response to this lake-induced cooling indicatesthat it was insufficient to modify the surroundingvegetation. A second decrease of ;2‰ in d

18 Ooccurred from 8900 to 8300 cal BP, suggesting a

Ž .climatic cooling Hu et al., 1999 .Ž .Anderson et al. 1997 suggested that a 1‰ de-

18 14 Žcrease in d O at 10,100–8200 C BP ;11,700–.9200 cal BP at Seneca Lake in the Fingers Lakes

Ž .region of New York site 4 in Fig. 3 was caused bya cooling event following the YD, as a result ofinflux of meltwater into the Great Lakes that reducedsummer temperatures. They correlated this excursionwith the PB at 9600 14C BP in Greenland ice coresŽ .Johnsen et al., 1992 . However, the PB event only

Ž .lasted for about 200 years e.g., Bjorck et al., 1996 ,¨whereas Anderson et al.’s isotope excursion lastedfor about 2000 radiocarbon years. This significantdifference in duration for a presumed same climaticevent suggests the importance of local factors incontrolling stable-isotope composition of lake car-bonates.

3.2. Vegetation eÕidence from paleoecologicalrecords

ŽAt Rattle Lake in northwestern Ontario site 10 in. Ž .Fig. 3; Fig. 9 , Bjorck 1985 found a pollen se-¨

quence from a PicearFraxinusrUlmusrArtemisiaŽ .warm to a PicearSalixrCyperaceae zone at

14 Ž .11,1000–10,200 C BP cold to a PicearPinusrŽ .JuniperusrBetula zone warm . The near disappear-

ance of Fraxinus and Ulmus at 11,1000 –10,200 14CBP suggests a cold climate, which is also supportedby very low pollen-accumulation rates characterized

Ž .by a dwarf-shrub tundra. Wright 1989 suggestedthat this climatic reversal is correlative with the YDevent.

ŽIn the southern Great Lakes region western Ohio,.Indiana, and Illinois , pollen diagrams reveal a dis-

tinct reversal in the pollen sequence for an intervalŽcorrelative with the YD Shane, 1980, 1987; Wright,

.1989; Shane and Anderson, 1993 . Most clearly atŽthe Pyle and Stotzel-Leis sites on the Till Plains Fig.

.10 , after a long dominant period Picea pollen de-

Ž .Fig. 9. Summary pollen diagram from Rattle Lake, northwestern Ontario that shows vegetation response decline of Fraxinus and Ulmus toŽ .the Younger Dryas cooling Bjorck, 1985; Wright, 1989 .¨


Fig. 10. Summary pollen diagrams from Pyle and Stotzel-Leis sites in the southern Great Lakes region that show clear vegetation responseŽ . Ž Ž . .Picea recurrence to the Younger Dryas cooling reprinted from Shane 1987 by permission of Taylor & Francis .

creased to about 5% at 13,000–11,000 14C BP aspollen of Fraxinus, Quercus, OstryarCarpinus, andother temperate hardwood trees dominated the pollenassemblage. Picea then abruptly increased to 30%,

Ž .and other boreal conifers Abies and Larix alsoincreased. These conifer pollen percentages thenabruptly decreased with the arrival of Pinus. ThePicea recurrence after an interval with relativelyabundant Fraxinus and OstryarCarpinus pollen

14 Žlasted from about 11,000 to 10,000 C BP Shane,.1987; Shane and Anderson, 1993 .

Ž .Anderson and Lewis 1992 summarized pollenrecords around the Great Lakes and along the St.

Lawrence River and identified five types of pollenanomalies from the normal pollen sequence during

Ž .the early Holocene Fig. 11 . They interpreted thatthese pollen anomalies were results of localized coldclimate caused by enhanced meltwater discharge fromproglacial lakes at 10,000–8000 14C BP. They re-ported that the strongest vegetation response to melt-water-induced cooling occurred at sites within oralong margins of large water bodies.

At Crawford Lake and Twiss Marl Pond, the YDclimatic reversal was documented by a negative ex-cursion in d

18 O and a change in lithology and ele-Ž .mental geochemistry Figs. 6–8 . However, this event


Ž . ŽFig. 11. Correlation of cold-tolerant pollen assemblages stippled bands along the Great Lakes–St. Lawrence regions from Anderson and. Ž 14 .Lewis, 1992 . Time scale is in radiocarbon year BP kas1000 year C BP .

Ž .occurred in the upper part of the spruce Piceazone. There was no significant forest response to theYD cooling. This was probably a result of the insen-sitivity of non-ecotonal vegetation. In the easternGreat Lakes region, the Picea dominance started as

14 Ž .early as over 12,000 C BP Jacobson et al., 1987 .In the following ;2000 years, the Picea belt mayhave moved latitudinally in response to climaticchanges. However, as Crawford Lake and TwissMarl Pond were situated in the middle of the broadspruce belt at 11,000 14C BP, there was no detectablechange in forest composition recorded by pollen. Aforest response has been detected in the northern andsouthern boundaries of the spruce belt at the YD

Ž .time, e.g., in Ohio and Indiana Shane, 1987 andŽ .northwestern Ontario Bjorck, 1985 . This implies¨

either that elsewhere pollen techniques are not sensi-tive enough to detect subtle forest change or that theforest persisted below its threshold in non-ecotonal

Ž .locations Shane, 1987 . Nevertheless, the vegetationchange was shown by the slight increase or persis-tence of herbs and decrease of pollen concentrations,although it was not as significant as isotopic andlithologic records. The brief climatic oscillation at9650 14C BP caused the recurrence of Picea pollen,because at that time the Picea–Pinus ecotone wasnear the sites.

Ž .Mayle et al. 1993b summarized the pollen evi-dence for the YD event in eastern North America

Žand proposed the peak or persistence of alder Al-.nus pollen as an indicator. However, this is clearly

not the case for sites in the Great Lakes region.

3.3. Lake leÕels, meltwater and climate

Lake levels in the Great Lakes have been recon-structed and continuously refined over the lastdecade, and their connection with meltwater dis-charge and with climatic oscillations has been dis-cussed on the basis of pollen and stable-isotope dataŽFig. 12; Lewis and Anderson, 1989, 1992; Ander-son and Lewis, 1992; Rea et al., 1994a,b; Lewis et

. Ž .al., 1994 . Rea et al. 1994b suggested that the YDcold event was coeval with the Lake Algonquinhighstand at 11,200–10,200 14C BP and was causedby the lake effect as proposed by Lewis and Ander-

Ž .son 1992 . On the basis of isotopic and lake-leveldata from Lake Huron and Georgian Bay, Rea et al.Ž .1994a,b found that the highstands, including LakeAlgonquin, were characterized by isotopically heavywater. In contrast, the intervening lowstands werecharacterised by cold, dilute, isotopically very lightwater, as indicated by their isotope and ostracode

Ž . Ž .data Rea et al., 1994b . Lewis et al. 1994 com-


Ž . Ž .Fig. 12. Huron Basin during the last deglaciation. Lake levels and oxygen isotopes from Moore et al. 1994 and Rea et al. 1994a . IsotopeŽ . Ž . Ž .data from cores in Lakes Huron 37P, M17 , Michigan 9V and Georgian Bay 3P . mS, main Stanley lowstand, eM, early Mattawa

highstand.

bined seismic, lithostratigraphic, and biotic data toreconstruct the lake-level history of the northernHuron Basin. They suggest a steady drop in lakelevels from the main Algonquin highstand from

14 Ž .10,800 to 10,000 C BP. Lewis et al.’s 1994Žlake-level history Fig. 10 in their paper; see also

.Fig. 2 in Moore et al., 2000 is slightly different inŽ .timing from Rea et al.’s 1994a; Fig. 12 .

3.4. Moraines and oscillations of the southern mar-gin of the ice sheet

During the last deglaciation, there were repeatedŽ .major oscillations advance and retreat of the south-

ern margin of the Laurentide Ice Sheet in the GreatLakes region, which left complex moraines. Ifmoraines occurred over a large geographic area, theymight not be caused simply by ice-sheet dynamicsbut by local or regional climatic fluctuations. Saar-

Ž .nisto 1974 recognized a series of end moraines inŽ .the Lake Superior region of Ontario site 7 in Fig. 3

and Michigan that were formed between 11,000 and10,100 14C BP. These moraines indicate slowdownor a pause of deglaciation during that time, precededand followed by rapid ice retreat. This episode wascalled Algonquin Stadial, a reversal in the late-gla-cial climatic warming, correlated with the YD eventŽ .Saarnisto, 1974 . The lack of pollen evidence forthe climatic reversal in this region was attributed to

non-ecotonal vegetation, because spruce forestrwoodland had already occupied most of the availablelandscape at that time, and to insensitivity of pollentechniques in resolving climatic changes of this mag-

Ž .nitude Saarnisto, 1974 .Ž . ŽLowell et al. 1999 confirmed the age 10,025

14 .C BP of the Lake Gribben forest bed south ofŽ .Lake Superior in Upper Michigan site 8 in Fig. 3

by dating nine wood samples, and they argued thatŽ .the ice advance Marquette readvance that buried

the forest by a prograding ice-contact fan was theresult of the YD cooling event. They identified acontemporaneous glacial margin that extended al-most 1000 km from Duluth, MN, to North Bay,Ontario, and argued that a surging glacier systemwould not produce a linear moraine system across

Ž .both a major basin Lake Superior and a majorŽ . Ž .upland Abitibi Upland . Lowell et al. 1999 also

reviewed the results of the last several decades ofresearch on the glacial geology of the Great Lakesand their possible connection with the broad-scaleYD cooling event.

3.5. Fossil beetles

Fossil insects at several sites in the Great Lakesregion indicate a stable climatic condition without amajor amelioration at the onset of the HoloceneŽ 14from 11,000 to 9000 C BP; Miller and Morgan,


.1982; Schwert et al., 1985; Morgan, 1987, 1992 . Ina recent synthesis and analysis of mutual climaticranges of 40 fossil beetle assemblages from 24 sites

Ž .in northeastern North America, Elias et al. 1996concluded that Aour results show fewer changes dur-ing the Late-glacial, with almost a plateau of summer

Ž .temperatures from 13 to 10 kaB p. 420 . An ex-panded analysis of insect fossils for eastern andcentral North America shows no indication of the

Ž .YD cooling Elias, 1997 .In Europe, fossil insects are regarded as one of

the most sensitive proxy indicators of late-glacialclimatic oscillations, showing clear evidence for the

Ž .YD cooling event in Britain Coope, 1977 . In Mar-Ž .itime Canada, Miller 1997 analyzed fossil insects

from three late-glacial sites in Nova Scotia andfound that evidence for the YD is not as strong asthat presented by palynologic and lithologic studiesŽ .Mott et al., 1986; Mayle et al., 1993a; Mott, 1994 .He interpreted this weak response of beetles to theYD cooling as caused by delayed response to cli-matic change, by response to climate indirectlythrough changes in vegetation and ground cover, andby probably the wide temperature and habitat toler-ances of most beetle species. He suggested a gradientof decreasing beetle sensitivity to YD climate cool-ing from Europe through Maritime Canada to theGreat Lakes region.

4. Evidence from the Rocky Mountains of NorthAmerica

4.1. AdÕance of alpine glaciers

During the last deglaciation after the full-glacialadvance of the Cordilleran Ice Sheet, alpine glaciers

Žreadvanced and left numerous moraines Luckmanand Osborn, 1979; Davis, 1988; Osborn et al., 1995;

.Osborn and Gerloff, 1997; Clark and Gillespie, 1997 .The dating and correlation exercises in recent yearshave accumulated evidence either for or against theYD climatic reversal in the Rocky Mountains. Rea-

Ž .soner et al. 1994 first convincingly established theage of the Crowfoot Advance in the Banff NationalPark, Alberta, between 11,330 and 10,070 14C BP byAMS dating of terrestrial macrofossils from down-stream glaciolacustrine records. They cored two lakes

adjacent to the Crowfoot moraine type locality anddated the sediments bracketing an inorganic sedi-ment interval. These inorganic sediments were asso-ciated with the Crowfoot Advance on the basis ofbulk geochemistry and clast lithology. The CrowfootAdvance is the best glacier evidence for the YD

Žclimatic reversal in the Canadian Rockies Osborn etal., 1995; Reasoner and Huber, 1999; Rutter et al.,

.2000 .Ž .Osborn and Gerloff 1997 reviewed the glacier

evidence from the northern Rockies, presented newdata on moraines in northwestern Montana thoughtto be Crowfoot equivalent, and suggested a YD age

Ž .for these deposits. Gosse et al. 1995 dated the InnerTitcomb Lakes moraine in the Wind River Moun-tains, WY, by AMS 10 Be measurements of bouldersurfaces on the moraine from 13,000 to 11,400 10 Be

Ž .BP equivalent to calendar years overlapping theYD event. Farther south in the Colorado Front Range,

Ž .Menounos and Reasoner 1997 AMS-dated an inter-val of clastic lake sediments that show characteristicsconsistent with glacier activity and suggested that

Ž .advances of alpine glaciers Satanta Peak Advancesoccurred between 13,200 and 11,100 cal BPŽ 14 .11,070–9970 C BP . The equilibrium-line-altitudeŽ .ELA depression associated with these glacial de-posits indicated that the glacier response to the YDcooling in the Rocky Mountains was minor, similarin extent to that of the Little Ice Age advance. Theinterpretation of some glaciolacustrine records iscontroversial. For the Crowfoot Advance, LeonardŽ .1998 argued that the correlation between lake-sedi-ment cores and moraines as suggested by Reasoner

Ž .et al. 1994 may not be secure and suggested thatthe type Crowfoot moraine predates the YD interval.

Similar controversy exists for glacier advancealong the west coast of North America. Clark and

Ž .Gillespie 1997 found that the Recess Peak morainesin the Sierra Nevada, CA, were deposited before11,190 14C BP, predating the YD. In another case at

Ž .Mount Rainier, WA, Heine 1998 found that alpineglaciers retreated during the YD interval, probablydue to a lack of available moisture, and suggestedthat the climate might also have been cold. Theseauthors suggested either regional variable responsesof local climate in North America during the YDinterval or differential responses of alpine glaciers tothe same climatic event.



4.2. Response of timberline Õegetation

Ž .Reasoner and Jodry 2000 recently presentedŽpollen records from two high-altitude lakes Black

.Mountain Lake and Sky Pond in the Colorado RockyMountains showing clear and rapid response of alpine

Žtimberline vegetation to the YD climate cooling Fig..13 . Their records show declines in arboreal pollen

Ž .percentages Picea, Abies, Pinus, and Quercus andpollen-accumulation rates during the YD interval,reflecting a downslope displacement of the alpinetimberline ecotone of 60–100 m in elevation. Thesevegetation responses further strengthen the case ofthe YD in the Rocky Mountains as documented by

Ž .glacier and glaciolacustrine records. Reasoner 1998compared the difference in vegetation responses tothe YD cooling between the Canadian and AmericanRockies. He suggested that the lack of pollen evi-dence for a BOA warm interval in the CanadianRockies was due to the short time gap betweendeglaciation of the trunk valley and the YD advanceof alpine glaciers, which left little time for theexpansion of arboreal taxa prior to the YD interval.

5. Southern interior North America

In the unglaciated regions of central North Amer-ica south of the Great Lakes and east of the RockyMountains, reconstructing late-glacial paleoenviron-ments has proven difficult because long-term, high-resolution records are sparse. Consequently, paleocli-matic information for these regions has to be inferredfrom other deposits, such as loess, sand dunes, and

Ž .soils. Recently, Holliday 2000 reviewed and re-trieved paleoclimatic information from archaeologi-cal, geochemical, and geological data from the

ŽSouthern High Plains of the Great Plains site 13 in.Fig. 3 . He found climatic oscillations during the

Pleistocene–Holocene transition, especially dry andŽwarm conditions during the Folsom period 10,900–

14 . Ž10,200 C BP as indicated by eolian activity wind.erosion, eolian deposition , which coincided with the

YD interval. The Folsom period was drier and possi-bly warmer than the preceding Clovis periodŽ 14 .11,200–10,900 C BP and the period after.

6. Causes of climatic oscillations in the interior ofNorth America

6.1. Local Õersus broad-scale climatic change

For the Great Lakes region, it has been postulatedthat the YD-age climatic reversal and perhaps otherdeglacial climatic events were a locally induced cli-matic change caused by an air mass passing over

Ž .cold proglacial lakes. Lewis and Anderson 1992attribute a gradual decrease of 3‰ in d

18 O from;11,000 to 10,500 14C BP in the Lake Erie basin tothe isotopically light meltwater inflow from the high-stand of proglacial Lake Algonquin in the LakeHuron basin. Using as analogue the cooling effect of

Ž .lowlands west of modern Hudson Bay Rouse, 1991 ,they propose that the lake cooling was responsiblefor the Picea recurrence in the southern Great Lakes

Ž .region Shane, 1987 and for pollen anomalies atŽ .other sites. Shane and Anderson 1993 pointed out,

however, that the strongest YD signal is found atsites distant from the lakes. On the basis of thestable-isotope and pollen record from a small lakenear Glacial Lake Agassiz during its second expan-

Ž .sion, Hu et al. 1997 showed that lake-inducedcooling in the early Holocene did not cause a signifi-cant change in vegetation. In the Great Lakes region,the available data suggest that the detectability ofpollen anomalies during the YD interval relies moreon site location relative to ecotonal vegetation, ratherthan on the distance from proglacial lakes or down-

Žwindrupwind directions Saarnisto, 1974; Shane,.1987; Wright, 1989; Yu, 2000 .

On the basis of isotopic and lake-level data fromthe sediments of Lake Huron and Georgian Bay, Rea

Ž .et al. 1994a,b found that the highstands, includingLake Algonquin at 11,200–10,200 14C BP, werecharacterised by isotopically heavy water, indicating

Fig. 13. Summary pollen percentage and pollen-accumulation rate diagrams of subalpine forest taxa at Black Mountain Lake and Sky PondŽ .in Colorado for the late-glacial and early Holocene period from Reasoner and Jodry, 2000 . These diagrams show clear and rapid response

of alpine timberline vegetation to the Younger Dryas cooling at about 12,900–11,500 cal BP.


a much smaller meltwater component relative toinput from local precipitation and run-off. In con-trast, the intervening lowstands were characterisedby cold, dilute, isotopically very light water. Thesenew results do not appear to support the Lewis and

Ž .Anderson 1989, 1992 hypothesis, although Rea etŽ .al. 1994b maintain that lake-effect cooling is prob-

ably related more to surface area than the meltwatervolume. Nevertheless, the relatively warm lake waterduring highstands of the Great Lakes certainly wouldhave had a less pronounced cooling effect. In fact,the lake areas of Lakes Ontario, Erie, and Michiganat 10,800 14C BP were smaller during the main LakeAlgonquin phase than at present, and Lake Huron

Žwas similar in size see Fig. 2b in Lewis and Ander-.son, 1992 .

The cooling effect of large lakes on adjacent landdoes occur. However, the main difficulty with theLewis and Anderson hypothesis is how this coolingeffect links to the YD climatic reversal, which re-quires more enhanced lake and ice-sheet coolingduring the YD interval than before and after. Theextensive ice sheet itself might very well have over-printed the localized cooling caused by proglaciallake water by modifying atmospheric circulation and

Ž .temperature Clark et al., 1999 . Because of its size,Hudson Bay presently generates a distinct air massoff its coast. During the YD interval, proglacial lakeswould have had a relatively small cooling effectcompared with the extensive ice sheet to the north.

Ž .Anderson and Lewis 1992 also postulated that asimilar meltwater-induced cooling could account forearly-Holocene pollen anomalies at 10,000–8000 14C

Ž .BP summarized in their paper Fig. 11 . However,their hypothesis does not predict the PB coolingfound at sites such as Crawford Lake and TwissMarl Pond and perhaps the Gage Street and NicholsBrooks sites, which were distant from any proglacial

Žlakes at 9500 BP see Fig. 1 in Anderson and Lewis,.1992 . In fact, the areas of the Great Lakes at 9500

BP, except Lake Superior, were much smaller thantoday, and water levels in the Huron and Michigan

Žbasins were near their lowest levels ever see Fig. 20.in Teller, 1987 .

In summary, the hypothesis of local meltwater-in-duced cooling cannot be applied to the YD-age andother climatic oscillations recorded in the Great Lakes

Ž .region. At some of these small lakes Fig. 7 , the

sequence and relative magnitude of climatic changesmatch in detail the records from the North Atlanticregion and indicate that these oscillations are likelyan expression of broad-scale, probably global, cli-matic changes. This similarity cannot be satisfacto-rily explained by any local or regional processes, andit must relate to a more fundamental climatic triggerin the coupled ocean–atmosphere–ice sheet climaticsystem, which would explain these widespread,

Žabrupt climatic events Broecker et al., 1985; Rind etal., 1986; Shane, 1987; Wright, 1989; Peteet et al.,

.1997; Yu and Eicher, 1998 .

6.2. Transmittal mechanism and response of conti-nental interior

Recent data indicate that the abrupt warming atŽ .the end of the last glaciation Bølling warming and

Ž .at the end of the YD Holocene warming in theNorth Atlantic both occurred several decades beforethe tropical warming indicated by an increase inatmospheric methane concentration, suggesting thatthe trigger of deglacial climatic oscillations was re-lated to North Atlantic Ocean rather than to changes

Žin the tropics Severinghaus et al., 1998; Severing-.haus and Brook, 1999 . Many theories to explain

these deglacial climatic oscillations invoke mecha-nisms associated with ice sheets and the North At-lantic Ocean. The mechanisms have long been re-lated to dynamics of northern hemispheric ice sheetsŽice-sheet forcing; McCabe and Clark, 1998; Clark et

. Žal., 1999 and of oceanic circulation Broecker et al.,.1990 or to external forcing from beyond the North

ŽAtlantic region Bond and Lotti, 1995; Goslar et al.,.2000 .

Ž .Alley and Clark 1999 recently synthesized evi-dence and theories on the deglaciation of the North-ern Hemisphere and on climatic oscillations duringthe last deglaciation. They emphasized the impor-tance of temporal perspectives in discussing mil-lennial-scale climatic events. The deglaciation wasultimately a response to increased insolation in Mi-

Žlankovitch cycles. The D-O oscillations ;1500.years spacing were usually superimposed on a longer

H–B Cycle with a spacing of 6000–17,000 years.The last H–B cycle occurred during the lastdeglaciation at the transition from the Bølling warm-

Ž .ing to the YD sH0; Andrews et al., 1994 . Alley


Ž .and Clark 1999 suggest that millennial-scale warm-ing and cooling events during the last deglaciationmay represent some combination of response to freeoscillation or oscillations in the climate system, withforced oscillations linked to changes in the mid-lati-tude ice sheets. A meltwater-triggered shift in ther-mohaline circulation in the North Atlantic during theYD would have had the potential to greatly cool thehigh-latitude North Atlantic, Greenland, and Europe.The YD climatic reversal was possibly, at least inpart, related to meltwater diversion in North AmericaŽ .Broecker et al., 1988 . However, it was one of manysimilar abrupt climatic changes during the lastglaciation and last deglaciation that may have had

Ž .multiple causes Alley and Clark, 1999; Alley, 2000 .Holocene D-O-like events usually have a variable

Ž .spacing of 700–2200 years Bond et al., 1997 ; the8.2-ka cooling event was one of these and has a

Žgeographic pattern similar to that of the YD Alley et.al., 1997 .

A widely accepted hypothesis is that iceberg andmeltwater pulses from wasting ice sheets diluted thesurface ocean with less dense water, thus shifting oreven shutting down the thermohaline circulationŽ . Ž .THC and North Atlantic Deep Water NADWformation, causing reduced heat transfer from sub-tropical to subpolar regions and cooling the highlatitudes around the North Atlantic, especially down-

Žwind, i.e., Europe Broecker et al., 1988, 1989;.Broecker, 1994, 1997 . This mechanism has been

invoked to explain the most prominent cold eventduring the last deglaciation, the YD at 12,700–11,600cal BP, by the routing change of the Laurentide icesheet meltwater from the Mississippi drainage sys-

Ž .tem to the St. Lawrence River Rooth, 1982 . Thecooling effect extended upwind in North America as

Ž .far as Ohio Wright, 1989 . Meanwhile areas fartherŽ .inland e.g., Minnesota were affected not only by

the proximity to the ice sheet in southern Canada butby the influence of the milder Caribbean air mass,thereby accounting for non-analogue vegetation that

Žcharacterized the region in the late-glacial see Sec-.tion 6.3 for details . Change in oceanic ventilation

was also associated with the 200-year-long PB atŽ;10,900 cal BP Lehman and Keigwin, 1992;

.Bjorck et al., 1996 . The final collapse of ice sheetsänd catastrophic drainage of glacial lakes might havetriggered the most prominent Holocene cooling event

Žat 8200 cal BP Alley et al., 1997; Barber et al.,.1999 , although another interpretation of this event

Ž .has been offered Hu et al, 1999 .Although at a global scale the exact causes of

rapid climatic oscillations, including the YD event,are still incompletely understood, accumulating evi-dence from different regions and climate modellingwill ultimately shed light on their causes and mecha-nisms and the Earth’s climatic dynamics as a whole.Here we assume that the climatic signals during thelast deglaciation originated from meltwater-inducedchange in North Atlantic thermohaline circulation, asindicated by most available empirical evidence, andwe discuss recent modeling results that point to themechanism of transmittal of climatic signals into thecontinental interior of North America.

The results from general circulation modelsŽ .GCMs suggest possible mechanisms of cause andtransmittal of climatic signals. On the other hand, thegeographic distribution and magnitude of the cli-matic events, as indicated by empirical records, areimportant for testing modeled mechanisms. Differentforcings would result in various climatic responses indifferent geographic regions. Using an energy bal-

Ž .ance climate model, Harvey 1989 simulated theclimatic response during the YD to several hypo-thetical causes, including fresh meltwater lens withextensive sea ice, iceberg flooding, reduction inNADW, reduction in northward heat transport, andreduction in atmospheric CO concentration. Rind2

Ž .and Overpeck 1993 systematically examined theclimatic responses to hypothetical causesrforcings

Ž .of climatic variability, including 1 inherent randomŽ .variability in the atmosphere; 2 inherent or forcedŽ .variability in the ocean system; 3 solar variability;

Ž .4 variability in volcanic aerosol loading in theŽ .atmosphere; and 5 variability in atmospheric trace

gases. By using GISS GCM, Rind and Overpeckfound that different geographic patterns resulted fromdifferent forcings. For example, the cooling of theNorth Atlantic due to THC shutdown will affectmostly those regions adjacent to and downwind ofthe North Atlantic. This is apparently the case inmodeling the impact of a cold North Atlantic on YD

Ž .cooling Rind et al., 1986 . In contrast, decreasedinsolation would affect all latitudes. However, be-cause the differential heating of land and oceancauses regional changes in atmospheric circulations


and advection patterns, the maximum cooling inresponse to decreased insolation tends to occur overinland regions far removed from oceanic influencesŽ .Rind and Overpeck, 1993; Lean and Rind, 1998 .This sensitive response of continental interiors tosmall solar change was suggested by empirical data

Ž .from the northern Great Plains Yu and Ito, 1999 .Ž .Hostetler et al. 1999 modeled the response of the

climate system during a Heinrich event by using anŽ .atmospheric GCM GENESIS and found that mod-

eled temperatures and wind fields exhibit spatiallyvariable responses over the Northern Hemisphere.Because different mechanisms have different signa-tures in time and space, they may allow for discrimi-

Žnation among climatic records Rind and Overpeck,.1993 .

Recently accumulated evidence for deglacial cli-Žmatic oscillations from the North Pacific Kennett

and Ingram, 1995; Behl and Kennett, 1996; Hendy.and Kennett, 1999 , the west coast of North America

ŽEngstrom et al., 1990; Mathewes, 1993; Peteet and.Mann, 1994; Benson et al., 1997 and Rocky Moun-

Žtains Reasoner et al., 1994; Gosse et al., 1995;Reasoner and Huber, 1999; Reasoner and Jodry,

.2000 indicate synchronous response of the climatesystem to the North Atlantic climatic changes. Wasthis signal transmitted from the North Atlantic Oceanto the North Pacific Ocean through the atmosphereor the ocean? Although the flow of water within theconveyor from the North Atlantic to North Pacific

Ž .takes about 1000 years Kennett and Ingram, 1995 ,too slow to explain the near-synchronous climaticresponses, several climatic simulations were con-ducted to explicitly test the transmittal mechanisms.

To explain the presence of the YD climatic rever-Ž .sal around the North Pacific, Peteet et al. 1997Ž .used an atmospheric general circulation model GISS

to test the sensitivity of the Northern Hemisphere airtemperatures to change in North Pacific sea-surfacetemperatures. They found that a colder North Pacificalone has a cooling effect on air temperatures overNorth America and suggested that a reduction inwater vapor due to cold ocean temperatures is thekey element that causes the cooling. They also founddry conditions in southwestern North America inresponse to the North Pacific cooling, an effectsupported by paleoclimatic data from Owens Lake

Ž .basin Benson et al., 1997 .

Using the ECHAM3rLSG-coupled ocean–atmo-Ž .sphere general circulation model OAGCM , Mikola-

Ž .jewicz et al. 1997 found that a cold North Atlanticcould cause North Pacific climatic variability. Dur-ing the YD interval the temporary shutdown ofNADW formation and a decrease in thermohalinecirculation resulting from meltwater input cooled theNorth Atlantic and reduced poleward heat transport.A maximum cooling occurred over the North At-lantic and Europe, but a marked cooling over theentire Northern Hemisphere also occurred. Thisdownstream cooling reached the North Pacific pri-marily through the atmosphere, according to simula-tions of the OAGCM and a radiocarbon tracer modelŽ .Mikolajewicz et al., 1997 . Changes in atmosphereaffect coastal upwelling at the North American westcoast, and the induced surface cooling caused betterventilation of the thermocline waters of the north-eastern Pacific.

As shown in Fig. 14, for the Great Lakes regionand the northeastern part of North America the cool-ing was an upstream effect of a cold North Atlantic,diminishing inland. This is typical geographic pat-

Žterns of oceanic climatic forcing Rind and Over-.peck, 1993 . Upstream cooling effects were also

suggested by the atmosphere GCM results, though atŽ .smaller magnitude Rind et al., 1986; Wright, 1989 .

However, the difference of the Mikolajewics et al.Ž .1997 simulations from previous model results isthat it used a coupled OAGCM and connected NorthAtlantic and North Pacific directly. The relativelysmall cooling at the North American west coast iscaused by the intensification of the northward windsalong the coast. The Mikolajewics et al. simulationsuccessfully explains the evidence from marinerecords in the North Pacific, but it still cannot fullyaccount for the evidence from the interior of NorthAmerica and the west coast reviewed in this paper.The simulation suggests a small temperature depres-sion in the region around 458N and 1008W. Thisregion includes the Colorado Rocky Mountains,where strong evidence for the YD is available from

Žglacial and pollen records Menounos and Reasoner,.1997; Reasoner and Jodry, 2000 .

Two empirical studies from western North Amer-Ž .ica region F and site 11 in Fig. 3 discuss the

transmittal mechanisms of the YD signal in relationŽ .to the GCM simulations. Benson et al. 1997 sug-


Fig. 14. Modeled change in mean near-surface air temperature around North America due to a shutdown of North Atlantic Deep WaterŽ .formation in the ECHAM3rLSG coupled OAGCM Mikolajewicz et al., 1997 . Here shown is the difference between meltwater-induced

Ž .NADW shutdown experiment similar to the Younger Dryas scenario and control run.

gested that the cooling of the North Pacific de-creased the temperature and moisture content of theair mass passing over the middle latitudes of westernNorth America and caused a drier climate there, asindicated by their paleoclimate records and climate

Ž .simulation model Peteet et al., 1997 . Reasoner andŽ .Jodry 2000 found a downslope displacement of the

alpine timberline vegetation during the YD in Col-orado Rocky Mountains and suggested that changein atmospheric circulation in the eastern North Pa-cific would enhance onshore air flow across the

ŽNorth American Cordillera Mikolajewics et al.,.1997 . This enhanced air flow increases orographic

precipitation at higher latitudes and provides favor-able conditions for the expansion of alpine glaciers,though the temperature change alone, determinedfrom the change in timberline vegetation, is suffi-cient to explain the magnitude of YD glacial ad-

Žvances in the Rocky Mountains Reasoner and Jodry,.2000 .

6.3. Amphi-Atlantic contrasts in late-glacial climates

The configuration of the continents framing theNorth Atlantic Ocean has strongly influenced thepatterns of air-mass distribution. Under conditions ofchanging Milankovitch insolation distribution, thisgeography has resulted in the initiation and growthof the Laurentide ice sheet, and during the deglacia-tion phase the dynamics of the ice sheet itself con-trolled the climatic history on both sides of theAtlantic.

North America east of the Rocky Mountains de-rives its moisture primarily from the tropical mar-itime air mass originating over the Caribbean area, asPacific maritime air loses most of its moisture in

Ž .crossing the Western Cordillera Fig. 15 . The samemaritime air from the Caribbean moves up the NorthAmerican east coast and across the Atlantic to Eu-rope with the prevailing westerly winds. This circula-tion pattern dominated during the last interglacialperiod, and temperatures then were even warmerthan today, as indicated by the small size of theGreenland ice sheet and by evidence that the arcticof northern Scandinavia and northwestern Siberia

Ž .was less cold than at present Velichko et al., 1997 .The Laurentide ice sheet began to form during aninterval of declining summer insolation and hadspread as far west as the Rocky Mountains by thetime of the last glacial maximum. During this phaseof expansion, the Laurentide ice sheet was nourishedprimarily by the moist Caribbean air mass. In Eura-sia, the accumulation of ice early in the last glacialperiod was centered in northwestern Siberia. How-ever, during the last glacial maximum the Scandina-vian ice sheet had expanded and effectively blockedthe flow of moisture to the west-Siberian ice, whichbecame less extensive than it had been earlierŽ .Velichko et al., 1997; Mangerud et al., 1999 .

At its maximum and during its wastage, the Lau-rentide ice sheet indirectly controlled the climate ofEurope. Its height and area, when combined with thecontemporaneous Cordilleran ice sheet, caused theWesterly Jet Stream to be divided in its course


Fig. 15. Schematic map illustrating the Caribbean air penetrationhypothesis that explains the lack of evidence for a cold YoungerDryas and other climatic oscillations in central interior NorthAmerica. During the deglacial period, the climatic signal from the

Ž .North Atlantic e.g., the YD was transmitted upstream as far asŽ .Ohio Rind et al., 1986; Wright, 1989 , affecting the Great Lakes

region as shown by evidence reviewed in this paper. At the sametime, the signal was also transmitted downstream to the Pacific

ŽOcean and penetrated as far as the Rocky Mountains Mikolaje-.wicz et al., 1997 . However, in region east of the Rocky Moun-

tains and west of the Great Lakes the cold arctic air mass wastrapped north of the Laurentide ice sheet, so warm Caribbean aircontinued to penetrate northward without interruption, producingtemperate conditions there. The Caribbean influence is indicatedby non-analogous species assemblages, i.e., persistence of temper-ate deciduous species in spruce forests, as reviewed in this paper.See text for further detail and Fig. 3 for locations of paleoclimaticsites.

across the continent, with the southern branch nour-ishing the desert basins of the American Southwestand the northern branch traversing northernmost

ŽCanada north of the ice sheet COHMAP Members,.1988 , bringing arctic air to the North Atlantic and

Siberia. Expanded sea ice shifted the oceanic polarfront far south to the latitude of Portugal, according

Žto the dominance of polar foraminifera Ruddiman.and McIntyre, 1981 , and permafrost developed in

Eurasia as far south as the Alps and in Siberia to aŽcomparable latitude 508; Baulin and Danilova,

.1984 . Cold semi-desert vegetation dominated by

chenopods and Artemisia in Europe and even in theMediterranean lowlands reflected the influence ofthe cold Siberian anticyclone and the low moisturecontent of the cool Atlantic air mass. Trees werelargely confined to refuges in the Italian and Balkan

Ž .highlands Watts et al., 1996; Willis, 1994 , sup-ported by orographic precipitation, although they

Žalso existed in the Hungarian lowlands Willis et al.,.2000

During the glacial period, substantial and rapidclimatic oscillations are recorded in the Greenlandice cores as well as in marine cores and terrestrialrecords. Among these the concentration of ice-rafted

Ž .detritus in marine cores known as Heinrich eventsrepresents the discharge of icebergs to the Labrador

ŽSea and thence around the northern gyre Heinrich,1988; Andrews and Tedescco, 1992; Broecker et al.,

.1992; Bond et al, 1992; Groussett et al., 1993 . Theresulting cap of fresh water caused a reduction indeep-water formation and the return of the oceanicpolar front to the south. The last main Heinrich-typeevent was the YD. During this cool interval, icebergswere supplemented by meltwater from the retreatingice sheet of the St. Lawrence River, and later joinedby the eastward discharge of Glacial Lake Agassiz.The YD cooling interrupted the BOA warm period

Ž .that had followed dissipation of the previous H-1Heinrich event. At this time the oceanic polar front,which had moved northward almost to its modern

Ž .position Ruddiman and McIntyre, 1981 , returned tothe latitude of Portugal. Near-glacial climatic condi-tions were restored in Europe and sustained by thecold Atlantic storms and the cold Siberian air thatwas still diverted from the North American arctic.Cold conditions also returned to northeastern NorthAmerica, where the cold North Atlantic waters cooledthat portion of the Caribbean air mass moving up theeast coast. The famous AnoreastersB of the modernNew England climate are storms characteristic ofthis air flow. They may have been strengthened bythe anticyclonic circulation around the ice sheet,which produced winds in the St. Lawrence Valleythat were strong enough to orient sand dunes as the

Ž .ice retreated to the north Filion, 1987 . Noreastersmay have been more common and intense in the YDbecause of the cooler waters and cooler air mass,extending as far west as Ohio. Model experimentsfor the YD suggest that low temperatures resulting


from an upwind effect extended that far west—butŽ .no farther Rind et al., 1986 .

Meanwhile, in southeastern North America andwest as far as the Rocky Mountains, the Caribbeanair mass played a direct and more immediate role inthe climate of North America during the glacialperiod, even up to the border of the ice sheet. In thesoutheast the very diverse modern tree flora, al-though not so exceptional as that in southeasternChina in a comparable continental position, rivalsthat recorded for the late Tertiary in Europe. This

Ž .comparison led Braun 1947 to postulate that theforests of the southern Appalachian Mountains haveremained little changed since the Tertiary and werenot affected by Pleistocene climatic changes. In aclassic review of Pleistocene biogeography, DeeveyŽ .1947 showed that boreal trees like spruce hadindeed invaded the region during glacial periods, andsince then dozens of pollen diagrams in the southeasthave shown that the forests were modified substan-

Ž .tially Watts, 1980 and that the varied topographyprovided habitats for both cool-temperate andwarm-temperate trees.

Farther west in the center of the continent, theforest vegetation and thus the climatic history duringthe glacial and deglacial periods is more difficult toreconstruct. The evidence is summarized in somedetail below because the contrasts with the Europeanhistory are so striking. Pollen studies show that theclimate in the mid-continent south of the Laurentideice sheet was dominantly temperate and relativelystable, contrasting strongly with the dominantly frigidbut oscillating conditions in Europe, despite the factthat in North America a huge ice sheet was close athand, whereas in Europe the Scandinavian ice sheetwas relatively small.

Only a narrow fringe of permafrost existed inNorth America south of the ice sheet, for most of thearea north of 508 was protected from freezing by theice sheet itself. Tundra was also restricted—in factspruce forest grew on the stagnant ice of terminal

Ž .moraines Florin and Wright, 1969 . The southwardexpansion of spruce forest to 408 latitude indicatesthe prevalence of cool summers, but from there northto 508 the apparent admixture of temperate decidu-ous trees like ash, oak, and elm implies that theCaribbean air mass had an ameliorating effect on theclimate, and that winters did not have the low tem-

perature extremes found today, for it is this factorthat limits the northern range of these deciduoustrees.

These relations are not fully explained by theusual numerical climatic reconstructions that com-pare fossil pollen assemblages with analogous pollensurface samples taken from areas of known meanclimatic conditions, which do not include minor butcritical climatic variables such as extremes. Fossilassemblages that are anomalous because they haveno modern analogues are not included in the analy-sis. The lack of analogues for mid-continental assem-blages during deglaciation reflects the interplay oftwo major climatic controls that are different fromthe present—stronger summer insolation patterns thatcontrol seasonality and atmospheric circulation, andthe existence of a major ice sheet in mid-latitudes. Itis hypothesized that the intervals of extreme cold inwinter such as those experienced today in the mid-continent did not occur during the glacial periodbecause arctic air was trapped north of the ice sheet—the very air that was diverted instead over the poleto the North Atlantic and Siberia, as mentionedabove. The arctic air that did flow south off the icesheet was adiabatically warmed and reduced in hu-

Žmidity, as in foehns or chinooks Bryson and Wend-.land, 1967 . Conditions in the periglacial area per-

mitted temperate deciduous trees to survive in thedominant spruce forest, which itself extendedthroughout the Midwest because of cool summersengendered by the ice sheet to the north.

Ž .For example, King 1973 showed that in Mis-souri the closed spruce pollen assemblage was suc-ceeded by an assemblage that he called Aspruce withdeciduous elementsB, dated at about 13,500 14C BP.It contained about 20% pollen of oak, elm, iron-wood, and hazel in an assemblage dominated by

Ž .spruce 30% . A similar combination was found byŽ .Andersen 1954 in southeastern Michigan, where

the 30% spruce pollen was combined with 30%temperate types, which, however, were attributed notto local occurrence but to redeposition from uniden-tified older interglacial sediments, a common alterna-tive interpretation of non-analogue assemblages. Ap-parently in these regions well south of the ice sheetthe winter climate was already ameliorating, contem-poraneous with the temperate BOA conditions aroundthe North Atlantic.


Farther north in the Midwest, a less diverse butsignificant admixture of temperate tree taxa alsooccurs in the spruce pollen zone. A compilation of20 pollen sites in the Minnesota area for the sprucepollen zone, which dates from about 16,000 to 11,00014C BP, shows about 2–5% each for ash, oak, elm,and ironwood. Even the underlying herb zone atmany sites has 2–5% oak and ash pollen, although itmay contain macrofossils of tundra plants. This zoneis also marked by pollen of the open-ground temper-

Ž . Žate plants Artemisia 10–40% and Ambrosia 5–.10% , in addition to the 40% spruce pollen. Many

sites also contain 1% pollen of Typha, another taxonrare in the boreal forest today. The modern borealforest and forest-tundra of Canada, their closest gen-eral analogues, lack these temperate taxa, and pollensurface samples from that region do not show such

Žvalues, nor do pollen traps Ritchie and Lichti-.Federovich, 1967 . It is commonly considered that

unexpected pollen types, if not accounted for byredeposition, represent distant transport to an area oflow local pollen production. Although this explana-tion is likely for pine pollen in tundra settings, it isnot valid for the spruce forest that extended even upto the ice front. Spruce produces abundant pollen,but, unlike pine pollen, its pollen is not easily trans-ported long distances by wind. Of course, the findingof macrofossils provides undeniable evidence forlocal occurrence of the plant, but macrofosssils ofthe critical temperate plants are notoriously difficultto find in lake deposits, on which most studies are

Ž .made. In Europe, Kullman 1998 reported macro-fossils of oak, hazel, and elm in an otherwise domi-nant boreal-forest assemblage in a peat deposit in theSwedish Mountains at a site well above the modernrange of these temperate taxa dated to the time verysoon after ice retreat. The occurrence of the temper-ate types in pollen diagrams had previously beenattributed to distant transport. This combination ofboreal and temperate plants has no analogue today inScandinavia and can be attributed to the same typeof anomalous climatic conditions called upon belowfor the Minnesota assemblage.

Pollen percentages may yield a deceiving pictureof vegetation, and a more accurate estimate of thepollen deposition of the anomalous types in questioncan be made by the determination of pollen influx,which requires an accurate chronology. One suitable

site for such an estimate is Kylen Lake in northeast-Ž .ern Minnesota Birks, 1981 , where the non-calcare-

ous terrain provides no basis for supposing that theradiocarbon dates and thus the chronology or influxcalculations are in error. Here in sediments dated13,400 to 12,400 14C BP, Ambrosia-type pollenŽ . y2 y17% has an influx of about 200 grains cm year ,

Ž . Ž .Artemisia 10% an influx of 400, and spruce 20%an influx of 800 grains cmy2 yeary1. Similar influxvalues were obtained for the spruce pollen zone at

ŽElk Lake in northwestern Minnesota Whitlock et al.,. Ž1993 , Rutz Lake in south-central Minnesota Wad-

.dington, 1969 , and Lake West Okoboji in north-Ž .western Iowa Van Zant, 1979 . These three sites

were occupied by prairie in the mid-Holocene, whenthe influx values of Ambrosia and Artemisia weregenerally not much higher than they were in thelate-glacial spruce zone, as prairie plants dispersemuch less pollen than trees. The pollen-trap collec-

Ž .tions of Ritchie and Lichti-Federovich 1967 in themodern boreal forest and tundra of Canada containonly trace amounts of Ambrosia, despite the highpollen production by this weed in modern agricul-tural areas of the Great Plains that can provide asource. These pollen types are unlikely to haveblown to the Minnesota area from the south duringthe glacial period, because spruce forest extended atthis time at least as far south as Kansas, and theclosest open land was in Texas, where the pollenassemblage was dominated by grasses and containedlittle Ambrosia.

A hypothetical climatic reconstruction for condi-tions that would permit the occurrence of Ambrosiaand temperate trees in a landscape dominated byboreal trees, even containing tundra types like Dryas,emphasizes the unique conditions of a landscapebordering a huge ice sheet at a time with substan-tially higher summer insolation than today—condi-tions for which we should not anticipate detailedmodern analogues. Winters would not have had thetemperature extremes that today limit the northernrange of temperate hardwoods, because the frigidarctic air was trapped north of the ice sheet. Wherethe landscape had a diversified topography, the spruceforest would have been confined to sheltered valleys,and the uplands would have been host to shrubs withscattered wind-resistant larch trees, as in the modernrolling landscapes of southern Labrador. Dry habitats


in the spruce forest, such as sandplains with littlecompetition, can be occupied even today by Dryas, atypical tundra plant. In other openings in the late-gla-cial spruce forest, drifting snow may have protectedthe ground cover from winter winds, and aftersnowmelt, when the surface was warmed and driedby adiabatic winds and by the high insolation oflate-glacial summers, they could have provided ahabitat for temperate herbs like Ambrosia.

Anomalous pollen assemblages during the glacialand deglacial periods in North America have their

Žcounterparts among the fossil beetles Morgan et al.,.1983 . The so-called disharmonious vertebrate fau-

Ž .nas of this time Lundelius et al., 1983 also indicateclimatic conditions found nowhere today on the con-tinent.

The Allerød-like amelioration of Midwestern cli-mate, occasioned by the steady flow of Caribbean airmodulated by the cool periglacial summers necessaryto sustain spruce, expanded northward after the last

Žmajor ice advance in Iowa Des Moines lobe, 14,00014 .C BP . It reached southern Minnesota by 12,00014C BP, when spruce pollen started a rapid decline in

Žfavor of early-successional tress ash, birch, and.alder , followed by elm and oak. In areas farther

east, e.g., from Ohio eastward, this Allerød-typewarming was abruptly interrupted by the YD, whichbrought the return of spruce, fir, and larch, followedimmediately by pine, which had been progressingrapidly westward from ice-age refuges in the Ap-palachian highlands. Pine was blocked in its expan-sion beyond Ohio because the climate was alreadytoo warm, not having been affected by the YDcooling, and to the north it was blocked by LakeMichigan and the ice sheet. When the ice sheet andits proglacial lakes withdrew, pine moved veryrapidly into Wisconsin and Minnesota, replacing theearly successional forest and, to the north, the spruceforest itself.

This vegetational sequence west of the GreatLakes reflects the direct influence of the Caribbeanair mass in that area throughout the deglacial period,while areas farther east and especially in Europeunderwent the strong climatic perturbations that werethemselves an indirect effect of the deglaciation pro-gression of the Laurentide ice sheet, which of coursewas built up by the Caribbean air mass in the firstplace. The contrast is also represented by the differ-

ence in the tree diversity in Europe and easternNorth America. The repeated episodes of major Lau-rentide glaciation and the accompanying severe cli-matic conditions in Europe reduced the temperatetree flora that had characterized the preglacial land-scape, whereas in southeastern United States notnearly so many taxa were lost.

An exception to the concept that the Laurentideice sheet had a limited role in modifying the late-gla-cial climate of its periglacial area is emphasized byconditions resulting from the unique event involvingthe collapse of the central ice dome over HudsonBay, resulting from the penetration of a calving frontup Hudson Strait about 7900 14C BP. When thecollapse occurred, the blockage of frigid arctic airnorth of the ice sheet was lost, and the Tyrell Seathat subsequently occupied the depressed HudsonBay lowland was filled with icebergs supplied byremaining portions of the ice sheet in Quebec andKeewatin. The frigid air mass that was generatedover the cold sea swept southward even in summer,resulting in an abrupt decrease in values of d

18 O ofcalcite in the varved sediments of Deep Lake innorthern Minnesota, according to the chronology and

Ž .hypothesis of Hu et al. 1999 . This event, which isweakly recorded if at all around the North Atlanticseaboard, is unlike the earlier late-glacial fluctuationssuch as the YD, for it was a local mid-continentdirect response to the Laurentide ice sheet control,rather than an upwind or downwind reflection of aNorth Atlantic event.

In summary, once the Laurentide ice sheet wasbuilt by moisture primarily supplied by the Caribbeanair mass, that ice sheet provided the principal controlon the climate of Europe. This was accomplishedduring full-glacial time by the rerouting of arctic airand during deglacial Heinrich events and the YDinterval by the yield of meltwater to the NorthAtlantic. Icebergs were of particular importance be-cause of their stored AcoldthB. In contrast, the Scan-dinavian ice sheet had little direct influence on Euro-pean climatic events; rather, it expanded and decayedin response to the Atlantic and Siberian climaticforces, just as did the glaciers in Britain and theAlps. In this reconstruction, the trapping of arctic airto the north of the Laurentide ice sheet and itsdiversion to the east also allowed Caribbean air topenetrate deeply into the interior of North America


Ž .Fig. 15 . The striking climatic oscillations in Europeduring the deglacial period thus contrasted stronglywith the more uniform and temperate conditions ofinterior North America east of the Rocky Mountainsand west of the Great Lakes.

7. Future research directions

7.1. Multiple proxy studies

The regional variations in evidence for the YDevent as summarized above suggest the differentexpression of the YD in paleoecological records,depending on geographic location and character of aparticular site. In some regions, the YD event may

Žnot be expressed as a cold interval see An et al.,.1993; Roberts et al., 1993; Kneller and Peteet, 1999 .

Some proxies may be silent due to their insensitivityto climatic change at the critical time, for example,insensitive response of non-ecotonal vegetationŽ .Shane, 1987; Wright, 1989 . Stable isotopes appear

Žto be very sensitive to climatic change Wright,.1984; Ammann, 2000 . Thus, more investigations

are needed on stable-isotope analysis of carbonatesfrom central North America. In circumstances wherebulk sedimentary carbonates are not reliable for pale-oclimatic research, carbonate shells of molluscs and

Žostracodes may be used e.g., Lister, 1988; vonGrafenstein et al., 1992, 1998, 1999; Yu and Eicher,

.1998 . At some lakes carbonates may be totallylacking, and in these cases, cellulose of aquaticmacrophytes or hydrogen isotopes of individual

Žbiomarkers may provide an alternative Krishna-.murthy et al., 1995; Sauer et al., 1999 .

The use of multiple proxy records has the poten-tial to add significant detail to the nature of climaticchange. The multiproxy records from the easternGreat Lakes region may correlate with records fromthe Atlantic Seaboard but may imply climatic changesof a different nature, such as changes in seasonalityand balance of thermal and moisture conditions,rather than simply temperature and precipitation. It isessential to have multiple proxy data to investigatemany different aspects of the YD and other climaticoscillations. The apparent contrast in late-glacial cli-mates between central North America and the At-lantic region could also be tested using multiple

proxy data, especially if proxy records are able toreveal climatic information concerning extremes orseasonality rather than usual average climates.

7.2. Site characteristics and selection strategy

Because complicated factors may be responsiblefor a paleoecological signal from large lakes, such asthe Great Lakes, small lake sites seem to be moresuitable for detecting late-glacial climatic changes.The clear evidence from Crawford Lake and othersmall sites in the eastern Great Lakes region impliesthat the site characteristics are important in providingsuitable isotopic records. Sites in climatically sensi-tive ecotonal regions are especially suitable for pale-oclimatic studies. High sampling resolution is re-quired for detecting brief century-scale events.

In humid temperate regions, oxygen isotopes canbe used as indicator of air temperatures, but insemi-arid and arid regions they may reflect climateindirectly through evaporation. Thus, non-climaticnoises, such as from local hydrological change,should be considered in a paleoclimatic investiga-tion. Otherwise, strong noise from local hydrologicaleffects may obscure the climatic signals from stableisotopes.

8. Summary

Ž .1 In the Great Lakes region paleoclimatic inter-pretation of oxygen-isotope records from severalsmall lakes indicates a classic climatic sequenceduring the last deglaciation that is comparable withrecords from Europe and Greenland. Some of theseclimatic oscillations, especially the YD cold reversal,have also been recorded in upland and aquatic vege-tation and glacier readvances.

Ž .2 Along the Rocky Mountains, the YD event isrecorded by alpine glacier advance at sites fromAlberta to Colorado and by shift in timberline vege-tation in Colorado.

Ž .3 In the interior region between the RockyMountains and the Great Lakes, evidence for theseclimatic oscillations are generally lacking. However,in the southern High Plains dry and warm climaticconditions as indicated by eolian activity appeared tocoincide with the YD interval.


Ž .4 The regional variations in climatic evidencesuggest that climatic oscillations may have differentexpression in paleo-records, depending on geo-graphic location and characteristics of a particularsite. Small lakes with limited local hydrologicalcomplications in climatically sensitive ecotonal re-gions are ideal for paleoclimatic investigation. Theuse of multiple proxy records has the potential toreveal the nature of climatic change, such as changein seasonality and extremes.

Ž .5 The geographic extent and magnitude of thedeglacial climatic oscillations across North Americasuggest that they are an expression of widespreadclimatic changes rather than locally induced events.Together with simulation results from general circu-lation models, the compiled evidence suggests thatclimatic signals were likely carried over the NorthHemisphere through the atmosphere, producing ei-

Ž .ther upwind the Great Lakes region or downwindŽ .effects the Rocky Mountains .

Ž .6 The lack of evidence for a cold YD and otherclimatic oscillations in interior North America eastof the Rocky Mountains and west of the Great Lakeswas probably caused by the trapping of cold arcticair mass north of the Laurentide ice sheet and byuninterrupted northward penetration of warmCaribbean air. This strongly contrasted climate alsoproduced more temperate conditions and non-analo-gous biological assemblages in this region.

Acknowledgements

ŽWe thank Femke Wallien Earth Sciences—.Elsevier Science BV for her suggestion of writing

this review; Ueli Eicher and contributors of data setsused in this review for making numerical data avail-able; and Brigitta Ammann, Bob Johnson, Mel Rea-soner, Pierre Richard and Bryan Shuman for helpfulsuggestions and comments on the manuscript. This isLimnological Research Center contribution 5 bs.

References

Ahlberg, K., Almgren, E., Wright Jr., H.E., Ito, E., Hobbie, S.,1996. Oxygen-isotope record of Late-Glacial climatic changein western Ireland. Boreas 25, 257–267.

Alley, R.B., 1998. Icing the North Atlantic. Nature 392, 335–336.Alley, R.B., 2000. The Younger Dryas cold interval as viewed

from central Greenland. Quat. Sci. Rev. 19, 213–226.Alley, R.B., Clark, P., 1999. The deglaciation of the northern

hemisphere: a global perspective. Annu. Rev. Earth Planet.Sci. 27, 149–182.

Alley, R.B., Meese, D.A., Shuman, C.A., Gow, A.J., Taylor,K.C., Grootes, P.M., White, J.W.C., Ram, M., Waddington,E.D., Mayewski, P.A., Zielinski, G.A., 1993. Abrupt increasein Greenland snow accumulation at the end of the YoungerDryas event. Nature 362, 527–529.

Alley, R.B., Mayewski, P.A., Sowers, T., Stuiver, M., Tayler,K.C., Clark, P.U., 1997. Holocene climatic instability: aprominent, widespread event 8200 yr ago. Geology 25, 483–486.

Ammann, B., 1989. Response times in bio- and isotope stratigra-phies to late-Glacial climatic shifts, an example from lakedeposits. Eclogae Geol. Helv. 82, 183–190.

Ammann, B., 2000. Biotic responses to rapid climatic changes:introduction to a multidisciplinary study of the Younger Dryasand minor oscillations on an altitudinal transect in the SwissAlps. Palaeogeogr., Palaeoclimatol., Palaeoecol. 159, 191–201.

Ammann, B., Lotter, A., 1989. Late-Glacial radiocarbon- andpalynostratigraphy on the Swiss Plateau. Boreas 18, 109–126.

Ammann, B., Lotter, A.F., Eicher, U., Gaillard, M.-J., Wohlfarth,B., Haeberli, W., Lister, G., Maisch, M., Niessen, F.,Schluchter, Ch., 1994. The Wurmian Late-glacial in lowland¨ ¨Switzerland. J. Quat. Sci. 9, 119–125.

Ammann, B., Birks, H.J.B., Brooks, S.J., Eicher, U., von Grafen-stein, U., Hofmann, W., Lemdahl, G., Schwander, J., Tobol-ski, K., Wick, L., 2000. Quantification of biotic responses torapid climatic changes around the Younger Dryas—a synthe-sis. Palaeogeogr., Palaeoclim., Palaeoecol. 159, 313–347.

An, Z.S., Porter, S.C., Zhou, W., Lu, Y., Donahue, D.J., Head,M.J., Wu, X., Ren, J., Zheng, H., 1993. Episode of strength-ened summer monsoon climate of Younger Dryas age on theLoess Plateau of Central China. Quat. Res. 39, 45–54.

Andersen, S.T., 1954. A late-glacial pollen diagram from southernMichigan, USA. Dan. Geol. Unders., II Raekke, Nr. 80,140–155.

Anderson, T.W., Lewis, C.F.M., 1992. Climatic influences ofdeglacial drainage changes in southern Canada at 10-8 kasuggested by pollen evidence. Geogr. Phys. Quat. 46, 255–272.

Anderson, T.W., MacPherson, J.B., 1994. Wisconsinan late-gla-cial environmental changes in Newfoundland: a regional re-view. J. Quat. Sci. 9, 171–178.

Anderson, W.T., Mullins, H.T., Ito, E., 1997. Stable isotoperecord from Seneca Lake, New York: evidence for a coldpaleoclimate following the Younger Dryas. Geology 25, 135–138.

Andrews, J.T., 1994. Wisconsinan Late-glacial environmentalchange on the southeast Baffin shelf, southeast Baffin Islandand northern Labrador. J. Quat. Sci. 9, 179–183.

Andrews, J.T., Tedesco, K., 1992. Detrital carbonate-rich sedi-ments, northwestern Labrador Sea: implications for ice-sheet

Ž .dynamics and iceberg rafting Heinrich events in the NorthAtlantic. Geology 20, 1087–1090.


Andrews, J.T., Erlenkeuser, H., Tedesco, K., Aksu, A., Jull,Ž .A.J.T., 1994. Late Quaternary stages 2 and 3 meltwater and

Heinrich events, NW Labrador Sea. Quat. Res. 41, 26–34.Atkinson, T.C., Briffa, K.R., Coope, G.R., 1987. Seasonal temper-

atures in Britain during the past 22,000 years, reconstructedusing beetle remains. Nature 325, 587–593.

Barber, D.C., Dyke, A., Hillaire-Marcel, C., Jennings, A.E., An-drews, J.T., Kerwin, M.W., Bilodeau, G., McNeely, R.,Southon, J., Morehead, M.D., Gagnon, J.-M., 1999. Forcing ofthe cold event of 8,200 years ago by catastrophic drainage ofLaurentide lakes. Nature 400, 344–348.

Baulin, V.V., Danilova, N.S., 1984. Late quaternary permafrost inSiberia. In: Velichko, A., Wright, H.E. Jr., Barnosky, C.Ž .Eds. , Late Quaternary Environments of the Soviet Union.University of Minnesota Press, Minneapolis, MN, pp. 60–78.

Behl, R.J., Kennett, J.P., 1996. Brief interstadial events in theSanta Barbara Basin, NE Pacific, during the last 60 kyr.Nature 379, 243–246.

Bender, M., Sowers, T., Dickson, M.-L., Orchardo, J., Grootes,P., Mayewski, P.A., Meese, D.A., 1994. Climate correlationsbetween Greenland and Antarctica during the past 100,000years. Nature 372, 663–666.

Benson, L., Burdett, J., Lund, S., Kashgarian, M., Mensing, S.,1997. Nearly synchronous climate change in the NorthernHemisphere during the last glacial termination. Nature 388,263–265.

Berglund, B.E., Bjorck, S., Lemdahl, G., Bergsten, H., Nordberg,¨K., Kolstrup, E., 1994. Late Weichselian environmental changein southern Sweden and Denmark. J. Quat. Sci. 9, 127–132.

Birks, H.J.B., 1981. Late Wisconsin vegetational and climatichistory of Kylen Lake, northeastern Minnesota. Quat. Res. 16,322–355.

Birks, H.H., Ammann, B., 2000. Two terrestrial records of rapidclimatic change during the glacial-Holocene transitionŽ .14,000–9,000 calendar years B.P. from Europe. Proc. Natl.Acad. Sci. U. S. A. 97, 1390–1394.

Bjorck, S., 1985. Deglaciation chronology and revegetation in¨northwestern Ontario. Can. J. Earth Sci. 20, 850–871.

Bjorck, S., Berglund, B.E., Digerfeldt, G., 1988. New aspects on¨the deglaciation chronology of southern Sweden. Geogr. Pol.55, 37–49.

Bjorck, S., Kromer, B., Johnsen, S., Bennike, O., Hammarlund,¨D., Lemdahl, G., Possnert, G., Rasmussen, T.L., Wohlfarth,B., Hammer, C.U., Spurk, M., 1996. Synchronized terrestrial-atmospheric deglacial records around the North Atlantic. Sci-ence 274, 1155–1160.

Bjorck, S., Walker, M.J.C., Cwynar, L.C., Johnsen, S., Knudsen,¨K.-L., Lowe, J.J., Wohlfarth, B., INTIMATE Members, 1998.An event stratigraphy for the Last Termination in the NorthAtlantic region based on the Greenland ice-core record: aproposal by the INTIMATE group. J. Quat. Sci. 13, 283–292.

Blunier, T., Schwander, J., Stauffer, B., Stocker, T., Dallenbach,Ä., Indermuhle, A., Tschumi, J., Chappellaz, J., Raynaud, D.,¨Barnola, J.-M., 1997. Timing of the Antarctic Cold reversaland the atmospheric CO increase with respect to the Younger2

Dryas event. Geophys. Res. Lett. 24, 2683–2686.

Blunier, T., Chappellaz, J., Schwander, J., Dallenbach, A., Stauf-¨fer, B., Stocker, T.F., Raynaud, D., Jouzel, J., Clausen, H.B.,Hammer, C.U., Johnsen, S.J., 1998. Asynchrony of Antarcticand Greenland climate change during the last glacial period.Nature 394, 739–743.

Bohncke, S.J.P., 1993. Late glacial environmental changes in theNetherlands: spatial and temporal patterns. Quat. Sci. Rev. 12,707–718.

Bond, G., Lotti, R., 1995. Iceberg discharges into the NorthAtlantic on millennial time scales during the last glaciation.Science 267, 1005–1010.

Bond, G., Heinrich, H., Broecker, W.S., Labeyrie, L., McManus,J., Andrews, J.T., Huon, S., Jantschik, R., Clasen, S., Simet,C., Tedesco, K., Klas, M., Bonani, G., Ivy, S., 1992. Evidencefor massive discharges of icebergs into the glacial NorthernAtlantic. Nature 360, 245–249.

Bond, G., Broecker, W., Johnsen, S., McManus, J., Labeyrie, L.,Jouzel, J., Bonani, G., 1993. Correlations between climaterecords from North Atlantic sediments and Greenland ice.Nature 365, 143–147.

Bond, G., Showers, W., Cheseby, M., Lotti, R., Almasi, P.,deMenocal, P., Priore, P., Cullen, H., Hajdas, I., Bonani, G.,1997. A pervasive millennial-scale cycle in North AtlanticHolocene and glacial climates. Science 278, 1257–1266.

Braun, E.L., 1947. Development of the deciduous forests ofeastern North America. Ecol. Monogr. 17, 211–219.

Broecker, W.S., 1994. Massive iceberg discharges as triggers forglobal climate change. Nature 372, 421–424.

Broecker, W.S., 1997. Thermohaline circulation, the Achilles heelof our climate system: will manmade CO upset the current2

balance? Science 278, 1582–1588.Broecker, W.S., Peteet, D.M., Rind, D., 1985. Does the ocean–

atmosphere system have more than one stable mode of opera-tion? Nature 315, 21–26.

Broecker, W.S., Andree, M., Wolfli, W., Oeschger, H., Bonani,G. et al., 1988. The chronology of the last deglaciation:implications to the cause of the Younger Dryas event. Paleo-ceanography 3, 1–19.

Broecker, W.S., Kennett, J.P., Flower, B.P., Teller, J.T., Trum-bore, S., Bonani, G., Wolfli, W., 1989. Routing of meltwaterfrom the Laurentide Ice Sheet during the Younger Dryas coldepisode. Nature 341, 318–321.

Broecker, W.S., Bond, G., Klas, M., Bonani, G., Wolfli, W.,1990. A salt oscillator in the glacial Atlantic? I. The concept.Paleoceanography 5, 469–477.

Broecker, W.S., Bond, G., McManus, J., Klas, M., Clark, E.,1992. Origin of the Northern Atlantic’s Heinrich events. Clim.Dyn. 6, 265–273.

Brook, E.J., Harder, S., Severinghaus, J., Steig, E.J., Sucher,C.M., 2000. On the origin and timing of rapid changes inatmospheric methane during the last glacial period. GlobalBiogeochem. Cycles 14, 559–572.

Brooks, S.J., Mayle, F.E., Lowe, J.J., 1997. Chironomid-basedlate glacial climatic reconstruction for southeast Scotland. J.Quat. Sci. 12, 161–167.

Bryson, R.A., Wendland, W.M., 1967. Tentative climatic patterns


for some late-glacial and post-glacial episodes in central NorthŽ .America. In: Mayer-Oakes, W.J. Ed. , Life, Land, and Water.

University of Manitoba Press, Winnipeg, Manitoba, pp. 271–298.

Chappellaz, J., Blunier, T., Raynaud, D., Barnola, J.-M., Schwan-der, J., Stauffer, B., 1993. Synchronous changes in atmo-spheric CH and Greenland climate between 40 and 8 kyr BP.4

Nature 366, 443–446.Clark, D.H., Gillespie, A.R., 1997. Timing and significance of

late-glacial and Holocene cirque glaciation in the SierraNevada, California. Quat. Int. 38r39, 21–38.

Clark, P.U., Alley, R.B., Pollard, D., 1999. Northern Hemisphereice-sheet influences on global climate change. Science 286,1104–1111.

COHMAP Members, 1988. Climatic changes of the last 18,000years: observations and model simulations. Science 241,1043–1052.

Coope, G.R., 1977. Fossil coleopteran assemblages as sensitiveŽ .indicators of climatic changes during the Devensian last cold

stage. Philos. Trans. R. Soc. London, Ser. B 280, 313–340.Coope, G.R., Lemdahl, G., 1995. Regional differences in the

Late-glacial climate of northern Europe based on coleopterananalysis. J. Quat. Sci. 10, 391–395.

Coope, G.R., Lemdahl, G., Lowe, J.J., Walkling, A., 1998. Tem-perature gradients in northern Europe during the last glacial-

Ž 14 .Holocene transition 14-9 C kyr BP interpreted fromcoleopteran assemblages. J. Quat. Sci. 13, 419–433.

Cwynar, L.C., Levesque, A.J., 1995. Chironomid evidence for lateglacial climatic reversals in Maine. Quat. Res. 43, 405–413.

Cwynar, L.C., Watts, W.A., 1989. Accelerator-mass spectrometerdates from the bogs of Ballybetagh, Ireland. Quat. Res. 31,377–380.

Cwynar, L.C., Levesque, A.J., Mayle, F.E., Walker, I., 1994.Wisconsinan Late-glacial environmental change in NewBrunswick: a regional synthesis. J. Quat. Sci. 9, 161–164.

Dansgaard, W., White, J.W.C., Johnsen, S.J., 1989. The abrupttermination of the Younger Dryas event. Nature 339, 532–533.

Dansgaard, W., Johnsen, S., Clausen, H., Dahl-Jensen, D., Gunde-strup, N., Hammer, C., Hvidberg, C., Steffensen, J., Svein-bjornsdottir, A., Jouzel, J., Bong, G., 1993. Evidence forgeneral instability in past climate from a 250 kyr ice-corerecord. Nature 364, 218–220.

Davis, P.T., 1988. Holocene glacier fluctuations in the AmericanCordillera. Quat. Sci. Rev. 7, 129–157.

Deevey, E.S., 1947. Biogeography of the Pleistocene. Geol. Soc.Am. Bull. 60, 1316–1416.

Denton, G.H., Hendy, C.H., 1994. Younger Dryas age advance ofFranz Josef Glacier in the southern Alps of New Zealand.Science 264, 1434–1437.

Eicher, U., 1980. Pollen-und Sauerstoffisotopenanalysen anspatglazialen Profilen vom Gerzensee, Faulenseemoos und¨vom Regenmoos ob Boltigen. Mitt. Natur. Ges. Berne, N.F.37, 65–80.

Eicher, U., 1987. Die spatglazialen sowie fruhpostglazialen Kli-¨ ¨maverhaltnisse in Bereiche der Alpen: Sauerstoffisotopenkur-¨ven kalkhaltiger Sedimente. Geogr. Helv. 2, 99–104.

Eicher, U., Siegenthaler, U., 1976. Palynological and oxygen

isotope investigations on Late-Glacial sediment cores fromSwiss lakes. Boreas 5, 109–117.

Elias, S.A., 1997. The mutual climatic range method of palaeocli-mate reconstruction based on insect fossils: new applicationsand interhemispheric comparisons. Quat. Sci. Rev. 16, 1217–1225.

Elias, S.A., Anderson, K.H., Andrews, J.T., 1996. Late Wisconsinclimate in northeastern USA and southeastern Canada, recon-structed from fossil beetle assemblages. J. Quat. Sci. 11,417–421.

Engstrom, D.R., Hansen, B.C.S., Wright Jr., H.E., 1990. A possi-ble Younger Dryas record in southeastern Alaska. Science250, 1383–1385.

Filion, L., 1987. Holocene development of parabolic dunes in thecentral St Lawrence lowland, Quebec. Quat. Res. 28, 196–209.

Florin, M.-B., Wright Jr., H.E., 1969. Diatom evidence for thepersistence of stagnant glacial ice in Minnesota. Geol. Soc.Am. Bull. 80, 695–704.

Fritz, P., 1983. Paleoclimatic studies using freshwater depositsand fossil groundwater in central and northern Canada. Palaeo-climates and Palaeowaters: A Collection of EnvironmentalIsotope Studies. International Atomic Energy Agency, Vienna,pp. 157–166.

Fritz, P., Anderson, T.W., Lewis, C.F.M., 1975. Late-Quaternaryclimatic trends and history of Lake Erie from stable isotopestudies. Science 190, 267–269.

Fritz, P., Morgan, A.V., Eicher, U., McAndrews, J.H., 1987.Stable isotope, fossil Coleoptera and pollen stratigraphy in lateQuaternary sediments from Ontario and New York State.Palaeogeogr., Palaeoclimatol., Palaeoecol. 58, 183–202.

Goslar, T., Kuc, T., Ralska-Jasiewiczowa, M., Rozanski, K.,Arnold, M., Bard, E., van Geel, B., Pazdur, M.F., Szerozyn-ska, K., Wicik, B., Wieckowski, K., Walanus, A., 1993.High-resolution lacustrine record of the Late GlacialrHolo-cene transition in Central Europe. Quat. Sci. Rev. 12, 287–294.

Goslar, T., Arnold, M., Bard, E., Kuc, T., Pazdur, M.F., Ralska-Jasiewiczowa, M., Rozanski, K., Tisnerat, N., Walanus, A.,Wick, B., Wieckowski, K., 1995. High concentration of atmo-spheric 14C during the Younger Dryas cold episode. Nature377, 414–417.

Goslar, T., Arnold, M., Tisnerat-Laborde, N., Czernik, J., Wieck-owski, K., 2000. Variations of Younger Dryas atmosphericradiocarbon explicable without ocean circulation changes. Na-ture 403, 877–880.

Gosse, J.C., Evenson, E.B., Klein, J., Lawn, B., Middleton, R.,1995. Precise cosmogenic 10 Be measurements in western NorthAmerica: support for a global Younger Dryas cooling event.Geology 23, 877–880.

Grigg, L.D., Whitlock, C., 1998. Late-glacial vegetation andclimate change in western Oregon. Quat. Res. 49, 287–298.

Grootes, P.M., Stuiver, M., 1997. Oxygen 18r16 variability inGreenland snow and ice with 10y3- to 105-year time resolu-tion. J. Geophys. Res. C 102, 26455–26470.

Grootes, P.M., Stuiver, M., White, J.W.C., Johnsen, S., Jouzel, J.,1993. Comparison of oxygen isotope records from the GISP2and GRIP Greenland ice cores. Nature 366, 552–554.

Grousset, F.E., Labeyrie, L., Sinko, J.A., Cremer, M., Bond, G. et


al., 1993. Patterns of ice-rafted detritus in the glacial NorthŽ .Atlantic 40–558N . Paleoceanography 8, 175–192.

Hammarlund, D., Edwards, T.W.D., Bjorck, S., Buchardt, B.,Wohlfarth, B., 1999. Climate and environment during the

Ž .Younger Dryas GS-1 as reflected by composite stable iso-tope records of lacustrine carbonates at Torreberga, southernSweden. J. Quat. Sci. 14, 17–28.

Hansen, B.C.S., 1995. A review of late-glacial pollen recordsfrom Ecuador and Peru with reference to the Younger Dryasevent. Quat. Sci. Rev. 14, 853–865.

Hansen, B.C.S., Engstrom, D.R., 1996. Vegetation history ofPleasant Island, southeastern Alaska, since 13,000 yr B.P.Quat. Res. 46, 161–175.

Harvey, L.D.D., 1989. Modelling the Younger Dryas. Quat. Sci.Rev. 8, 137–149.

Hebda, R.J., 1983. Late-glacial and postglacial vegetation historyat Bear Cove Bog, northeast Vancouver Island, B.C. Can. J.Earth Sci. 61, 3173–3192.

Heine, J.T., 1998. Extent, timing, and climatic implications ofglacier advances on Mount Rainer, Washington, USA, at thePleistocenerHolocene transition. Quat. Sci. Rev. 17, 1139–1148.

Heinrich, H., 1988. Origin and consequences of cyclic ice raftingin the Northeast Atlantic Ocean during the past 130,000 years.Quat. Res. 29, 143–152.

Hendy, I.L., Kennett, J.P., 1999. Latest Quaternary North Pacificsurface-water responses imply atmosphere-driven climate vari-ability. Geology 27, 291–294.

Heusser, C.J., 1973. Environmental sequence following the Fraseradvance of the Juan de Fuca lobe, Washington. Quat. Res. 3,284–306.

Heusser, C.J., 1977. Quaternary palynology of the Pacific slope ofWashington. Quat. Res. 8, 282–306.

Hoek, W.Z., 1997. Late-glacial and early Holocene climatic eventsand chronology of vegetation development in the Netherlands.Veg. Hist. Archaeobot. 6, 197–213.

Holliday, V.T., 2000. Folsom drought and episodic drying on theSouthern High Plains from 10,900–10,200 14C yr B.P. Quat.Res. 53, 1–12.

Hostetler, S.W., Clark, P.U., Bartlein, P.J., Mix, A.C., Pisias,N.J., 1999. Atmospheric transmission of North Atlantic Hein-rich events. J. Geophys. Res. D 104, 3947–3952.

Hu, F.S., Wright Jr., H.E., Ito, E., Lease, K., 1997. Climaticeffects of glacial Lake Agassiz in the midwestern UnitedStates during the last deglaciation. Geology 25, 207–210.

Hu, F.S., Slawinski, D., Wright Jr., H.E., Ito, E., Johnson, R.G.,Kelts, K.R., McEwan, R.F., Boedigheimer, A., 1999. Abruptchanges in North American climate during early Holocenetimes. Nature 400, 437–440.

Hughen, K.A., Overpeck, J.T., Peterson, L.C., Trumbore, S.,1996. Rapid climate changes in the tropical Atlantic regionduring the last deglaciation. Nature 380, 51–54.

Huntley, B., Birks, H.J.B., 1983. An Atlas of Past and PresentPollen Maps for Europe: 0–13,000 Years Ago. CambridgeUniv. Press, Cambridge, UK.

Isarin, R.F.B., Bohncke, S.J.P., 1999. Mean July temperatures

during the Younger Dryas in northern and central Europe asinferred from climate indicator plant species. Quat. Res. 51,158–173.

Isarin, R.F.B., Renssen, H., 1999. Reconstructing and modellingLate Weichselian climates: the Younger Dryas in Europe as acase study. Earth Sci. Rev. 48, 1–38.

Isbele, G.A., Hooghiemstra, H., van der Borg, K., 1995. Acooling event during the Younger Dryas chron in Costa Rica.Palaeogeogr., Palaeoclimatol., Palaeoecol. 117, 73–80.

Iversen, J., 1954. The Late-Glacial flora of Denmark and itsrelation to climate and soil. Dan. Geol. Unders. Ser. II 80,87–119.

Ivy-Ochs, S., Schluchter, C., Kubik, P.W., Synal, H.-A., Beer, J.,Kerschner, H., 1996. The exposure age of an Egesen moraineat Julier Pass, Switzerland measured with the cosmogenicradionuclides 10 Be, 26Al, and 36Cl. Eclogae Geol. Helv. 89,1049–1063.

Jacobson Jr., G.L., Webb III, T., Grimm, E.C., 1987. Patterns andrates of vegetation change during the deglaciation of eastern

Ž .North America. In: Ruddiman, W.F., Wright, H.E. Jr. Eds. ,North America and Adjacent Oceans during the Last Deglacia-tion. Geological Society of America, Boulder, CO, pp. 277–288.

Johnsen, S.J., Clausen, H.B., Dansgaard, W., Fuhrer, K., Gunde-strup, N., Hammer, C.U., Iversen, P., Jouzel, J., Stauffer, B.,Steffensen, J.P., 1992. Irregular glacial interstadials recordedin a new Greenland ice core. Nature 359, 311–313.

Kennett, J.P., Ingram, B.L., 1995. A 20,000-year record of oceancirculation and climate change from the Santa Barbara basin.Nature 356, 357–362.

King, J.E., 1973. Late Pleistocene palynology and biogeographyof the western Missouri Ozarks. Ecol. Monogr. 43, 539–565.

Kneller, M., Peteet, D., 1999. Late-glacial and early Holoceneclimate changes from a central Appalachian pollen and macro-fossil record. Quat. Res. 51, 133–147.

Koc Karpuz, N., Jansen, E., 1992. A high resolution diatomrecord of the last deglaciation from the S.E. Norwegian Sea:documentation of rapid climatic changes. Paleoceanography 7,499–520.

Krishnamurthy, R.V., Syrup, K.A., Baskaran, M., Long, A., 1995.Late glacial climate record of midwestern United States fromthe hydrogen isotope ratio of lake organic matter. Science 269,1565–1567.

Kuhry, P., Hooghiemstra, H., van Geel, B., van der Hammen, T.,1993. The El Abra Stadial in the Eastern Cordillera of Colom-

Ž .bia South America . Quat. Sci. Rev. 12, 333–344.Kullman, L., 1998. Non-analogous tree flora in the Scandes

Mountains, Sweden, during early Holocene—macrofossil evi-dence of rapid geographic spread and response to paleocli-mate. Boreas 27, 155–161.

Lean, J., Rind, D., 1998. Climate forcing by changing solarradiation. J. Clim. 11, 3069–3094.

Lehman, S.J., Keigwin, L.D., 1992. Sudden changes in NorthAtlantic circulation during the last deglaciation. Nature 356,757–762.

Lemcke, G., Sturm, M., 1996. d18 O and trace element measure-


ments as proxy for the reconstruction of climate changes atŽ .Lake Van Turkey : preliminary results. In: Dalfes, H.N.,

Ž .Kukla, G., Weiss, H. Eds. , Third Millennium BC ClimateChange and Old World Collapse. NATO ASI Ser., Ser. 1, vol.49, pp. 653–678.

Leonard, E.M., 1998. Is the type-Crowfoot Moraine, Alberta,really a Younger Dryas equivalent? Geol. Soc. Am. Bull. 30Ž .7 , 66, Abstracts with Programs.

Leroy, S.A.G., Zolitschka, B., Negendank, J.F.W., Seret, G.,2000. Palynological analysis in the laminated sediment of

Ž .Lake Holzmaar Eifel, Germany : duration of Late glacial andPreboreal biozones. Boreas 29, 52–71.

Levesque, A.J., Mayle, F.E., Walker, I.R., Cwynar, L.C., 1993. Apreviously unrecognized late-glacial cold event in eastern NorthAmerica. Nature 361, 623–626.

Levesque, A.J., Cwynar, L.C., Walker, I.R., 1994. A multiproxyinvestigation of late-glacial climate and vegetation change atPine Ridge Pond, southwest New Brunswick, Canada. Quat.Res. 42, 316–327.

Lewis, C.F.M., Anderson, T.W., 1989. Oscillations of levels andcool phases of the Laurentian Great Lakes caused by inflowsfrom glacial Lakes Agassiz and Barlow-Ojibway. J. Paleolim-nol. 2, 99–146.

Ž .Lewis, C.F.M., Anderson, T.W., 1992. Stable isotope O and Cand pollen trends in eastern Lake Erie, evidence for a locally-induced climatic reversal of Younger Dryas age in the GreatLakes basin. Clim. Dyn. 6, 241–250.

Lewis, C.F.M., Moore, T.C., Rea, D.K., Dettman, D.L., Smith,A.M., Mayer, L.A., 1994. Lakes of the Huron basin—theirrecord of runoff from the Laurentide ice-sheet. Quat. Sci. Rev.13, 891–922.

Leyden, B.W., 1995. Evidence of the Younger Dryas in CentralAmerica. Quat. Sci. Rev. 14, 833–840.

Lister, G.S., 1988. A 15,000-year isotopic record from LakeZurich of deglaciation and climatic change in Switzerland.¨Quat. Res. 29, 129–141.

Lotter, A.F., 1999. Late-glacial and Holocene vegetation historyand dynamics as shown by pollen and plant macrofossilanalyses in annually laminated sediments from Soppensee,central Switzerland. Veg. Hist. Archaeobot. 8, 165–184.

Lotter, A.F., Eicher, U., Birks, H.J.B., Siegenthaler, U., 1992.Late-glacial climatic oscillations as recorded in Swiss lakesediments. J. Quat. Sci. 7, 187–204.

Lowe, J.J., 1994. North Atlantic Seaboard Programme IGCP-253:climate changes in areas adjacent to the North Atlantic duringthe last glacial-interglacial transition. J. Quat. Sci. 9, 95–198Ž .Special Issue Editor .

Lowe, J.J., NASP Members, 1995. Palaeoclimate of the NorthAtlantic Seaboards during the last glacialrinterglacial transi-tion. Quat. Int. 28, 51–61.

Lowell, T.V., Larson, G.J., Hughes, J.D., Denton, G.H., 1999.Age verification of the Lake Gribben forest bed and theYounger Dryas advance of the Laurentide Ice Sheet. Can. J.Earth Sci. 36, 383–393.

Luckman, B.H., Osborn, G.D., 1979. Holocene glacier fluctua-tions in the middle Canadian Rocky Mountains. Quat. Res. 11,52–77.

Lundelius Jr., E.L., Graham, R.W., Anderson, E., Guilday, J.,Holman, J.A., Steadman, D.W., Webb, S.D., 1983. Terrestrial

Ž .vertebrate faunas. In: Potter, S.C. Ed. , Late-Quaternary Envi-ronments of the United States. The Late Pleistocene, vol. 1.University of Minnesota Press, Minneapolis, MN, pp. 311–353.

Maenza-Gmelch, T.E., 1997a. Vegetation, climate, and fire duringthe late-glacial-Holocene transition at Spruce Pond, HudsonHighlands, southeastern New York, USA. J. Quat. Sci. 12,15–24.

Maenza-Gmelch, T.E., 1997b. Late-glacial—Early Holocene veg-etation, climate, and fire at Sutherland Pond, Hudson High-lands, southeastern New York, USA. Can. J. Bot. 75, 431–439.

Magny, M., Ruffaldi, P., 1995. Younger Dryas and early Holocenelake-level fluctuations in the Jura Mountains, France. Boreas24, 155–172.

Mangerud, J., Anderson, S.T., Berglund, B.E., Donner, J.J., 1974.Quaternary stratigraphy of Norden, a proposal for terminologyand classification. Boreas 3, 109–128.

Mangerud, J., Svendsen, J.I., Astakhov, V.I., 1999. Age andextent of the Barents and Kara ice sheets in Northern Russia.Boreas 28, 46–80.

Marcoux, N., Richard, P.J.H., 1995. Vegetation et fluctuationsclimatiques postglaciaires sur la cote septentrionale gaspesi-enne, Quebec. Can. J. Earth Sci. 31, 79–96.

Mathewes, R.W., 1973. A palynological study of postglacialvegetation changes in the University Research Forest, south-western British Columbia. Can. J. Bot. 51, 2085–3103.

Mathewes, R.W., 1993. Evidence for Younger Dryas-age coolingon the North Pacific coast of America. Quat. Sci. Rev. 12,321–331.

Mathewes, R.W., Heusser, L.E., Patterson, R.T., 1993. Evidencefor a Younger Dryas-like cooling event on the British Columbiacoast. Geology 21, 101–104.

Mayewski, P.A., Meeker, L.D., Whitlow, S., Twickler, M.S.,Morrison, M.C., Bloomfield, P., Bond, G.C., Alley, R.B.,Gow, A.J., Grootes, P.M., Meese, D.A., Ram, M., Taylor,K.C., Wumkes, W., 1994. Changes in atmospheric circulationand ocean ice cover over the North Atlantic during the last41,000 years. Science 263, 1747–1751.

Mayewski, P.A., Meeker, L.D., Twickler, M.S., Whitlow, S.,Yang, Q., Lyons, W.B., Prentice, M., 1997. Major featuresand forcing of high-latitude northern hemisphere atmosphericcirculation using a 110,000-year-long glaciochemical series. J.Geophys. Res. 102C, 26,345–26,366.

Mayle, F.E., Cwynar, L.C., 1995. Impact of the Younger Dryascooling event upon lowland vegetation of Maritime Canada.Ecol. Monogr. 65, 129–154.

Mayle, F.E., Levesque, A.J., Cwynar, L.C., 1993a. Acceleratorspectrometer ages for the Younger Dryas event in AtlanticCanada. Quat. Res. 39, 355–360.

Mayle, F.E., Levesque, A.J., Cwynar, L.C., 1993b. Alnus as anindicator taxon of the Younger Dryas cooling in eastern NorthAmerica. Quat. Sci. Rev. 12, 295–305.

Mayle, F.E., Bell, M., Birks, H.H., Brooks, S.J., Coope, G.R.,Lowe, J.J., Sheldrick, C., Shijie, L., Turney, C.S.M., Walker,M.J.C., 1999. Climate variations in Britain during the last


Ž .glacial-Holocene transition 15.0–11.5 cal ka BP : comparisonŽ .with the GRIP ice-core record. J. Geol. Soc. London, 156 II ,

411–423.McCabe, A.M., Clark, P.U., 1998. Ice-sheet variability around the

North Atlantic Ocean during the last deglaciation. Nature 392,373–377.

Menounos, B., Reasoner, M.A., 1997. Evidence for cirque glacia-tion in the Colorado Front Range during the Younger DryasChronozone. Quat. Res. 48, 38–47.

Mikolajewics, S.H., Crowley, T.J., Schiller, A., Voss, R., 1997.Modelling teleconnections between the North Atlantic andNorth Pacific during the Younger Dryas. Nature 387, 384–387.

Ž .Miller, R.F., 1997. Late-glacial Allerod Younger DryasColeoptera from central Cape Breton island, Nova Scotia,Canada. Can. J. Earth Sci. 34, 247–259.

Miller, R.F., Morgan, A.V., 1982. A postglacial coleopterousassemblage from Lockport Gulf, New York. Quat. Res. 17,258–274.

Moore, T.C., Rea, D.K., Mayer, L.A., Lewis, C.F.M., Dobson,D.M., 1994. Seismic stratigraphy of Lake Huron–GeogianBay and postglacial lake level history. Can. J. Earth Sci. 31,1606–1617.

Moore Jr., T.C., Walker, C.G., Rea, D.K., Lewis, C.F.M., Shane,L.C.K., Smith, A.J., 2000. Younger Dryas interval and out-flow from the Laurentide ice sheet. Paleoceanography 15,4–18.

Morgan, A.V., 1987. Late Wisconsin and early Holocene paleoen-vironments of east-central North America based on assem-blages of fossil Coleoptera. In: Ruddiman, W.F., Wright, H.E.

Ž .Jr. Eds. , North America and Adjacent Oceans During theLast Deglaciation. Geol. Soc. Am., The Geology of NorthAmerica, vol. K-3, pp. 353–370.

Morgan, A.V., 1992. Comparisons between Europe and NorthAmerica for the Younger Dryas: the insect evidence. Geol.Assoc. Can.–Mineral. Assoc. Can. 17, A79, Programs withAbstracts.

Morgan, A.V., Morgan, A., Ashworth, A.C., Matthews Jr., J.C.,1983. Late Wisconsin fossil beetles in North America. In:

Ž .Potter, S.C. Ed. , Late-Quaternary Environments of the UnitedStates. The Late Pleistocene, vol. 1. University of MinnesotaPress, Minneapolis, MN, pp. 354–363.

Mott, R.J., 1994. Wisconsinan Late glacial environmental changein Nova Scotia: regional synthesis. J. Quat. Sci. 9, 155–160.

ŽMott, R.J., Stea, R.R., 1993. Late-glacial AllerodrYounger.Dryas buried organic deposits, Nova Scotia, Canada. Quat.

Sci. Rev. 12, 645–657.Mott, R.J., Grant, D.R., Stea, R., Occhietti, S., 1986. Late-glacial

climatic oscillation in Atlantic Canada equivalent to theAllerodrYounger Dryas event. Nature 323, 247–250.

Osborn, G., Gerloff, L., 1997. Latest Pleistocene and earlyHolocene fluctuations of glaciers in the Canadian and northernAmerican Rockies. Quat. Int. 38r39, 7–19.

Osborn, G., Clapperton, C., Davis, P.T., Reasoner, M., Rodbell,D.T., Seltzer, G.O., Zielinski, G., 1995. Potential glacial evi-dence for the Younger Dryas event in the Cordillera of Northand South America. Quat. Sci. Rev. 14, 823–832.

Pennington, W., 1986. Lags in adjustment of vegetation to climate

caused by the pace of soil development: evidence from Britain.Vegetation 67, 105–118.

Peteet, D.M., 1987. Younger Dryas in North America—modeling,data analysis, and re-evaluation. In: Berger, W.H., Labeyrie,

Ž .L.D. Eds. , Abrupt Climatic Change. Reidel, Dordrecht, pp.185–193.

Peteet, D., 1995. Global Younger Dryas? Quat. Int. 28, 93–104.Peteet, D.M., Mann, D.H., 1994. Late-glacial vegetation, tephra,

and climatic history of southwest Kodiak Island, Alaska.Ecoscience 1, 255–267.

Peteet, D.M., Vogel, J.S., Nelson, D.E., Southon, J.R., Nickmann,R.J., Heusser, L.E., 1990. Younger Dryas climatic reversal innortheastern USA? AMS ages for an old problem. Quat. Res.33, 219–230.

Peteet, D.M., Daniels, R.A., Heusser, L.E., Vogel, J.S., Southon,J.R., Nelson, D.E., 1993. Late- glacial pollen, macrofossils,and fish remains in the northeastern U.S.A.: the YoungerDryas oscillation. Quat. Sci. Rev. 12, 597–612.

Peteet, D.M., Daniels, R.A., Heusser, L.E., Vogel, J.S., Southon,J.R., Nelson, D.E., 1994. Wisconsinan late-glacial environ-mental change in southern New England: a regional synthesis.J. Quat. Sci. 9, 151–154.

Peteet, D., Del Genio, A., Lo, K.K.-W., 1997. Sensitivity ofnorthern hemisphere air temperatures and snow expansion toNorth Pacific sea surface temperatures in the Goddard Institutefor space studies general circulation model. J. Geophys. Res.102, 23781–23791.

Porter, S.C., An, Z., 1995. Correlation between climate events inthe North Atlantic and China during the last glaciation. Nature375, 305–308.

Ralska-Jesiewiczowa, M., Goslar, T., Madeyska, T., Starkel, L.Ž .Eds. , 1998. Lake Gosziaz, Central Poland. Institute of Botany,Polish Academy of Sciences, Krakow.

Rea, D.K., Moore, T.C., Anderson, T.W., Lewis, C.F.M., Dobson,D.M., Dettman, D.L., Smith, A.J., Mayer, L.A., 1994a. GreatLakes paleohydrology: complex interplay of glacial meltwater,lake levels, and sill depths. Geology 22, 1059–1066.

Rea, D.K., Moore Jr., T.C., Lewis, C.F.M., Mayer, L.A., Dettman,D.L., Smith, A.L., Dobson, D.M., 1994b. Stratigraphy andpaleolimnologic record of lower Holocene sediments in north-ern Lake Huron and Georgian Bay. Can. J. Earth Sci. 31,1586–1605.

Reasoner, M.A., 1998. Synchronous response of late Pleistoceneto early Holocene pollen records from high-altitude sites in theRocky Mountains; evidence from terrestrial macrofossil based

Ž .chronologies. Geol. Soc. Am. 30 7 , 66–67, Abstracts withPrograms.

Reasoner, M.A., Huber, U.M., 1999. Postglacial palaeoenviron-ments of the upper Bow Valley, Banff National Park, Alberta,Canada. Quat. Sci. Rev. 18, 475–492.

Reasoner, M.A., Jodry, M.A., 2000. Rapid response of alpinetimberline vegetation to the Younger Dryas climate oscillationin the Colorado Rocky Mountains, USA. Geology 28, 51–54.

Reasoner, M.A., Osborn, G., Rutter, N.W., 1994. Age of theCrowfoot advance in the Canadian Rocky Mountains: a glacialevent coeval with the Younger Dryas oscillation. Geology 22,439–442.


Richard, P.J.H., 1994. Wisconsinan late-glacial environmentalchange in Quebec: a regional synthesis. J. Quat. Sci. 9,´165–170.

Richard, P.J.H., Larouche, A.C., Lortie, G., 1992. Paleophyto-´geographie et paleoclimats postglaciaires dans l’ouest du Bas-´ ´Saint-Laurent, Quebec, Canada. Geogr. Phys. Quat. 46, 151–´172.

Rind, D., Overpeck, J., 1993. Hypothesized cause of decade-to-century-scale climate variability: climate model results. Quat.Sci. Rev. 12, 357–374.

Rind, D., Peteet, D., Broecker, W., McIntyre, A., Ruddiman, W.,1986. The impact of cold North Atlantic sea surface tempera-tures on climate: implications for the Younger Dryas coolingŽ .11–10 k . Clim. Dyn. 1, 3–33.

Ritchie, J.C., Lichti-Federovich, S., 1967. Pollen dispersal phe-nomena in arctic-subarctic Canada. Rev. Palaeobot. Palynol. 3,255–266.

Roberts, N., Taieb, M., Barker, P., Damnati, B., Icole, M.,Williamson, D., 1993. Timing of the Younger Dryas event inEast Africa from lake-level changes. Nature 366, 146–148.

Rooth, C., 1982. Hydrology and ocean circulation. Prog. Oceanogr.11, 131–149.

Rouse, W.R., 1991. Impacts of Hudson Bay on the terrestrialclimate of the Hudson Bay Lowlands. Arct. Alp. Res. 23,24–30.

Rousseau, D.D., Preece, R., Limondin-Lozouet, N., 1998. Britishlate glacial and Holocene climatic history reconstructed fromland snail assemblages. Geology 26, 651–654.

Ruddiman, W.F., McIntyre, A., 1981. The mode and mechanismof the last deglaciation: oceanic evidence. Quat. Res. 16,125–134.

Rutter, N.W., Weaver, A.J., Rokosh, D., Fanning, A.F., Wright,D.G., 2000. Data-model comparison of the Younger Dryasevent. Can. J. Earth Sci. 37, 811–830.

Saarnisto, M., 1974. The deglaciation history of the Lake Superiorregion and its climatic implications. Quat. Res. 4, 316–339.

Sauer, P.E., Eglinton, T.I., Hayes, J.M., Sessions, A.L., Schim-melmann, A., 1999. Compound-specific hydrogen isotopemeasurements of organic matter: a new paleoclimate tool andorganic material tracer. Extended Abstracts of the 19th Inter-national Meeting on Organic Geochemistry, pp. 83–84.

Schwert, D.P., Anderson, T.W., Morgan, A., Morgan, A.V., Kar-row, P.F., 1985. Changes in late Quaternary vegetation andinsect communities in southwestern Ontario. Quat. Res. 23,205–226.

Severinghaus, J.P., Brook, E.J., 1999. Abrupt climate change atthe end of the last glacial period inferred from trapped air inpolar ice. Science 286, 930–933.

Severinghaus, J.P., Sowers, T., Brook, E.J., Alley, R.B., Bender,M.L., 1998. Timing of abrupt climate change at the end of theYounger Dryas interval from thermally fractionated gases inpolar ice. Nature 391, 141–146.

Shane, L.C.K., 1980. Detection of a late-glacial climatic shift incentral mid-western pollen diagrams. American Quaternary

Ž .Association, 6th Biennial Meeting Orono, Maine , Abstracts.p. 171.

Shane, L.C.K., 1987. Late-glacial vegetational and climatic his-

tory of the Allegheny Plateau and the till plains of Ohio andIndiana. Boreas 16, 1–20.

Shane, L.C.K., Anderson, K.H., 1993. Intensity, gradients andreversals in late glacial environmental change in east-centralNorth America. Quat. Sci. Rev. 12, 307–320.

Siegenthaler, U., Eicher, U., 1986. Stable oxygen and carbonŽ .isotope analyses. In: Berglund, B.E. Ed. , Handbook of

Holocene Palaeoecology and Palaeohydrology. Wiley, Toronto,pp. 407–422.

Siegenthaler, U., Eicher, U., Oeschger, H., 1984. Lake sedimentsŽ 18as continental d O records from the glacialrpost-glacial

transition. Ann. Glaciol. 5, 149–152.Sowers, T., Bender, N., 1995. Climate records covering the last

deglaciation. Science 269, 210–214.Steig, E.J., Brook, E.J., White, J.W.C., Sucher, C.M., Bender,

M.L., Lehman, S.J., Morse, D.L., Waddington, E.D., Clow,G.D., 1998. Synchronous climate changes in Antarctica andthe North Atlantic. Science 282, 92–95.

Street-Perrott, F.A., Perrott, R.A., 1990. Abrupt climate fluctua-tions in the tropics: the influence of Atlantic Ocean circula-tion. Nature 343, 607–612.

Stuiver, M., Grootes, P.M., 2000. GISP2 oxygen isotope ratios.Quat. Res. 53, 277–284.

Stuiver, M., Grootes, P.M., Braziunas, T.F., 1995. The GISP2d

18 O climate record of the past 16,500 years and the role ofsun, ocean, and volcanoes. Quat. Res. 44, 341–354.

Taylor, K.C., Lamorey, G.W., Doyle, G.A., Alley, R.B., Grootes,P.M., Mayewski, P.A., White, J.W.C., Barlow, L.K., 1993.The AflickeringB switch of late Pleistocene climate change.Nature 361, 432–436.

Taylor, K.C., Mayewski, P.A., Alley, R.B., Brook, E.J., Gow,A.J., Grootes, P.M., Meese, D.A., Saltzman, E.S., Severing-haus, J.P., Twickler, M.S., White, J.W.C., Whitlow, S., Zielin-ski, G.A., 1997. The Holocene Younger Dryas transitionrecorded at Summit, Greenland. Science 278, 825–827.

Teller, J.T., 1987. Proglacial lakes and the southern margin of theLaurentide ice sheet. In: Ruddiman, W.F., Wright, H.E. Jr.Ž .Eds. , North America and Adjacent Oceans during the LastDeglaciation. Geol. Soc. Am., The Geology of North America,vol. K-3, pp. 39–69.

Thompson, L.G., Mosley-Thompson, E., Davis, M.E., Lin, P.-N.,Henderson, K.A., Cole-Dai, J., Bolzan, J.F., Liu, K.B., 1995.Late glacial stage and Holocene tropical ice core records fromHuascaran, Peru. Science 269, 46–50.´

Van der Hammen, T., Hooghiemstra, H., 1995. The El AbraStadial, a Younger Dryas equivalent in Colombia. Quat. Sci.Rev. 14, 841–851.

Van Geel, B., Coope, G.R., Van der Hammen, T., 1989. Palaeoe-cology and stratigraphy of the Late glacial type section at

Ž .Usselo The Netherlands . Rev. Palaeobot. Palynol. 60, 25–129.

Van Zant, K., 1979. Late glacial and postglacial pollen and plantmacrofossils from Lake West Okoboji, northwestern Iowa.Quat. Res. 12, 358–380.

Velichko, A.A., Kononov, Y.M., Faustova, M.A., 1997. The lastglaciation of Earth: size and volume of ice-sheets. Quat. Int.41, 43–51.


von Grafenstein, U., Erlenkeuser, H., Muller, J., Kleinmann-Ei-senmann, A., 1992. Oxygen isotope records of benthic ostra-cods in Bavarian lake sediments. Naturwissenschaften 79,145–152.

von Grafenstein, U., Erlenkeuser, H., Kleinmann, H., Muller, J.,Trimborn, P., 1994. High-frequency climatic oscillations dur-ing the last deglaciation as revealed by oxygen-isotope records

Ž .of benthic organisms Ammersee, southern Germany . J. Pale-olimnol. 11, 349–357.

von Grafenstein, U., Erlenkeuser, H., Muller, J., Jouzel, J.,Johnsen, S., 1998. The cold event 8200 years ago documentedin oxygen isotope records of precipitation in Europe andGreenland. Clim. Dyn. 14, 73–81.

von Grafenstein, U., Erlenkeuser, H., Brauer, A., Jouzel, J.,Johnsen, S.J., 1999. A mid-European decadal isotope-climaterecord from 15,500 to 5000 years B.P. Science 284, 1654–1657.

Waddington, J.C.B., 1969. A stratigraphic record of the polleninflux to a lake in the Big Woods of Minnesota. Spec.Pap.-Geol. Soc. Am. 123, 263–282.

Walker, M.J.C., 1984. Pollen analysis and Quaternary research inScotland. Quat. Sci. Rev. 4, 369–404.

Walker, I.R., Mott, R.J., Smol, J.P., 1991. Allerod-Younger Dryaslake temperature from midge fossils in Atlantic Canada. Sci-ence 253, 1010–1012.

Walker, M.J.C., Bohncke, S.J.P., Coope, G.R., O’Connell, M.,ŽUsinger, H., Verbruggen, C., 1994. The Devensian Weich-

. Žselian Late glacial in North-west Europe Ireland, UK, North.Belgium, The Netherlands, North Germany . J. Quat. Sci. 9,

109–118.Watts, W.A., 1977. The Late Devensian vegetation of Ireland.

Philos. Trans. R. Soc. London, Ser. B 380, 273–293.Watts, W.A., 1980. The late Quaternary vegetation history of the

southeastern United States. Annu. Rev. Ecol. Syst. 11, 387–409.

Watts, W.A., 1985. Quaternary vegetation cycles. In: Edwards,Ž .K.J., Warren, W.P. Eds. , The Quaternary History of Ireland.

Academic Press, London, pp. 155–185.Watts, W.A., Allen, J.R.M., Huntley, B., 1996. Vegetation history

and palaeoclimate of the last glacial period at Lago Grande diMonticchio, southern Italy. Quat. Sci. Rev. 15, 133–153.

White, J.W.C., Steig, E.J., 1998. Timing is everything in a gameof two hemispheres. Nature 394, 717–718.

Whitlock, C., Bartlein, P.J., Watts, W.A., 1993. Vegetation his-Ž .tory of Elk Lake. In: Bradbury, J.P., Dean, W.E. Eds. , Elk

Lake, Minnesota: Evidence for Rapid Climate Change in theNorth-Central United States. Spec. Pap.-Geol. Soc. Am., vol.276, pp. 251–274.

Willis, K.J., 1994. The vegetational history of the Balkans. Quat.Sci. Rev. 13, 769–788.

Willis, K.J., Rudner, E., Sumegi, P., 2000. The full-glacial forestsof central and southeastern Europe. Quat. Res. 53, 203–213.

Wilson, S.E., Walker, I.R., Mott, R.J., Smol, J.P., 1993. Climateand limnological changes associated with the Younger Dryasin Atlantic Canada. Clim. Dyn. 8, 117–187.

Wohlfarth, B., 1996. The chronology of the last termination: areview of radiocarbon-dated, high-resolution terrestrial strati-graphies. Quat. Sci. Rev. 15, 267–284.

Wright Jr., H.E., 1984. Sensitivity and response time of naturalsystems to climatic change in the late Quaternary. Quat. Sci.Rev. 3, 91–131.

Wright Jr., H.E., 1989. The Amphi-Atlantic distribution of theYounger Dryas paleoclimatic oscillation. Quat. Sci. Rev. 8,295–306.

Yu., Z.C., 1997. Late Quaternary Paleoecology of the SouthernNiagara Escarpment, Ontario, Canada: A Multiple Proxy In-vestigation of Vegetation and Climate History. PhD Disserta-tion, Department of Botany University of Toronto, Toronto.380 pp.

Yu, Z.C., 2000. Ecosystem response to late-glacial and early-Ho-locene climate oscillations in the Great Lakes region of NorthAmerica. Quat. Sci. Rev. 19, 1721–1745.

Yu, Z.C., Eicher, U., 1998. Abrupt climate oscillations during thelast deglaciation in central North America. Science 282,2235–2238.

Yu, Z.C., Ito, E., 1999. Possible solar forcing of century-scaledrought frequency in the northern Great Plains. Geology 27,263–266.

Yu, Z.C., McAndrews, J.H., Eicher, U., 1997. Middle Holocenedry climate caused by change in atmospheric circulation pat-terns: evidence from lake levels and stable isotopes. Geology25, 251–254.

Zhou, W.J., Donahue, D.J., Porter, S.C., Jull, T.A., Li, X.,Stuiver, M., An, Z.S., Matsumoto, E., Dong, G.R., 1996.Variability of monsoon climate in East Asia at the end of thelast glaciation. Quat. Res. 46, 210–239.

Zicheng Yu received his BSc in Physi-Ž .cal Geography 1985 and MSc in Envi-

Ž .ronmental Studies 1988 from PekingŽ .University China , where he taught

from 1988 to 1991. He obtained hisŽ . Ž .second MSc 1992 and PhD 1997 in

Botany from the University of TorontoŽ .Canada . The topic of his PhD researchwas on abrupt climatic changes and dif-ferential biological and hydrological re-sponses using multiple proxy data fromlake sediments in the Great Lakes re-

gion. He worked for about 1 year as a post-doctoral fellow in theLimnological Research Center at the University of MinnesotaŽ .USA on high-resolution climatic record of the last 2000 years inthe northern Great Plains using stable isotopes and trace elements.After that, he spent 2 years as a post-doctoral fellow at the

Ž .University of Alberta Canada working on ecosystem modelingof boreal forested peatlands and carbon dynamics–climate connec-tions. Currently, he is a Research Scientist with the ClimateChange Network at the Canadian Forest Service in Edmonton,working on carbon and vegetation dynamics of boreal upland andlowland forest. His research interests are interdisciplinary and atmultiple temporal scales on vegetation succession and dynamics,past climate change and variability, ecological response to climatefluctuations, carbon dynamics of northern peatlands and climaticconnection, and process-based ecosystem modeling.


H.E. Wright Jr. attended Harvard Uni-Ž .versity AB 1939, PhD 1943 and has

been on the faculty at the University ofMinnesota since 1947, most recently asRegent’s professor of Geology, Ecol-ogy, and Botany and Director of theLimnological Research Center. His re-search expanded from arid-region geo-morphology in New Mexico to studiesof the late-Quaternary history of land-scapes in different climatic and geomor-phic settings, including glaciation of

Minnesota, forest-fire history in Minnesota and Labrador, vegeta-tional and climatic history of the Near East and Greece, glaciationin the Andes of Peru and Bolivia, peatland patterns in Sweden,stable-isotope stratigraphy of Irish lake sediments, and lake devel-opment and vegetational history of the mountains and plains ofsouthwestern Siberia.

Documents

Response of interior North America to abrupt climate oscillations in the North Atlantic region during the last deglaciation