Interhemispheric Linkage of Paleoclimate during the Last Glaciation
Author(s): George H. Denton, Thomas V. Lowell, Calvin J. Heusser, Patricio I. Moreno, Bjørn
G. Andersen, Linda E. Heusser, Christian Schlüchter, David R. Marchant
Source: Geografiska Annaler. Series A, Physical Geography, Vol. 81, No. 2, Glacial and
Vegetational History of the Southern Lake District of Chile (1999), pp. 107-153
Published by: Blackwell Publishing on behalf of the Swedish Society for Anthropology and
Geography
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INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE
DURING THE LAST GLACIATION
BY
G.H. DENTON', C. J. HEUSSER2, T.V. LOWELL3, P.I. MORENO4, B.G. ANDERSEN5,
LINDA E. HEUSSER6, C. SCHLUCHTER7 and D.R. MARCHANT8
'Department of Geological Sciences and Institute for QuaternaryStudies,
University of Maine, Orono, Maine, USA
2100 Clinton Road, Tuxedo, New York, USA
3Department of Geology, University of Cincinnati, Cincinnati, Ohio, USA
4Institute for Quaternary Studies, University of Maine, Orono, Maine, USA
5Institute for Geology, University of Oslo, Oslo, Norway
6Lamont-Doherty Earth Observatory, Palisades, New York, USA
7Institute of Geology, University of Bern, Bern, Switzerland
8Department of Geology, Boston University, Boston, Massachusetts, USA
Denton, G.H., Heusser, C1L,Lowell, T.V., Moreno,P.I., Andersen,B.G., Heusser,L.E.,Schliichter,C. andMarchant,D.R.,
1999: Interhemispheric linkage of paleoclimate during the last
glaciation. Geogr. Ann., 81 A (2): 107-153.
ABSTRACT. Combined glacial geologic and palynologic data
from the southern Lake District, Seno Reloncavf, and Isla Grande
de Chilo6 in middle latitudes (40°35'-42°25'S) of the Southern
Hemisphere Andes suggest (1) that full-glacial or near-full-glacial climate conditions persisted from about 29,400 to 14,550 14C
yr BP in late Llanquihue time, (2) that within this late Llanquihue
interval mean summer temperaturewas depressed 6°- 8°C compared to modem values during major glacier advances into the
outer moraine belt at 29,400,26,760,22,295-22,570,
and 14,55014,805 14C yr BP, (3) that summer temperature depression was as
great during early Llanquihue as during late Llanquihue time, (4)
that climate deteriorated from warmer conditions during the early
part to colder conditions during the later part of middle Llanquihue time, (5) that superimposed on long-term climate deterioration are Gramineae peaks on Isla Grande de Chilo6 that represent
cooling at 44,520-47,110 14C yr BP (T-11), 32,105-35,764 14C yr
BP (T-9), 24,895-26,019 14C yr BP(T-7), 21,430-22,774 14C yr BP
(T-5), and 13,040-15,200 14C yr BP (T-3), (6) that the initial phase
of the glacial/interglacial transition of the last termination involved at least two major steps, one beginning at 14,600 14C yr BP
and another at 12,700-13,000 '4C yr BP, and (7) that a late-glacial
climate reversal of < 2-3° C set in close to 12,200 14C yr BP, after
an interval of near-interglacial warmth, and continued into
Younger Dryas time. The late-glacial climate signal from the
southern Chilean Lake District ties into that from proglacial Lago
Mascardi in the nearby Argentine Andes, which shows rapid ice
recession peaking at 12,400 14C yr BP, followed by a reversal of
trend that culminated in Younger-Dryas-age glacier readvance at
11,400-10,200 14C yr BP.
Many full- and late-glacial climate shifts in the southern Lake
District match those from New Zealand at nearly the same Southern Hemisphere middle latitudes. At the last glacial maximum
(LGM), snowline lowering relative to present-day values was
nearly the same in the Southern Alps (875 m) and the Chilean Andes (1000 m). Particularly noteworthy are the new YoungerDryas-age exposure dates of the Lake Misery moraines in
Geografiska Annaler * 81 A (1999) · 2
Arthur's Pass in the Southern Alps. Moreover, pollen records
from the Waikato lowlands on North Island show that a major
vegetation shift at close to 14,700 14C yr BP marked the beginning
of the last glacial/interglacial transition (Newnham et al. 1989).
The synchronous and nearly uniform lowering of snowlines in
Southern Hemisphere middle-latitude mountains compared with
Northern Hemisphere values suggests global cooling of about the
same magnitude in both hemispheres at the LGM. When compared with paleoclimate records from the North Atlantic region,
the middle-latitude Southern Hemisphere terrestrial data imply
interhemispheric symmetry of the structure and timing of the last
glacial/interglacial transition. In both regions atmospheric warming pulses are implicated near the beginning of Oldest Dryas time
(-14,600 14C yr BP) and near the Oldest Dryas/Billing transition
(-12,700-13,000 14C yr BP). The second of these warming pulses
was coincident with resumption of North Atlantic thermohaline
circulation similar to that of the modern mode, with strong formation of Lower North Atlantic Deep Water in the Nordic Seas.
In both regions, the maximum Bolling-age warmth was achieved
at 12,200-12,500 14C yr BP, and was followed by a reversal in climate trend. In the North Atlantic region, and possibly in middle
latitudes of the Southern Hemisphere, this reversal culminated in
a Younger-Dryas-age cold pulse.
Although changes in ocean circulation can redistributeheat between the hemispheres, they cannot alone account either for the
synchronous planetarycooling of the LGM or for the synchronous
interhemispheric warming steps of the abruptglacial-to-interglacial transition. Instead, the dominant interhemispheric climate
linkage must feature a global atmospheric signal. The most likely
source of this signal is a change in the greenhouse content of the
atmosphere. We speculate that the Oldest Dryas warming pulse
originated from an increase in atmospheric water-vapor production by half-precession forcing in the tropics. The major thermohaline switch near the Oldest Dryas/Bolling transition then could
have triggered another increase in tropical water-vapor production to near-interglacial values.
Introduction
A fundamental problem of global climate dynam107
G.H. DENTON ET AL.
ics is to identify mechanisms that drove late Quaternary glacial cycles. Strong statistical linkages
imply that these cycles are ultimately related to the
slow changes in the eccentricity of Earth's orbit
and in the tilt and orientation of its spin axis. But
what physical mechanisms translate the seasonality changes caused by these astronomical factors
into global climate changes? And why is there a
strong non-linear behavior of the late Quaternary
climate system that produces a dominant 100-kyr
signal? To complicate matters, the geologic record
shows numerous large and abruptclimate changes.
One of the most prominent of these occurred during the last termination, when global climate
switched from a glacial to an interglacial mode.
Beyond that, however, isotope and methane
records in Greenland ice cores reveal widespread
abrupt changes throughout the last glacial cycle
(Johnsen et al. 1992). In addition, a pervasive millennial-scale oscillation of the climate system is
commonly recognized (Denton and Karlen 1973;
Grimm et al. 1993; Bond et al. 1997). To explain
these observations, abrupt reorganizations of the
global ocean-atmosphere system have been suggested as fundamental to Earth's climate system
(Broecker and Denton 1990; Imbrie et al. 1992,
1993). Such reorganizations are thought to constitute jumps among stable modes of operation-shifts
that cause changes in the greenhouse gas content or
the albedo of the atmosphere. The challenge, then,
is to understand how orbital forcing interacts with
an ocean-atmosphere system that has a tendency to
undergo abrupt mode changes.
The atmosphere and ocean somehow work together to change the hydrologic cycle of the planet
during abrupt reorganizations of the climate system. Although the exact linkages remain elusive, a
strong possibility is that the roots of at least some
abrupt changes lie within the thermohaline circulation of the ocean (Broecker and Denton 1990).
Numerical modelling studies imply that several
quasi-stable patterns of thermohaline circulation
can exist because the dense water that sinks to the
deep ocean is produced in several locations. Particularly prominent are sources in the Southern
Ocean near Antarctica and in the northern North
Atlantic Ocean. By altering the transport of heat
from low to high latitudes, as well as across the
equator, shifts among these patterns can affect regional or even hemisphere-wide climate. Although
changes in ocean circulation can redistribute heat
on the planet, they cannot by themselves produce
the overall planetary cooling of the last glacial
108
maximum (LGM), nor the global warming of the
last glacial-interglacial transition. Does the explanation of these planetary changes lie instead in the
atmospheric inventory of water vapor (Broecker
1994, 1997a, b)? If so, then how are thermohaline
switches coupled to the production of atmospheric
water vapor, which occurs largely in the tropics? A
coupling between thermohaline switches and tropical ocean-atmosphere dynamics during some
late-glacial abruptclimate changes is suggested by
the tight linkage of North Atlantic atmosphere and
sea-surface temperatures (thermohaline variations) (Bond et al. 1993) with tropical Atlantic upwelling (tradewind strength) (Hughen et al. 1996,
1998) and monsoonal activity in the Arabian Sea
(Schulz et al. 1998). Do the origins of abrupt
changes always lie within the thermohaline circulation of the ocean, or is it possible that some lie
solely in the non-linear production of tropical atmospheric water vapor?
Carefully dated paleoclimate records can reveal
critical characteristics of the climate system during
abrupt changes. The chronology of such records
may reveal the anatomy of a system rich in climate
feedbacks, thus clarifying mechanisms of interhemispheric coupling. Of particular importance
are whether the dramatic rapid changes in the
North Atlantic region are manifested globally,
whether their amplitude varies between the two polar hemispheres, whether there are leads and lags
between the hemispheres, and finally, the phase relations among components of the climate system.
To assess these interhemispheric linkages, we
carried out field investigations in middle latitudes
of the Southern Hemisphere in both South America
and New Zealand. Our prime field area was the
southern Lake District, Seno Reloncavi, and Isla
Grande de Chiloe in southern Chile between 40035'
and 42°25'S (Fig. 1). Here the wide, north-south
trending Valle Central occurs between the Cordillera de los Andes to the east and the Cordillera de
la Costa to the west. Within the northernpart of the
field area, the valley floor is commonly less than
200 m in elevation and features Lago Puyehue,
Lago Rupanco, and Lago Llanquihue. At the LGM
these lakes were filled with piedmont ice lobes fed
by an extensive icefield and mountain-glacier system in the adjacent Andes. Today the southern part
of the valley within the field area is submerged below sea level in the marine basins of Seno Reloncavi, Golfo de Ancud, and northern Golfo Corcovado. These marine basins also were occupied at
the LGM by lobes of piedmont glaciers that adGeografiska Annaler * 81 A (1999) - 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
O
Ca,
Fig. 1. Index map of the field area in the southern Lake District of the Chilean Andes. Valle Central is a large north-south trending
depression filled with lakes and gulfs. The schematic map of the Llanquihue moraine system is derived from the four glacial geomorphological map plates in Andersen et al. (1999). The ice extent for the LGM shown in the inset is from Hollin and Schilling (1981).
The numbers refer to sites discussed in the text as follows: 1, Canal de la Puntilla site; 2, Puerto Octay site; 3, FrutillarBajo site; 4, Puerto
Phillippi site; 5, Fundo Llanquihue site; 6, crossroad site; 7, railroad bridge site; 8, Calle Santa Rosa site; 9, Northwest Bluff site; 10,
Bella Vista Park site; 11, Bella Vista Bluff site; 12, Canal Tenglo site; 13, Punta Penas site; 14, western Puerto Montt site, 15, Huelmo
site; 16, Calbuco site; 17, Taiquem6 site; 18, Dalcahue site. These sites are described in Denton et al. (1999b), Heusser et al. (1999),
and Moreno et al. (1999). In addition, 19 is the Cuesta Moraga site in the Chilean Andes (Heusser 1990) and 20 is the Lago Mascardi
site near Mt. Tronador in the Argentine Andes (Ariztegui et al. 1997).
Geografiska Annaler · 81 A (1999) · 2
109
G.H.DENTONETAL.
vanced westward from the Andes onto both the
mainland and Isla Grande de Chiloe. As a result,
Valle Central and Isla Grande de Chiloe both feature well-preserved moraine-and-outwash systems
of the Llanquihue (last) and older glaciations.
Overall, the spatial distribution of Llanquihue-age
moraines shows that the lowland piedmont lobes at
the LGM reached fartherwestward with increasing
distance to the south, until they overrode southern
Isla Grande de Chilo6 and passed into the open Pacific Ocean. The map pattern and radiocarbon
chronology of these Llanquihue moraines are given in Andersen et al. (1999) and Denton et al.
(1999).
Andean glaciers east of the southern Lake District are now confined largely to the high volcanoes
that surmount the lower crystalline peaks. However, east of Isla Grande de Chiloe small glaciers are
widespread on the lower crystalline Andean peaks
(frontispiece), and icefields mantle the higher volcanoes. Much of Valle Central and eastern Isla
Grande de Chiloe has been cleared of forest since
European settlement. But the natural vegetation is
Lowland Deciduous Evergreen Forest in the lowlands of the southern Lake District, and was Valdivian Evergreen Forest in eastern Isla Grande de
Chiloe. In the Lake District, this transition occurs
at about 41°S at intermediate and low elevations in
the mountains. In the Andes east of the southern
Lake District, successively higher vegetation zones
are composed of North Patagonian Evergreen Forest, and Subantarctic Deciduous Beech Forest to
treeline at 1250 m. Above treeline are high-elevation Andean shrubs and herbs. To the west of the
southern Lake District, and in western Isla Grande
de Chiloe, the coastal range exhibits the following
vegetation belts: North Patagonian Evergreen Forest, SubantarcticEvergreen Forest, and Magellanic
Moorland. As well as being defined by increasing
elevation in mountain ranges on both flanks of the
field area, these vegetation belts occur in distinctive patterns southward through the Chilean channels to Cape Horn. Descriptions of the vegetation
belts, their climate regimes in and near the field area, and their occurrence farther south in Chile at
lower elevations appear in Heusser et al. (1999)
and Moreno et al. (1999).
We chose the area of the southern Chilean Lake
District, Seno Reloncavi, and Isla Grande de Chiloe for detailed study for five reasons. First, the area
lies in the middle latitudes of the Southern Hemisphere, where the orbital seasonality signal is nearly out of phase with that of similar latitudes of the
110
Northern Hemisphere that have produced most
paleoclimate records of the last glacial/interglacial
transition. Second, the field area is within the zone
of dominant Southern Hemisphere westerlies and
thus ideally situated to monitor any northwardshift
of these westerlies during the LGM. Moreover, the
field area is on the western flank of the Andes, and
hence here the westerlies have not been affected by
topographic obstacles (such as the large ice sheets
that are so important in the westerlies belt of the
Northern Hemisphere). Third, the field area is far
removed from the North Atlantic basin, so commonly deemed the critical location on the planet for
triggering climate change. Also, unlike the North
Atlantic region, the area is distant from major
sources of deepwater formation that can greatly affect local ocean heat transport. Fourth, the field
area is at low elevations between central Isla
Grande de Chiloe (42°40'S) and the northern border of the Lake District (37°S), where terminal moraine belts dating to the LGM are accessible for
study and radiocarbon dating. Fifth, the field area
is in the runout zone for numerous pyroclastic
flows from Andean volcanoes. These pyroclastic
flows mantle the moraine belts and outwash plains.
They also afford sediment for interdriftdeposits, as
well as for the infilling of lakes and mires.
To piece together an overall paleoclimate
record, we combined glacial-geologic data with
pollen analysis of sediment cores from mires on
Llanquihue moraines and from interdrift organic
deposits. Four glacial morphologic maps were constructed for the Llanquihue moraine belts between
40°35' and 42°25'S, covering the regions of Lago
Puyehue, Lago Rupanco, Lago Llanquihue, Seno
Reloncavi, Golfo de Ancud, and northern Golfo
Corcovado (Andersen et al. 1999). The chronology
of the glacial deposits is based on more than 450
new radiocarbon dates of samples from stratigraphic sections and from the base of mires on moraine belts (Denton et al. 1999). The reconstruction
of the paleovegetation is based on pollen analysis
of numerous cores from surface mires and of several organic interdrift horizons (Heusser et al.
1999; Moreno et al. 1999), with extensive additional radiocarbon control. The cores were obtained
with a Wright square-rod piston corer. Multiple
overlapping cores were collected within 1 m of
each other at many sites to avoid problems associated with core breaks; these closely spaced cores
were easily correlated by magnetic susceptibility.
Master cores at Fundo Llanquihue (Heusser et al.
1999) and Canal de la Puntilla (Moreno et al. 1999)
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
a.
Southern Lake District, Chile
(40o30'- 42°25'S; 72o25'- 73045'W)
Mean Summer Temperature
b.
Southern Lake District, Chile
(40030'- 42025'S; 72°25'- 73°45'W)
Mean Summer Temperature
14°C
8°C I 10°C 12°C
160C
I
I
C.
Lago Mascardi, Argentina
(41017'S; 71°35'W)
Median Grain Size (pm)
3
4
5
6
7
8
-
.
<
i
-
<
0
-
-
.
II
0
.
-'s
o
~
-
_Ct
(r
L.
N
<-Glacier
1-
I
I
advance-
I
Drald bev R. D. Klv Jr. t1997
Fig. 2. Paleoclimate records for the southern Lake District of Chile and for Lago Mascardi in the adjacent Argentine Andes. Panels (a)
and (b) show a schematic representationof paleoclimate relative to the approximate climate limits of present-day vegetation belts whose
majorconstituents can be recognized in radiocarbon-datedpollen records. The reconstruction for middle Llanquihue time is largely from
Taiquem6 (Heusser et al. 1999), and hence the labels represent Taiquem6 pollen zones. The detailed fluctuations in middle Llanquihue
time come from the percentage of grass pollen (Heusser et al. 1999). We do not know the terminal position of piedmont ice margins
through this portion of Llanquihue time. The reconstruction for late Llanquihue time between 30,000 and 14,550 4C yr BP shows peaks
of glacier expansion into the outer Llanquihue moraine belts (Denton et al. 1999b). These peaks are placed on the diagram by assuming
that the glacier maxima correspond with the most extreme vegetational environment. Such an assumption is generally consistent with
the similar estimates for the lowering of snowline and treeline at the times of the most severe climate deterioration during the LGM
(Porter 1981; Heusser et al. 1999; Moreno et al. 1999). A possible exception, and hence a potential weakness in panels (a) and (b), is
that the glacial record shows a maximum close to 29,400 14C yr BP, whereas the pollen records from Taiquem6 and Dalcahue do not
show a fully developed SubantarcticParklanduntil several thousand years later (Heusser et al. 1999). The detailed pollen diagrams from
the Fundo Llanquihue and Canal de la Puntilla sites both show continued severe conditions during much of the LGM, as depicted by
the dashed line between glacier maxima. The portion of the late Llanquihue record between 14,600 and 10,000 14C yr BP is from pollen
records at Fundo Llanquihue (Heusser et al. 1999), Canal de la Puntilla (Moreno et al. 1999), Huelmo (Moreno 1998), and Taiquemo6
(Heusser et al. 1999). Two major warming steps terminated the LGM. We further subdivide the first step into an early warming pulse
at 14,600 14C yr BP (rise of Nothofagus at Canal de la Puntilla and Huelmo) and a later warming pulse at 14,000 14Cyr BP (invasion of
Subantarctic Parkland by thermophilic trees at many sites). The marked warming to near-Holocene values centered at 12,200-12,500
14C yr BP, followed by a late-glacial climate reversal, is shown at all pollen sites. Because of fire disturbance at many sites the record
for the end of this reversal is shown by a dashed line, representing the situation at Taiquemo6.However, a Younger Dryas signal is shown
in the sediment record from Lago Mascardi, a proglacial lake near Mt. Tronador in the Argentine Andes (Ariztegui et al. 1997) only
15 km east of our key sites of Fundo Llanquihue and Canal de la Puntilla (Fig. 1). The Lago Mascardi sediment record can be locked
into the pollen records from Fundo Llanquihue and Canal de la Puntilla, because all show a distinctive warm peak at 12,200-12,500
4C yr BP, followed by a reversal in climate trend. The subsequent record from Lago Mascardi suggests a Younger-Dryas-age readvance
of the Tronador ice cap (Ariztegui et al. 1997). This is similar to the situation on Isla Grande de Chilo6, where the Taiquem6 pollen
record escaped the influence of fire and shows episodic cooling between 11,360 and 10,355 14C yr BP (Heusser et al. 1999).
were analysed for pollen at intervals of 1 cm or less,
and the chronology of each was controlled with
more than 25 radiocarbon dates. Subsidiary cores
were analysed at more widely spaced intervals and
the chronology controlled with fewer radiocarbon
dates (Heusser et al. 1999). The organic layers revealed in exposures at the Canal Tenglo and DalcGeografiska Annaler · 81 A (1999) - 2
ahue sites were also analysed at intervals of 1 cm
or less; organic horizons in the other sections were
less closely controlled (Heusser et al. 1999).
Chilean Andes paleoclimate record
Figure2 shows a schematicrepresentationof the
111
G.H.DENTONETAL.
paleoclimatechangesinferredfrommorainechronologies and pollen recordsfrom the area of the
southernChileanLake District, Seno Reloncavi,
andIslaGrandede Chiloe.Theresultsaredisplayed
in relationto paleotemperature
estimatesfrompolin termsof fluctuationsof
len diagrams,interpreted
broadvegetationbeltsthroughtime (Heusseret al.
1999;Morenoet al. 1999).Climateparametersfor
each broad modern vegetationbelt in southern
Chilearereviewedin Heusseret al. (1999).Theraoutermorainebeltsof Llanquihue
diocarbon-dated
age arerepresentedin this diagramso thatthe glacier maximacorrespondto the mostextremevegetationalenvironment,an assumptionconsistentnot
only withthe chronologiesof thepollensequences
andthe morainebelts, butalso with the similarestimatesfor the loweringof snowlineandtreelineat
the times of the most severeclimatedeterioration
during the LGM (Porter 1981; Hubbard1997;
Heusseretal. 1999;Morenoetal. 1999).Finally,the
diagramindicatesthe roughboundariesfor early,
middle,andlateLlanquihuetime,basedlargelyon
usagedevelopedin NorthAmerica(Dreimanisand
Goldthwait1973). Marineisotope stage (MIS) 4
correspondsroughlyto earlyLlanquihuetime,MIS
3 to middleLlanquihuetime,MIS2 to lateLlanquihue time, andMIS 1 to the Holocene.
A potentialproblemwith our reconstructionis
thatthe SubantarcticParklandthatoccurredat all
the pollen sites at the LGM does not have an unambiguous modern analog. This parklandwas
closest in characterto theMagellanicMoorland,as
the two communitiesshare many taxa. We think
that the differencebetween the two communities
results from the thick and extensive alluvial fill,
cappedby outwashplains,of ValleCentralandof
easternIsla Grandede Chiloe.A similargeologic
settingdoes notexist in thepresent-dayMagellanic
Moorlandof theouter,cold andwet, rockycoastof
southernmostChile. The result is that in glacial
times grassesand compositesbecamewidespread
within the SubantarcticParkland,where conditions were suitablefor theirgrowthandreproduction on well-drainedoutwashplains;at the same
time,elementsof MagellanicMoorlandfloramade
upthelowlandvegetationin boggyareaswithinthe
poorly drainedmorainebelt. In other words, the
present-dayMagellanicMoorlanddoes not serve
as a strictmodem analogfor the ice-age Subantarctic Parklandbecause of geologic differences,
and yet the two are thoughtto representthe same
climateregimebecausetheyhaveso manycharacteristicplanttaxain common.
112
An importantassumptionof our conceptual
frameworkis thatthelow-slopingformerpiedmont
glaciers of the southernLake District, Seno Reloncavi, Golfo de Ancud, and Golfo Corcovado
werereliableindicatorsof climatechanges.Inotherwords,we assumethatthebehaviorof suchlobes
was dominatedby changes in surfaceconditions
imposed by climate (large variationsof ablation
and accumulationzones on such low-slopingsurfaces), as opposedto changes in basal conditions
(frombasalhydrologyanddeformingbeds).
A contraryview holdsthatsuchlow-slopingglacier lobes are inherentlyunstablewherethey rest
on deformingsediment.This concept arose from
the behaviorof ice streamsthatdrainthe interior
ice reservoirof the marine-basedWest Antarctic
Ice Sheetinto the Ross Ice Shelf (Alley andWhillans 1991a;MacAyeal1993).There,of course,the
ice surfacedoes not have an ablationzone, the ice
streamshave low profiles,and at least Ice Stream
B rests on till. It has been postulatedthatmost of
thefastflowof Ice StreamB resultsfromdeformationof this basaltill layer(Alley et al. 1987;Alley
and Whillans 1991b).These ice streamsundergo
episodic flow. For example,Ice StreamC is stagnanttoday but apparentlywas fast-flowingabout
200 yearsago (Alley andWhillans1991a). Generations of ice streamsseem to replaceeach other,
and it may be thatthe WestAntarcticIce Sheet is
still collapsingslowly as theseice streamsmigrate
headwardinto the interiorice reservoir.An extension of thisconceptis thatsectorsof theLaurentide
Ice Sheetrestingon sedimentarybedrock,particularlyalongthelow-slopinglobatesouthernmargin,
exhibitedsimilarunstablebehaviorthoughtto be
associatedwithdeformingsubglacialbeds.In fact,
Clark(1994) suggestedthat inherentinstabilities
wherethe LaurentideIce Sheet overrodedeforming sedimentled to irregularmarginalfluctuations
which, throughtheirimpacton thermohalinecirculation, caused the abruptclimate oscillations
seen in paleoclimaterecordsfromtheNorthAtlantic region.MacAyeal(1993) postulatedthatcyclic
behaviorof the interiorLaurentideIce Sheetfrom
alternatingthawed and frozen basal conditions
caused surges of carbonate-bearingice through
HudsonStrait,resultingin depositionof detrital,
carbonatelayersin the NorthAtlanticOceanduring Heinrichevents.The key point is that,by this
concept,therapidmarginaloscillationsof theLaurentideIce Sheetarenotthoughtto haveoriginated
in climate.It shouldalso be pointedout thatother
mechanismshave previously been proposedfor
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC
LINKAGEOF PALEOCLIMATE
DURINGTHELASTGLACIATION
some advancesof individuallobes of the southern
LaurentideIce Sheet (Prest 1970; Wright 1973;
Claytonet al. 1985).
Claytonet al. (1989) counteredthatnearlyintact
stratigraphicsequencesof coherentglacial units,
including some with delicate sedimentarystructures, is strongevidence againstpervasivedeformationof the bed beneathformerLaurentideice
lobes in the GreatLakesregion.Instead,they suggested that these lobes are low sloping because
they are supportedby high basalpore-waterpressure within the underlyingsediments,with rapid
flow from sliding at the ice-sediment contact.In
thisway thebasalslidingis akinto a thrustfaultfacilitatedby high waterpressure.Becausebasaltill
canbe draggedbeneaththeice, deformationwithin
the till is a consequenceof, not the cause of, rapid
ice flow.In fact,it now turnsoutthatthis argument
mayalso applyto the low-slopingflankof theWest
AntarcticIce Sheetin theRoss Embaymentsector,
whereKamband Engelhardt(1998) reportedthat
the flow mechanism beneath fast-moving Ice
StreamB is largelybasalslidingat the ice-till contact,ratherthandeformationwithinthe underlying
till layer.
In view of this situation,we mappedandradiocarbondatedthe morainebelts of six formerAndean piedmontglaciersthroughnearly2° of latitude to ensure that the glacier fluctuationswere
widespreadandhence relatedto regionalclimate,
ratherthanrestrictedto the local dynamicsof one
lobe (Andersenet al. 1999;Dentonet al. 1999). In
like manner,we analysedpollen cores over a similar range of latitudeto ensure that the recorded
vegetationchanges were representativeof the region ratherthana single site (Heusseret al. 1999;
Morenoet al. 1999).These two datasets reinforce
our assumptionthatthe majorglacierfluctuations
were a responseto climatechange.Moreover,we
were not able to find a layer of pervasivelydeformed sedimentwithin the Lago Llanquihueor
Seno Reloncavibasins(TurbekandLowell 1999).
The firstimportantpointto drawfromFig. 2 is
that the full-glacial summertemperaturereconstructedfrom paleovegetationwas depressed68°C comparedto modem valuesat this latitudeof
the Chilean Andes, an estimate consistent with
snowline depression of about 1000 m from its
presentpositionof 1900-2100 m in theVolcainCalbuco andLago Todoslos Santossectorof the Andes (Porter 1981; Hubbard 1997). Pollen sites
within 200 m elevation of present-daysea level
alongnearly2° of latitudeuniformlyshowan open
Geografiska Annaler · 81 A (1999) - 2
SubantarcticParklandenvironmentat the LGM.
The paucityof trees suggeststhatthese low-lying
sites wereall thenneartreeline.The implicationis
thattreelineloweredabout 1000 m or more at the
LGMfromits present-daylimit at close to 1250 m
in the adjacentAndes. Thus the temperaturedepressionaffectednot only higherelevationsnear
2000 m (snowline),butalso thelow-elevationfloor
of Valle Centralwithin a few hundredmetersof
present-daysea level (treeline).
A secondmajorpointderivedfromFig. 2 is that
the depression of summer temperaturewas as
greatduringearlyLlanquihueas duringlate Llanquihuetime. This conclusioncomes fromthe moraine morphologyand pollen stratigraphyat the
Taiquemosite in Isla Grandede Chiloe. The glacial morphologicmapnearthe Taiquem6site (fig.
5 in Andersen et al. 1999) depicts a complex
north-south trending Llanquihue-age moraine
belt, which is little-weatheredin sharpcontrastto
the older distal morainebelt. The age of a readvance into this morainebelt at the Dalcahuesite
(18 in Fig. 1) comes froma meanradiocarbondate
of 14,805 14Cyr BP for 34 samplesfrom a paleolandsurfaceburiedby a coarsening-upwardglaciofluvialsequencethatculminatesin basal lodgment till with large (up to 1 m diameter)striated
granitebouldersof Andeanprovenance(Dentonet
al. 1999). The paleolandsurfaceis at the top of a
layerof organicvolcanicfine sandandsilt with an
extensiveseries of radiocarbondates thatextends
back in sequence to 30,070 14Cyr BP (A-7685)
(Heusseret al. 1999;Dentonet al. 1999).Henceat
Dalcahue,andby extensionat all localities in the
Llanquihuemorainebelt at andwest of Dalcahue,
the advanceat 14,805 '4C yr BP was the most extensive of late Llanquihuetime. An organicclast
datedto 14,820+450 (QL-4532),as well as older
organic clasts dating to 19,840+180 '4C yr BP
(UGA-6979) and 21,080+220 '4C yr BP (UGA6972), are reworkedinto ice-contactgravelsnear
the upperlip of a prominentice-contactslope 1.5
kmnorthof theDalcahuesite (Dentonet al. 1999).
This morainecan be tracednorthwardand shown
to lie 4 km east of the Taiquem6pollen site (Andersenet al. 1999).The Taiquem6mire,in turn,is
located within a topographicdepression on the
outermostLlanquihuemoraine ridge in eastern
Isla Grandede Chiloe. Thus, in the region of the
Taiquem6and Dalcahue sites, all but the outermost ridges in this morainebelt are late-Llanquihue in age anddatefromclose to 14,805 '4Cyr BP
(Dentonet al. 1999).
113
G.H. DENTON ET AL.
The pollen record of the Taiquem6 mire comes
from a core 6.55 m long that penetrates to the glacial sediments of the outermost Llanquihue moraine (fig. 20 in Heusser et al. 1999). The chronology of the Taiquem6 pollen record is controlled by
30 radiocarbon dates. The lower 55 cm of the core
dates to > 49,892 '4C yr BP (AA-14770). The level
at 525 cm is dated to 47,110±289314C yr BP (AA14767), which is most likely a minimum age. The
point here is that a substantial portion of the lower
part of Taiquem6 core is close to or beyond the limit of radiocarbon dating.
The pollen record for the Taiquem6 core, given
in fig. 20 of Heusser et al. (1999), shows that a
closed-canopy North Patagonian Evergreen Forest
had come into existence by about 13,000 '4C yr BP
during the last glacial/interglacial transition, with
the peak of late-glacial forest development at
12,200 14C yr BP. This forest featured a suite of
thermophilic trees, including Lomatia, Myrtaceae,
and Maytenus, Gramineae, and even the cold-tolerant elements of the North Patagonian Evergreen
Forest, are at very low levels. Such interglacial conditions are not again encountered deeper in the
Taiquem6 core. There is no evidence from the pollen record, the lithology, or the radiocarbon chronology for hiatuses in the Taiquem6 core. Therefore, we infer that the outer Llanquihue-age moraine at Taiquem6 is younger than the penultimate
interglaciation. Given these age constraints, we
consider it highly likely that this outer moraine was
deposited during early Llanquihue time (probably
at the MIS 4 maximum), as indicated in Fig. 2. We
do not have direct information on the age of the outer Llanquihue-age moraine belt from Taiquem6
northwardto the middle reaches of the Llanquihueage moraines west of Seno Reloncavi. But the radiocarbon chronology given in Denton et al.
(1999) shows that the outer moraine belt fringing
northern Seno Reloncavi, Lago Llanquihue, and
Lago Rupanco is late Llanquihue in age. However,
middle or early Llanquihue advances into subsidiary moraine belts west of Lago Llanquihue are suggested by virtually unweathered till or outwash beneath radiocarbon-datedinterdriftorganic silt-andsand deposits in exposures at Puerto Octay, at Frutillar Bajo, and west of Puerto Varas (Denton et al.
1999).
A third major point illustrated in Fig. 2 is that the
climate deteriorated from warmer conditions during the early partto colder conditions during the later partof middle Llanquihue time (MIS 3). The pollen record from Taiquemo again affords most of the
114
pertinent evidence (fig. 20 in Heusser et al. 1999).
Throughout much of middle Llanquihue time, from
> 49,892 14Cyr BP (AA-14770) until about 26,019
14Cyr BP (AA-14758), the Taiquem6 mire in eastern Isla Grande de Chiloe recorded elements of
Subantarctic Evergreen Forest, except during the
times of the Gramineae peaks described below.
During zones T-12 and T-14, prior to 47,110 14Cyr
BP, Nothofagus occurs in association with arboreal
elements such as Podocarpus nubigena, Pilgerodendron-type, and Pseudopanax laetevirens, along
with low frequencies of Gramineae, Myrtaceae,
Maytenus, and Drimys winteri. This assemblage
implies the presence of Subantarctic Evergreen
Forest in the early phases of middle Llanquihue
time. Later in middle Llanquihue time, a diminished SubantarcticEvergreen Forest persisted until
between 44,520 and 35,764 '4C yr BP; but Subantarctic Parkland expanded thereafter, with the last
remnants of evergreen forest disappearing during
zone T-8 between 32,105 14Cyr BP and 26,019 14C
yr BP. Subantarctic Parkland then dominated until
about 13,000 '4C yr BP.At the nearby Dalcahue site,
remnants of Subantarctic Evergreen Forest lasted
until 25,176 14Cyr BP, then to be replaced by Subantarctic Parkland (Heusser et al. 1999).
A striking feature of the Taiquem6 pollen record
is the succession of Gramineae peaks through middle and late Llanquihue time (Table 1) (figs 20 and
29 in Heusser et al. 1999). The Gramineae peaks
during pollen zones T-9, T- 1, and T- 13 were each
accompanied by declines in Podocarpus nubigena
and Pilgerodendron type. The vegetation during
each of these zones is interpreted as being characteristic of phases of SubantarcticParklandenvironment superimposed on the general deterioration of
Subantarctic Evergreen Forest. The durations of
the youngest two of these middle Llanquihue grass
pollen events are 44,520-47,110 14Cyr BP for zone
T-11 and 32,105-35,764 '4C yr BP for zone T-9.
Grass pollen events also occurred during late Llanquihue time at 24,895-26,019 '4C yr BP for zone T7, 21,430-22,774 14C yr BP for zone T-5, and
13,040-15,200 '4C yr BP for zone T-3 (Heusser et
al. 1999).
In the southern Lake District, the pollen records
for middle Llanquihue time extend back only to
36,960-39,340 '4C yr BP at the Frutillar Bajo and
Puerto Octay sites, and intermittently to 39,660 14C
yr BP at the Canal Tenglo site (Heusser et al. 1999).
These records come not from a continuous core as
at Taiquem6, but from interdrift organic silt-andfine-sand deposits. They cover, then, only the later
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
Table 1. Comparison of the chronology of Heinrich events in the North Atlantic Ocean (40-60°N), Gramineae maxima at Taiquemo,
IslaGrandede Chilo6(42°10'S),andglaciermaximain the southernLakeDistrict-Isla Grandede Chilo6on the westernborderof
the ChileanAndes(40°35'-42°25'S).All ages arein 14Cyr BP.
Heinrichevents
NorthAtlanticOcean
Gramineaeevents
IslaGrandede Chiloe
Glaciermaxima
ChileanAndes
(Elliot et al. 1998)
(Heusser et al. 1999)
(Denton et al. 1999b)
H-1 13,200-15,000
H-2 20,200-22,200
T-3 13,040-15,200
T-5 21,430-22,774
H-3 26,000-27,700
H-4 34,200-35,200
(32,200-35,200)
H-5 44,220
T-7 24,895-26,019
T-9 32,105-35,764
14,550-14,805
22,295-22,570
21,000 (?)
26,760
-
T-11 44,520-47,110
1. The table shows the durations (rather than just the peaks) of Heinrich lithic events. The duration of Heinrich events 1, 2, 3, and 4
aretakenfromthe averageradiocarbon
ages of the baseandtopof the lithiclayersin severalcoresover20° of latitudein the North
AtlanticOcean,as plottedin fig. 5 of Elliotet al. (1998). An alternatechronologyfor the durationof H-4 (givenin parentheses)is
fromthe combinedlithicand 18O0signalin fig. 9 of Elliotet al. (1998). Previousage estimatesfor the H-4 eventareabout35,100
14Cyr. BP(Cortijo et al. 1997) and about 35,500 14Cyr. BP(Bond et al. 1993). The radiocarbon age estimate for H-5 is from Elliot
et al. (1998).Radiocarbon
datesfromNorthAtlanticmarine-sediment
coresarecorrectedby 400 yearsforthesurface-water
marine
reservoireffect.
2. The Gramineae events show the duration, not just the peaks, of pulses of enhanced grass pollen in the Taiquemo vegetation record,
IslaGrandede Chiloe.
3. Thetableshowsonly the peaksof Chileanglacieradvances,not the durationsof the expansionevents.
4. Thedateof 21,000 14Cyr. BP(?) listedunderGlacierMaximain the ChileanAndesrefersto the alternatechronologyof Dentonet
al. (1999).
phase of the much longer middle Llanquihue
record farthersouth at Taiquem6, when Subantarctic Evergreen Forest was giving way to Subantarctic Parkland. The three records from stratigraphic
sections in the southern Lake District show the
continued presence of Subantarctic Parkland during the later phase of middle Llanquihue time, beginning at least by 36,960-39,340 '4C yr BP. These
data indicate that during the later part of middle
Llanquihue time, there was a transition zone in
eastern Isla Grande de Chiloe between Subantarctic Evergreen Forest and Subantarctic Parkland,
while Subantarctic Parkland dominated the southern Lake District.
The trend evident from the vegetation record at
Taiquem6 is of increasingly cold climate beginning
after pollen zone T-12 (fig. 20 of Heusser et al.
1999) and continuing into the beginning of late
Llanquihue time. That overall cold and wet climate
persisted throughout middle Llanquihue time is indicated by the low percentages of the thermophilic
arboreal species of the North Patagonian Evergreen Forest, even during early phases of middle
Llanquihue time, and by the continued presence of
Lepidothamnusfonkii, Astelia pumila, and Donatiafascicularis. Superimposed on this overall cooling trend are stadial-interstadial fluctuations. Subantarctic Parkland conditions marked the stadials
Geografiska Annaler * 81 A (1999) · 2
(T-13, T-11, and T-9 in fig. 20 of Heusser et al.
1999), and reversion toward Subantarctic Evergreen Forest characterized the interstadials (T-14,
T-12, T-10, and T-8 in fig. 20 of Heusser et al.
1999). Subantarctic Evergreen Forest diminished
in dominance during successive interstadials. This
fluctuating deterioration of climate does not leave
a sharp delineation between middle and late Llanquihue time. Rather, the impression from the increase in frequency of Gramineae in the Taiquem6
pollen record is that the climate deterioration continued toward a culmination of cold conditions
about 21,900 14Cyr BP in late Llanquihue time.
The extent of Andean piedmont ice lobes during
middle Llanquihue time is not well constrained.
The Lago Llanquihue piedmont lobe did not advance into the outer moraine belt at the Puerto Octay site between 29,363 and prior to 39,340 14Cyr
BP or at the Frutillar Bajo site between 26,760 and
prior to 36,960 14Cyr BP. Nor is there evidence for
nearby Andean ice at the Canal Tenglo section
alongside Seno Reloncavi between 29,385 and prior to 39,660 14Cyr BP (Denton et al. 1999). These
three sites are all at or near the ice-contact slopes
that rise above Lago Llanquihue or Seno Reloncavi
at the proximal margin of the outer Llanquihue moraine belt. The first definitive evidence that middle
Llanquihue climate had deteriorated enough to
115
G.H. DENTON ET AL.
send Andean piedmont glaciers into the outer Llanquihue-age moraine belt is outwash deposited
about 29,400 '4C yr BP at the Puerto Octay and Canal Tenglo sites.
A fourth major point shown in Fig. 2 is that climate fluctuated within a narrow range close to fullglacial conditions from about 29,400 to about
14,550 '4C yr BP. Throughout this long interval
Subantarctic Parkland environments persisted in
the southern Lake District (Heusser et al. 1999;
Moreno et al. 1999). Andean piedmont glacier
lobes repeatedly advanced into the outer Llanquihue moraine belt during this long interval. As mentioned above, during these peaks of the LGM, we
estimate a mean summer temperatureabout 6-8°C
lower than at present. The only discrepancy between the pollen and glacial records is near the beginning of the LGM. The glacial record implies
that the earliest major advance occurred close to
29,400 14Cyr BP. However, the pollen records from
Isla Grande de Chiloe suggest that full-glacial conditions may not have been achieved until about
26,000 '4C yr BP. The last remnants of Subantarctic
Evergreen Forest disappeared at Taiquemo between 32,105 14Cyr BP and 26,019 14Cyr BP, and
at Dalcahue about 25,176 14Cyr BP (Heusser et al.
1999).
The anatomy of the youngest piedmont glacier
advances is depicted in Fig. 2 and in Table 1 from
details in Denton et al. (1999). The first of these is
dated to 22,460 14C yr BP for the Lago Rupanco
lobe; between 20,890 14Cyr BPand 23,020 14Cyr
BP for the Lago Llanquihue lobe; 22,570 14Cyr BP
for the Seno Reloncavi lobe; and 22,295 14Cyr BP
for the northern part of the Golfo Corcovado lobe.
Basal dates from mires on the moraine belts near
Lago Llanquihue and Seno Reloncavi indicate recession after the culmination of this advance (Denton et al. 1999). Note that new radiocarbon dates
reported in Denton etal. (1999) have caused the advance of most lobes to be placed at 22,295-22,570
14Cyr BP, ratherthan at the somewhat younger ages
reported in Lowell et al. (1995). See Denton et al.
(1999) for the current inventory of radiocarbon
dates associated with this advance, as well as alternative interpretationsof available dates. There may
also have been advances to the edge of the kame
terraces alongside the western shore of Lago Llanquihue shortly before 17,800 '4C yr BP and again
shortly before 15,730 14Cyr BP.
The youngest readvance of the LGM is radiocarbon dated near Lago Llanquihue to 14,650 14C
yr BP (Puerto Phillippi site), 14,869 14Cyr BP(Llan116
quihue site), 14,882 14C yr BP (Northwest Bluff
site), 14,540 14C yr BP (Bella Vista Bluff site),
14,550-14,613 14Cyr BP (railroadbridge site), and
14,820 14Cyr BP (Calle Santa Rosa site) (Denton et
al. 1999). Near Seno Reloncavf it is dated to 15,220
14Cyr BP at Isla Maillen, 14,879 14Cyr BP at Punta
Penas, and shortly after 15,040 14Cyr BP at the top
of the ice-contact slope beside Canal Tenglo. For
the Golfo de Ancud lobe, this youngest readvance
is dated to shortly after 14,900 14Cyr BP at the Calbuco site. For the northern part of the Golfo Corcovado lobe, numerous radiocarbondates from the
Dalcahue site place this readvance at close to
14,805 14Cyr BP. Therefore, we conclude that these
four piedmont glacier lobes fluctuated in near synchrony (within the limits of radiocarbon dating)
during the youngest major pulse of the LGM. Note
that because of numerous new radiocarbon dates
reported in Denton et al. (1999), the date for the
thin upper peat bed at the railroadbridge site is older than reported in Lowell et al. (1995). Furthermore, the ages of the lower, thick organic bed and
underlying organic silt are also older than reported
previously in Mercer (1976) and Hoganson and
Ashworth (1992).
Two older advances into the outer Llanquihue
moraine system during the LGM occurred at about
29,400 14Cyr BP and at 26,760 14Cyr BP (Denton
et al. 1999) (Fig. 2 and Table 1). The first of these
older advances is recorded for only two piedmont
lobes (Llanquihue and Seno Reloncavf), and the
second for only one lobe (Llanquihue). Nevertheless, we consider it likely thatboth advances are regional, given the nearly simultaneous behavior of
several lobes during the youngest two maxima of
the LGM.
A fifth major point in Fig. 2 involves the structure of the climate changes that, taken together,
constitute the last glacial-interglacial transition.
An examination of this structurerequires a detailed
chronology of the youngest glacial maximum. As
described above, the advance to this maximum culminated at 14,550-14,805 '4C yr BP for piedmont
glacier lobes in the field area. For the northernpart
of the field area, this youngest advance of the LGM
terminated behind the outermost late Llanquihue
moraine belt. However, in the southern portion of
the field area, this final advance was the maximum
of late Llanquihue time. We interpret the lack of
younger radiocarbon dates associated with sediments deposited at this maximum to mean that ice
recession was underway immediately.
Pollen analyses of three organic beds deposited
Geografiska Annaler * 81 A (1999) · 2
LINKAGEOF PALEOCLIMATE
INTERHEMISPHERIC
DURINGTHELASTGLACIATION
just before the culmination of this youngest advance at the Llanquihue, Bella Vista Bluff, and
Punta Penas sites all show cold, wet Subantarctic
Parkland conditions (Heusser et al. 1999). For example, the pollen analysis of the organic silt bed at
the Bella Vista Bluff site in Puerto Varas, which
covers the interval from 14,540 to 15,640 '4C yr BP,
indicates continuous cold Subantarctic Parkland
conditions right up until 14,540 14Cyr BP, when organic accumulation ceased as a result of the lakelevel rise that heralded readvance of the Lago Llanquihue piedmont lobe. This is consistent with other
detailed pollen evidence for cold moorland habitats in the southern Chilean Lake District throughout the latter part of the LGM from the Canal de la
Puntilla (Moreno 1997; Moreno et al. 1999) and
Fundo Llanquihue (Heusser et al. 1999) sites. The
open environment and increased precipitation implied by the moorland pollen taxa suggest a northward shift of the westerlies storm tracks (Moreno
et al. 1999).
These paleoenvironmental inferences from pollen analysis are consistent with earlier conclusions
drawn from fossil beetle data (Hoganson and Ashworth 1992; Ashworth and Hoganson 1993). One
of the key fossil-beetle sites is an organic silt layer
dated to 15,715 440 14Cyr BP (GX-5275) within
the kame terrace at the Bella Vista Park site (Hoganson and Ashworth 1992), situated about 175 m
north of the Bella Vista Bluff site. The age and
stratigraphicposition indicate that this organic bed
at the park site corresponds with the organic silt
bed at the bluff site. Another key sample locality is
the Puerto Varas railroad bridge site, where a silt
and sand unit is capped by a 26-cm-thick layer of
organic silt, peat, and wood (see Denton et al.
1999). Mercer (1976) reported a date of 16,270±
360 14C yr BP (RL-1 13) of wood from near the base
of the silt and sand unit. Hoganson and Ashworth
(1992) gave an age of 14,060±450 14C yr BP (GX5507) for the capping organic layer. Ten samples of
fossil beetles come from regular intervals within
the silt and sand unit, and two samples from within
the capping organic layer (Hoganson and Ashworth 1992). The implication based on these two
dates is that the beetle samples come from regular
intervals between 16,270 14Cyr BP and 14,060 14C
yr BP, the critical last several thousand years of the
LGM (Hoganson and Ashworth 1992, table 1 and
fig. 4). However, our new dates show that the entire
silt and sand unit was deposited at 17,350-17,880
14C yr BP (Denton et al. 1999), and hence that the
ten beetle samples from the silt and sand unit refer
Geografiska Annaler · 81 A (1999) · 2
only to that time. In addition, wood from the upper
few centimeters of the capping organic layer yielded a mean age of 14,613 '4C yr BP (Denton et al.
1999). In view of the fact that the capping organic
bed is developed into the top of the silt unit, it is not
clear whether the two beetle samples from the organic bed refer to 14,613 '4C yr BP or to 17,35017,880 '4C yr BP. With this revised chronology, the
fossil-beetle data are invaluable in showing cold,
wet environmental conditions near Lago Llanquihue at 16,000-18,170 14Cyr BP (Canal de Chanchan site near Puerto Octay), 17,350-17,880 14Cyr
BP (railroadbridge site in Puerto Varas), and 15,715
14Cyr BP (Bella Vista Park site in Puerto Varas); all
of these times occur within the long interval of the
LGM.
We earlier presented evidence for significant
amelioration of climate at the end of the last glaciation. That evidence was the invasion of the lowlands of the southern Chilean Lake District by thermophilic arboreal elements of the North Patagonian Evergreen Forest. This arboreal diversification
began with the spread of Myrtaceae, Nothofagus
cf. dombeyi, Lomatia, Maytenus and other relatively thermophilic arborealtaxa, which became prominent by 13,900 14Cyr BP at the Canal de la Puntilla
site, by 13,500 14Cyr BP at the Fundo Llanquihue
site, and by 13,700 '4C yr BP at the Alerce site
(Lowell et al. 1995). To determine more closely the
beginning of the glacial/interglacial transition, we
subsequently investigated in detail several cores in
a transect at the Canal de la Puntilla site (Moreno
et al. 1999) and a single core with a high sedimentation rate at the Huelmo site (Moreno 1998). In
both cases the chronology was established with numerous new AMS radiocarbon dates. This detailed
examination shows that the first indication of
warming was a rapid rise of Nothofagus at 14,600
14Cyr BP at both sites. Warming continued with the
invasion of thermophilic tree species of the North
Patagonian Evergreen Forest about 14,100 '4C yr
BP at Canal de la Puntilla and 14,200 '4C yr BP at
Huelmo. These vegetation changes are taken to be
regional because they are recorded nearly simultaneously at two widely separated sites. The chronologies show that both vegetational changes are rapid
at each site.
Thanks to these detailed and well-dated stratigraphies, we can place age brackets on the climate
change that marked the end of glacial conditions in
the southern Chilean Lake District. The pollen
record from the Bella Vista Bluff section in the terrace at Puerto Varas shows cold conditions until the
117
G.H. DENTON ET AL.
culmination of the last glacial advance of the LGM
at 14,550-14,805 14Cyr BP (Heusser et al. 1999).
The first indication of subsequent climate amelioration is the rise of Nothofagus at 14,600 14Cyr BP
at Canal de la Puntilla and at Huelmo (Moreno
1998; Moreno etal. 1999), followed by invasion of
thermophilic species at about 14,000 14Cyr BP at
many sites in the Lake District (Heusser et al.
1999). Note that because of extensive new radiocarbon dating described in Denton et al. (1999),
this age for the beginning of the last termination is
slightly earlier than we reported in Lowell et al.
(1995).
The pollen records from throughout the southern Chilean Lake District and on Isla Grande de
Chiloe register a consistent late-glacial pattern of
vegetational development (Heusser et al. 1996,
1999; Moreno 1997, 1998; Moreno et al. 1999).
The inferences drawn from the detailed pollen
records from the Fundo Llanquihue and Canal de
la Puntilla sites are supported by a network of cores
analysed with less temporal resolution. The interval between 14,000 and 13,000 14Cyr BP is a transition zone in which thermophilic arboreal elements of the North Patagonian Evergreen Forest
invaded the lowlands, but in such relatively low
numbers that grasslands and elements of Magellanic Moorland remained. The persistence of
moorland taxa suggests that the westerlies storm
belt remained in the inferred northern glacial position. A closed-canopy North Patagonian Evergreen
Forest was then established at 12,700-13,000 14C
yr BP. The rapidity and magnitude of this event at
12,700-13,000 14C yr BP are particularly well
shown in the influx diagram from Canal de la Puntilla (Moreno et al. 1999), but are also evident in
other diagrams. One explanation for the disappearance of moorland taxa is the postulated southward
shift of the westerlies storm track. The evergreen
forest reached its fullest development about
12,000-12,200 '4C yr BP, and may at that time have
contained some elements of the Valdivian Evergreen Forest.After 12,000 '4C yr BP, opening of the
forest and the appearance of cold-tolerant elements
of the North Patagonian Evergreen Forest (Podocarpus nubigena, Pseudopanax laetevirens) suggest climate cooling that continued until at least
10,500 14Cyr BP, when the Fundo Llanquihue and
many other vegetational records are disturbed by
the influence of fire. However, farthersouth on Isla
Grande de Chiloe, the Taiquem6 record escaped
the influence of fire, and here North Patagonian Evergreen Forest communities opened to allow the
118
significant presence of Podocarpus and Pilgerodendron-type during episodic cooling between
11,360 and 10,355 14Cyr BP(Heusser et al. 1999).
We mentioned above that the overall deterioration of climate at the coldest times of the LGM was
about 6-8°C in mean summer temperaturerelative
to the present-day value (Heusser et al. 1999;
Moreno et al. 1999). About 3°C of this difference
was recovered in the first warming pulses between
14,600 '4Cyr BPand 12,700-13,000 14CyrBP.Another 3°C was recovered between 12,700-13,000
14Cyr BP and 12,200-12,500 '4C yr BP. Our data
suggest that warming at the beginning of these intervals was rapid. Finally, we infer that the late-glacial climate reversal after 12,000-12,200 '4C yr BP
involved a relatively small decline in temperature
(probably < 2-3°C, Heusser et al. 1999), compared
to the high values achieved at 12,200-12,500 14C
yr BP.
New Zealand paleoclimaterecord
To test the Chilean paleoclimate record as a barometer of change in the zone of Southern Hemisphere westerlies, rather than simply as a piece of
a complex matrix of varying regional climate
changes, we examined the available records and we
obtained additional radiocarbon dates from New
Zealand at the same Southern Hemisphere middle
latitudes but on the opposite side of the Pacific
Ocean. Although this test is preliminary, we conclude that major climate shifts in New Zealand are
similar to those in the Chilean Andes. The basis for
this conclusion follows.
Grassland and shrubland were dominant during
the LGM in most of New Zealand south of about
37°S (the latitude of Auckland in northern North
Island) (McGlone 1995), with continuous forest
persisting only in the extreme north (Fig. 3). Only
small patches of Nothofagus, Libocedrus, and podocarp trees remained in South Island. Forest
stands may have persisted in hilly areas of North Island south of 37°S, but grassland and shrubland
dominated the lower, rolling terrain. Such extensive deforestation probably reflects severe frosts
from outbreaks of cold subantarcticair masses over
most of New Zealand during the LGM (McGlone
1988). The oceanic reconstruction from Deep Sea
Drilling Project (DSDP) Site 594 on the south flank
of Chatham Rise 300 km east of South Island is
consistent with severe terrestrial conditions, because it shows cold subantarctic water close to
South Island during the LGM (Nelson et al. 1993).
Geografiska Annaler
- 81 A (1999) 2
·
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
170°
168°
I
172°
I
I
174°
176°
178°
I
I
I
Explanation
Ice today
[-j
I
Ice at last glacial maximum
|Alpine vegetation
.r
Grassland - shrubland
t
I! Tall, lowland - montane
lowlands
coniferous broadleaf forest
Fig. 3. The extent of mountain glaciers in the Southern Alps is from
Hollin and Schilling (1981) and the
distribution of vegetation is from
McGlone et al. (1993), both during
the LGM at approximately 18,000
14C yr BP. Also listed are the key
stratigraphic sites mentioned in the
text. The radiocarbondates associated with these sites and not listed in
the text or in Denton and Hendy
(1994) are as follows. At the Kamaka site, an organic-rich bed is overlain by laminated lacustrine deposits
and then by coarse outwash. The
laminated deposits reflect glacier expansion. They have yielded the following succession of AMS radiocarbon dates of small enclosed wood
and plant fragments, from bottom to
top: 22,420±230 14C yr BP (AA22271), 22,460±240 14Cyr BP (AA22275), 22,360±190 14C yr BP
(ETH-16387),22,620+200 '4CyrBP
(ETH-16388),22,160±310 4C yrBP
(AA-22275), 22,490±330 14Cyr BP
(AA-22279), 22,100+790 14Cyr BP
(AA-22282), 22,570±200 14Cyr BP
(AA-22283), 22,860+200 14Cyr BP
(AA-22285), 22,730±200 14Cyr BP
(AA-22286), 22,680+220 14Cyr BP
(AA-22290), 22,650+240 14Cyr BP
(AA-22291), 21,990±220 14Cyr BP
(ETH-16390), and 22,380+190 14C
yr BP (ETH-16391.
+
|Scattered forest areas
| t|7 Present-day coastline
Coastline at last glacial maximum
I Radiocarbon or exposure
sample site
Pollen site
e
Raupo
Kamaka
Kumara--"
Abut Head
Omoeroa Bluff ~--
-Pukaki
o DSDP Site 594
9
0
Scale
100
200
300
km
°..
192 4s°
17n-
172°
174°
176°
178°
Drafted by Richard D. Kelly Jr., Augusta, Maine, 1997.
In anycase, the harshterrestrialclimateduringthe
LGMpromotedinstabilityof the deforestedlandscape, with widespreadmechanical weathering,
slopeerosion,andloess deposition;extensivealluvial plains developedbetween the SouthernAlps
andthe east coast of SouthIsland.
Figure 4 displays the glacier advances in the
SouthernAlps thatoccurredwithinthis LGMenvironment.Duringthe maximumof the lateOtiran
glaciation(LGM),equilibrium-linealtitudesnear
Lake Pukaki were depressed 875 ±50 m below
present-dayvalues during deposition of the Mt
John(Kumara-22)
morainesand750 + 50 m during
Geografiska Annaler * 81 A (1999) · 2
River
deposition of the Tekapo(Kumara-33)moraines
(Porter 1975). Radiocarbondates from several
sites indicate multiple advances into late Otiran
morainebeltsthataretakento be theequivalentsof
the outerof the two late Otiranmorainebelts near
LakePukaki(Suggate1978).An organicbedoverlain by till occurswithinthe outerOtiranmoraine
belt at Mt Hercules on the western flank of the
SouthernAlps. Three small wood samples from
within this bed yielded new ages of 23,870 ± 330
14C yr BP (A-6188), 23,560 370 14C yr BP (A6591), and 23,510 350 '4C yr BP (A-6592). The
glacial advancerepresentedby the overlyingtill
119
G.H. DENTON ET AL.
Fig. 4. Paleoclimate records for New Zealand. The left panel shows the radiocarbon-dated glacial deposits indicative of ice advances
into the outer Otiranmoraine belts of the LGM. These are plotted with regardto the associated drop in equilibrium line altitude associated
with correlative moraine belts near Lake Pukaki (Porter 1975). The existence and chronology of the Younger-Dryas-age glacier readvance is from Denton and Hendy (1994), Lowell et al. (1995), and Ivy-Ochs et al. (1999). The right three panels show the record for
core DSDP Site 594 (Nelson et al. 1993, Heusser and van der Geer 1994). The marine pollen reflects changing vegetation in New Zealand. The CaCO3 record is a reflection of the amount of silt input from New Zealand. The benthonic 6180 signal is largely a reflection
of Northern Hemisphere ice-sheet volume. On the right is shown the numbers of marine isotope stages from Nelson et al. (1993).
likely occurred shortly after this time. Farthernorth
in the Grey River Valley, a stratigraphic section at
Kamaka registers an advance of the northern of the
two major piedmont lobes of the TaramakuGlacier
system at the LGM. Here an organic silt bed lies beneath laminated glacial lacustrine sediments and
outwash that herald advance of this lobe to the outer moraine system (Suggate 1965). New AMS radiocarbon dates of 14 small wood pieces from the
lacustrine sediments yielded a mean value of
22,400 14Cyr BP, consistent with an earlier date of
22,300+350 14Cyr BP (NZ-116) from the organic
silt bed (Suggate 1965).
A prominent maximum of the western of the two
piedmont lobes of the former Taramaku Glacier
system occurred near Kumara at close to 17,700
14Cyr BP. This maximum was the greatest of late
Otiran time for this lobe. The age comes from five
new AMS radiocarbon dates of small samples of
grass, wood, bark, and beetles from near the top of
an organic silt bed that underlies till with a non-erosive conformable contact. These dates are 17,380+
130 14CyrBP(ETH-13398),17,720±120 '4C yrBP
120
(ETH-11104), 17,900±140 14C yr BP (ETH13399), 18,160±140 14Cyr BP(ETH-13400), and
18,360 + 140 14C yr BP(ETH-13402). Two other radiocarbon samples previously collected within the
organic silt bed yielded ages of 18,450 ±300 14C yr
BP (NZ-4408) and 17,250+250 14C yr BP (NZ4407) (Moar 1980).
The terrace at Raupo (Kumara-22terraceof Suggate and Moar 1970) in the Grey Valley northeast
of Kamaka reveals an interstadial organic silt bed
with an aggregate thickness of 60 cm that separates
two outwash units. Previous radiocarbon ages of
samples from within this bed are 18,600+290 14C
yr BP (NZ-891) and 18,750+180 14Cyr BP (NZ737) (Suggate and Moar 1970). Additional radiocarbon samples that we collected yielded ages of
19,740± 150 14C yr BP (A-6550) for the base of the
organic silt bed, 18,940+170 14Cyr BP (A-6551)
for the middle, and 18,780 140 14CyrBP(A-6552)
for the top. Renewed outwash deposition in the
Grey Valley after 18,780 '4C yr BP probably represents glacier advance into the outermost late Otiran
moraine system. Also, glacier ice must have exGeografiska Annaler - 81 A (1999) · 2
DURINGTHELASTGLACIATION
LINKAGEOF PALEOCLIMATE
INTERHEMISPHERIC
tended into this moraine system when outwash was
deposited below the organic silt bed.
Near Abut Head on the Tasman seacoast, an advance into the outer moraine belt resulted in the
deposition of a large moraine on a thin organic bed
now exposed in a sea cliff. Radiocarbon samples
from the undisturbed uppermost part of this bed
give the following new ages for this advance:
16,615+95 14CyrsBP(A-9063), 16,575 +8 14CyrBP
(A-9064), 16,920±100 14C yr BP (A-9065), and
16,525 ±90 14Cyr BP (A-9066).
The advance to the widespread Kumara-33 moraines (or their equivalents), nested just behind the
outer Otiranmoraine belt, is dated only at Omoeroa
Bluff, cut into a moraine ridge on the Tasman Sea
coast near Franz Josef Glacier. Here two radiocarbon samples of a thin peat bed buried by drift of this
youngest advance yielded results of 13,950±140
4C yr BP (NZ-479) (Wardle 1978) and 15,300±
120 14C yr BP (Suggate 1990). Behind these moraines (or their equivalents) in most valleys lie basins, now or formerly filled with lake or sea water,
which reflect rapid collapse of ice tongues. Ice
cleared these basins prior to 13,500 14C yr BP in the
Tasman Valley near Lake Pukaki and 13,400 14Cyr
BP in the Paringa River valley (Suggate 1968).
From this preliminary chronology, Suggate (1965)
and Suggate and Moar (1970) estimated that rapid
ice recession close to 14,000 14C yr BP marked the
end of the LGM. As discussed below, however, palynological data from the Waikato lowlands in
North Island place the first decisive warming of the
last glacial/interglacial transition a bit earlier, at
shortly after 14,700 14C yr BP, based on the radiocarbon ages of numerous distinctive volcanic tephras registered in pollen records (Newnham et al.
1989).
Late-glacial readvances are recorded in many
valleys of the Southern Alps. Readvance of Franz
Josef Glacier toward the Waiho Loop moraine occurred at 11,050± 14 14C yr BP (Denton and Hendy
1994, 1995). A late-glacial moraine remnant in upper Cropp River Valley was dated to 10,250 14Cyr
BP by Basher and McSaveney (1989). We have obtained an error-weighted mean age of 10,055+±29
14Cyr BPfor five additional wood samples from this
remnant (Lowell et al. 1995). Because Mabin
(1995) cautioned against accepting the concept of
a regional Younger-Dryas-age readvance on this
limited evidence, exposure dates were obtained on
boulders from a prominent late-glacial moraine in
Arthur'sPass at the head of a tributaryof the former
TaramakuGlacier system. The dating results sugGeografiska Annaler * 81 A (1999) · 2
gest that the moraine is Younger Dryas in age (IvyOchs et al. 1999). This moraine represents snowline lowering of about 260-360 m relative to the
Little Ice Age value (Ivy-Ochs et al. 1999).
Figure 4 also depicts the pollen, carbonate, and
6180 records from marine sediment cores from
DSDP Site 594 on the south flank of Chatham Rise
300 km east of South Island (Nelson et al. 1993;
Heusser and van der Geer 1994). These records are
linked to the changing paleoenvironment of South
Island, because DSDP Site 594 is situated offshore
and downwind from the Southern Alps, from alluvial fans of the CanterburyPlains, and from major
rivers that drain the Southern Alps. The pollen
record shows the changing percentages of trees and
shrubs as reflected in pollen carried offshore by the
prevailing westerly winds and by rivers draining
the east coast of South Island. The carbonate record
is taken to represent increases in silt input (and a
corresponding dilution of CaCO3) during cold intervals from terrestrialerosion and outwash aggradation, with the resulting generation of loess and
offshore sediment transport. The pollen and carbonate records are closely correlated with each other, presumably because the extent of forest cover
went hand in hand with landscape stability and
hence sediment production on South Island. Both
of these terrestrial indicators seem to correspond
with the benthonic 6180 record. In turn, the benthonic 6180 record is determined by changes in
both temperatureand ice volume, with the latterbeing dominant. Because the Antarctic Ice Sheet
showed only modest volume change through the
last glacial cycle, the benthonic 6180 signal largely
follows Northern Hemisphere ice-volume changes.
The beginning of reforestation at the end of the
LGM is documented in North Island in the Waikato
lowlands. Grass and shrubs dominated these lowlands at the LGM (McGlone et al. 1993). Reforestation of the Waikato lowlands by a podocarp-hardwood community was rapid after deposition of the
Rerewhakaaitu tephra at 14,700 14Cyr BP (Newnham et al. 1989). Podocarp-hardwood forest had
spread across much of North Island by about
14,000 14Cyr BP, followed by progressive reforestation of the rest of New Zealand (McGlone et al.
1993). It is importantto note that the vegetation just
prior to 14,700 14Cyr BP was essentially the same
as that earlier in the LGM at 18,000 14Cyr BP depicted in Fig. 3 (McGlone etal. 1993). The massive
changes in the forest cover of North Island at the
end of the LGM, shortly after 14,700 14C yr BP,
were synchronous with the rapid rise of Nothofa121
G.H. DENTON ET AL.
gus, followed by the invasion of thermophilic trees
into the Chilean Lake District, at the end of the
LGM.
Landscape stabilization accompanied reforestation in central North Island, as documented by the
relative preservation of highly erodible tephras.
The landscape was unstable throughout the LGM
at elevations above 300 m (Leamy et al. 1973; McGlone et al. 1984; Kennedy 1988). The Rerewhakaaitu tephra, dated to 14,700 14Cyr BP (Froggatt
and Lowe 1990), is the earliest to be preserved on
highly eroded LGM landscapes above 300 m. By
13,080 14C yr BP (Rotorua Tephra; Froggatt and
Lowe 1990), extensive areas above 500 m were stable. By 11,850 14C yr BP (Waiohau Tephra; Froggatt and Lowe 1990), landscapes were stable up to
900 m elevation (McGlone 1995).
Middle-latitudeSouthern Hemisphere data in
a global context
In order to place the middle-latitude Southern
Hemisphere data from Chile and New Zealand into
a global perspective, we first present background
about the asymmetric shape of the 100,000-yr glacial cycles of late Quaternarytime, as our data bear
on the mechanisms that terminatedthe last such cycle. We also review critical elements of the classic
North Atlantic/European paleoclimate record because it is so important for our interhemispheric
comparison. We then go on to address the main implications of our data in the context of this background. We conclude by posing a series of questions raised by these implications, and suggest an
initiative to answer these questions.
Background
Late Quaternary 100,000-yr glacial cycles. Explaining the asymmetric 100,000-yr glacial cycles
of late Quaternarytime is one of the most difficult
challenges of paleoclimate research (Imbrie et al.
1993). The transition toward these 100,000-yr cycles began about 950,000 years ago, but took more
than 300,000 years to complete (Imbrie et al.
1993). By 600,000-650,000 years ago a clear
asymmetric 100,000-yr cycle had emerged. The
subsequent glacial extremes of MIS 16, 12, 10, 6,
and 2 reflect the storage of excess ice in huge
Northern Hemisphere ice sheets just prior to terminations that abruptly ended the buildup phases of
these asymmetric cycles (Raymo 1997).
The overall asymmetric shape of the last
122
100,000-yr ice-volume cycle, depicted in Fig. 5,
features a prolonged growth phase terminated by a
much shorter collapse phase. The existence of such
a prolonged growth phase implies to us that the icesheet system (and probably the entire climate system) became largely detached from pervasive orbital forcing early in the glacial cycle (except for
the long-term oscillations superimposed on the
prolonged growth phase) (see also Imbrie et al.
1993). Also superimposed on the long buildup
phase, at least in the North Atlantic region, are
shorter asymmetric oscillations known as Bond cycles (Bond et al. 1993; Broecker 1994; Rasmussen
et al. 1997). Each such cycle represents a period of
climate cooling that ended with a massive Heinrich
discharge of icebergs (Table 1), in turnfollowed by
abrupt increase of North Atlantic sea-surface temperatures. Each of these abrupt North Atlantic
warmings failed to terminate the long growth phase
of the 100,000-yr cycle. Instead, climate reverted
to cold conditions and ice-sheet buildup continued
until the decisive Heinrich 1 ice collapse (Table 1),
which was coincident with the beginning of the irreversible decline in ice-sheet volume registered in
TR163-31B (Fig. 5).
A striking feature of the late Quaternarymarine
6180 record is that the terminations of glacial
100,000-yr cycles had similar magnitudes during
times of both high-amplitude and low-amplitude
insolation changes. For example, increases both in
northern summer insolation and in tropical equinox insolation were considerably smaller during
the last termination than during the penultimate
termination, and yet both terminations had the
same amplitude. Because of the effect of long eccentricity cycles on the amplitude of precession,
the same situation occurred about 400,000 years
ago during TerminationV between MIS 12 and 11.
This is consistent with the postulate that terminations represent ocean-atmosphere reorganizations
between preferred glacial and interglacial modes
of the climate system (Broecker and Denton 1990).
And yet terminations recur at approximately the
100,000-yr intervals set by eccentricity (which
controls the amplitude of precession). The implication is that during terminations the increased effect of insolation somehow triggers a fundamental
reorganization of a non-linear system but does not
control the magnitude of the reorganization.
Asymmetric 100,000-yr climate cycles of late
Quaternarytime vary in length. In fact, the time intervals between terminations range from 84,000 to
120,000 years (Raymo 1997). Moreover, terminaGeografiska Annaler .81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
Core TR163-31B (3°37'S; 83058'W)
610 (%o)
5.5
6.0
V
I
lT
r
t
t ?
T
5.0
I
4.5
1-
t
4.0
I
t
t
t
V-19-30 (3023'S; 82°21 'W)
Sea Level (m below present)
3.5
o
t
3.0
150
100
t ,
I
50
I f
0
.
o0
·
*
. ** 00
*
o
.C^S
B
oi°
**o to
*.*. *^S
o
%-12,860t250
c-
15,080:200
16,120t260
*-
o°
*
6
o
<
17,200±220
*
*
0
*-
**
~
-
19,580+300
<20,960±370
.--20,890O360
*
9
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14,020:1180 -
14,340t220
<
o
21,480±280
-21,450t250
"I
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o.
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----23,930O320
25,690±460
N4
* P. wuellerstorfi
* Uvigerina
Orfd
by R. D. Kety Jr.
a - ----28,690600
% a
917
Fig . 5. The right panel, adapted from Shackleton (1987), shows a sea-level estimate based on the 5 'O records of V 19-30 in the equatorial
eastern Pacific Ocean (3°23'S, 83°21'W) and RC-17-177 in the equatorial western Pacific Ocean (1°45'N, 159°27'E). To construct this
curve, a correction for changing deep-water temperatures was made by calculating the isotopic difference between V- 19-30 (benthonic)
and RC 17-77 (planktonic). This difference was then subtracted from the benthonic V- 19-30 record. The results were scaled to the New
Guinea sea-level data to yield an estimate from isotopic data of the global sea-level record. Even though it is certainly not a perfect
representation of sea level, the resulting curve nevertheless highlights the long, gradual decline of sea level (increase of continental ice
volume), followed by an abrupttermination, that marked the latest 100,000-yr glacial cycle. The left panel shows the timing of the last
termination in detail. The figure displays the isotope data set from equatorial eastern Pacific core TR163-31B (3°37'S, 83°58'W) from
Shackleton et al. (1988), with original isotope data points kindly provided by NJ. Shackleton. The radiocarbon dates have a marine
reservoir correction of 580 14Cyears. The results from two species of benthonic foraminifers are plotted. They are taken to be a reasonable reflection of changes in continental ice volume. They show that the fundamental change between rising and declining isotope
values, which marks the beginning of the last termination, occurred close to 14,500 14Cyr BP (not adjusted for unknown mixing time
of the ocean at that date).
tions that lead to full deglacial conditions similar to
those of the present day seem to occur only after accumulation of considerable excess ice in huge
Northern Hemisphere ice sheets (Raymo 1997). In
other words, the major prerequisite for a sharp,
complete termination seems not to be excessive orbital forcing, but rather excessive ice volume.
North Atlantic/European late-glacial climate.
Mangerud et al. (1974) defined a set of chrons for
Geografiska Annaler
.81 A (1999) - 2
the subdivision of late-glacial time. The Boiling
chron was placed at 12,000-13,000 14Cyr BP, the
Older Dryas chron at 11,800-12,000 14C yr BP, the
Aller6d chron at 11,000-11,800 14Cyr BP, and the
Younger Dryas chron at 10,000-11,000 14Cyr. BP.
One of our major points given below is that the climate trigger for the last termination occurred during the millennia before the B6iling chron. Therefore, for the purposes of this paper, we use the term
Oldest Dryas chron to encompass the interval be123
G.H. DENTON ET AL.
a.
Southern Alps, New Zealand
(42°-44°30'S; 169°- 172°W)
MountainGlaciers
ELADepression (m)
1000
, .
,
,
,
500.
I
.
.
.
b.
c.
Eastern NorthAtlanticOcean
SU81-18
(37°46'N;10°11'W)
Chile
Southern LakeDistrict
(40030'-42°25'S; 72o25'-73045'W)
MeanSummerTemperature
6180 (%o)
8°C 10°C 12°C 14°C 16°C
0I
I
I
I
i
2
I
1
0
c
C
-
Preboreal
0 2
___0
,--- Z"
=z-:
.-11,e1n
-
_S.-12-W
CL
32,73313--1^t3.t
.
13,40t10.
,n i
i
__a
a
O,
T Z:.
0,-,
'-
1-
2
uYounger
Dryas
1i23,3e1701 OlderDry
B11lingOldest Dryas
i
o
n
0
l
i ..i..
0
-
< --
. I . .
. i
t N
o
,)
Drafted by R. D. Kelly Jr. 1997
d.
e.
f.
Switzerland
MountainGlaciers
Gerzensee
06°(46°- 48°N;
10°E) (46°50'N;07°33'E)
ELA Depression (m)
1000
8180 (%o)
0 -10
500
-8
-6
Q.
_Prboreal
Younger Dry_
X-
Aller8d
I
tW Older
Dryas' _rc
--B1111ng -U'~u_
(U
"
-
Oldest
Dryas
.
I. .
-- .
.
-
I.
-
I
-....
,
<-Glacier advance-
Fig. 6. Comparison of terrestrialrecords from middle latitudes of the Southern Hemisphere with paleoclimate records from the North
Atlantic region. Panels (a) and (b) are from Figs 2 and 4. Panel (c) is adapted from Bard etal. (1987). It shows the oxygen-isotope values
and AMS radiocarbon ages for the planktonic species Globigerina bulloides in core SU8 1-18 in the eastern North Atlantic Ocean off
Portugal. The AMS radiocarbon dates have a standardmarine reservoir correction of 400 14C years. The curve reflects a combination
of ice volume and temperaturechanges. It shows the same structureof paleoclimate signals as the records from middle latitudes of the
Southern Hemisphere in panels (a), (b), and (d). Particularly important is that the last termination, defined on the basis of the isotope
record, begins at 14,500 14C yr BP (Bard et al. 1987). This is the same result found in the detailed isotope record from core TR163-3 1B
in Fig. 7 from Shackleton et al. (1988). It correlates with the initial abrupt warming in the southern Chilean Lake District (panel (b))
and in New Zealand (Newnham et al. 1989). Panel (d) is from Ariztegui et al. (1997) and represents a curve of glacier activity in the
Argentine Andes that has similarities with the Greenland stable isotope record in panel (e) from Johnsen et al. (1992). Panel (f) shows
the Gerzensee lacustrine 6180 record from the forelands of the Swiss Alps (Eicher and Siegenthaler 1976), which likewise exhibits a
correspondence with the Greenland record. Also shown in panel (/) are the approximate equilibrium-line elevations on mountain glaciers
during the Zurich Stade just before massive recession began about 14,600 14C yr BP (Schliichter 1988; Ivy-Ochs et al. 1996), for the
time of maximum Bolling warmth (Maisch 1982), and for Egesen time (Maisch 1995; Ivy-Ochs et al. 1996). The European chronozones
shown on panels (c) and (f) are from Mangerud et al. (1974) as modified in the text. An important point here is that the Swiss Alps experienced massive deglaciation during Oldest Dryas time correlative with deglaciation in the southern Chilean Lake District (panel (b)).
124
Geografiska Annaler · 81 A (1999) · 2
DURINGTHELASTGLACIATION
LINKAGEOF PALEOCLIMATE
INTERHEMISPHERIC
tween 13,000 and 15,000 14C yr BP, much in the
sense of Welten (1982). This means that we place
the boundary of the Upper Pleniglacial and the Late
Glacial subages of Hammen et al. (1967) at 15,000
14Cyr BP, not at 13,000 14Cyr BP. The Late Glacial
subage would then include the Oldest Dryas, B6lling, Older Dryas, Aller6d, and Younger Dryas
chrons. However, when we discuss the climate signature of central Europe, we use the system of pollen zonation of Firbas (1949, 1954), which Lotter
et al. (1992) considered as strict biozones.
The basic shape for the North Atlantic/European
late-glacial climate signal is manifested in British
Isles beetle remains (Atkinson et al. 1987), along
with 6180 switches in both Greenland ice and Swiss
lacustrine marl (Dansgaard et al. 1984) (Fig. 6).
This shape consists of an abrupt change near or
shortly after the Oldest Dryas/B611ing transition
from cold conditions to an interval 200 years in
length that was close to interglacial warmth. This
warm interval was followed by progressive cooling,
with oscillations, throughlate Boiling, Older Dryas,
and Allerod time, culminating in the Younger Dryas
cold episode (Oeschger 1991). There is remarkable
coherence to this regional climate signal, such that
even the superimposed short-lived Gerzensee, Older Dryas, and early Preboreal 6180 oscillations can
be traced easily across this portion of the globe. The
only major discrepancy is that, whereas the 6180
and beetle records show an early B611ing-warmpeak
followed by a cooling trend, most pollen-recordsexhibit continuous vegetation development from
herbs and shrubs to a succession of forest trees. The
only widespread reversal in the pollen record took
place in the Younger Dryas and to a much less degree in the Older Dryas (Watts 1980).
The discrepancy between the magnitude of early
Bolling warmth in Europe implied by vegetation as
opposed to isotope and insect records may be more
apparent than real. An example comes from Lobsigensee on the Swiss Plateau (Ammann and Lotter
1989; Ammann 1989a, b; Elias and Wilkinson
1983, 1985). The Oldest Dryas part of the Lobsigensee records features a boreal and boreal-montane insect assemblage, along with Betula nana.
During the early B611ing warming, the boreal insect assemblage is replaced by a plant-independent
temperate assemblage that reflects a mean July
temperature close to interglacial values, just as in
Great Britain. A shift in water plants supports such
marked warming. The early Bolling climate was
warm enough to support broad-leafed deciduous
trees such as oak or hazel, which do not appearuntil
Geografiska Annaler · 81 A (1999) · 2
3000 years later because of migrational lags. It is
interesting to note that the 6180 record shows lateglacial climatic deterioration beginning in latest
Boiling time and culminating in a Younger Dryas
reversal. The vegetation record shows only a small
increase in non-arboreal pollen in Younger Dryas
time, reflecting some openings in the forest cover.
The alpine moraine record shows widespread
Egesen moraines of Younger Dryas age (Ivy-Ochs
et al. 1999). In sharp contrast to the situation in
Great Britain,Younger Dryas cooling is not reflected in the insect record at Lobsigensee.
The decisive warming that occurred throughout
the North Atlantic region near the Oldest Dryas/
Boiling transition was abrupt. The 6180 transition
occurred within a century or less in Greenland ice
cores (Oeschger et al. 1985) and in Swiss sediment cores (Siegenthaler et al. 1984). In the
British Isles a cold and continental climate ended
suddenly, with a 10°C rise of mean annual temperature in an interval of 300 to 800 years. By
12,500 14Cyr BP British climate was as warm as
that of today (Atkinson et al. 1987). At Ballybetagh in Ireland, re-establishment of vegetation began about 12,600 14Cyr BP;the warmest interval
of B611ing-Older Dryas-Aller6d time was at
11,900-12,400 14Cyr BPand was followed by climate deterioration leading to the Younger Dryas
reversal (Cwynar and Watts 1989). In central Europe north of the Alps, a treeless steppe-tundra
with some dwarf birch yielded swiftly to reforestation near the biozone Ia/Ib (Oldest Dryas/
B6lling) boundary (Lotter et al. 1992). This
change is shown by an abrupt rise in the arboreal/
non-arboreal ratio in numerous pollen profiles; it
was followed by the spread of trees across Europe
during biozones lb, Ic, and II (B6iling, Older
Dryas, Aller6d). For example, vegetation in the
Swiss and adjacent French Alps changed markedly in parallel with b180 shifts in precipitation near
the biozone Ia/Ib (Oldest Dryas/B61iing) boundary (Eicher and Siegenthaler 1976; Eicher et al.
1981). Here an increase in juniper coincides with
the initial 6180 rise. At some sites reforestation by
birch began during this early juniper phase, but at
most sites it occurred shortly after the juniper rise.
This birch reforestation is by far the most dramatic
late-glacial event recorded in pollen diagrams on
the northern margin of the Alps. Subsequently, a
rise of pine marks the beginning of biozone II
(Aller6d); open pine forest with some birch and
juniper characterizes biozone III (Younger
Dryas); and finally a transition from open to
125
G.H. DENTON ET AL.
closed pine forest marks the biozone III/IV
(Younger Dryas/Preboreal) boundary (Lotter
1991; Lotter et al. 1992).
A similar situation prevailed in the Pyrenees adjacent to the Atlantic Ocean and the Mediterranean
Sea (Jalut et al. 1992). Here the Oldest Dryas/
Bolling boundary is marked by a rapid increase in
birch, a beginning of pine expansion and, at some
localities, the spread of juniper.Again the rapid rise
of arboreal pollen, which represents the most pronounced change in the vegetation records, signals
the real change from an open glacial landscape to
an extensive woodland. Subsequently, a rise in pine
characterizes the Allerod biozone, whereas the
subsequent Younger Dryas biozone is notable for
expansion of Artemesia and juniper, with concurrent decline of birch and pine.
The circulation regime of the surface and deep
waters of the North Atlantic Ocean also changed
fundamentally near the Oldest Dryas/Bolling transition. The polar front retreatedrapidly (Ruddiman
and McIntyre 1981; Bard et al. 1987); the planktonic foraminifers in core Troll 3.1 show that warm
surface water leaked into the Norwegian Sea about
13,400 14Cyr BP and reached near-modern values
by about 13,100 14C yr BP, signaling a major
change in surface circulation (Lehman and Keigwin 1992a).
The circulation regime of the deep ocean
showed renewed strong thermohaline production
of North Atlantic Deep Water (NADW), including
Lower North Atlantic Deep Water (LNADW)
formed by convection in the Nordic Seas. Carbonisotope measurements of foraminifers from highsedimentation core RC11-83 from Cape Rise off
South Africa show that the first significant contribution of NADW flow to the Southern Ocean occurred between 12,700 and 13,100 14C yr BP
(Charles and Fairbanks 1992). Because it is closely
tied to surface conditions in the Nordic Seas, rapid
turnover of LNADW most likely began with the
warm surface water incursion documented in the
Troll 3.1 core (Lehman and Keigwin 1992a). This
is consistent with Cd/Ca and 613Cratios from core
EN-1200-GGCI near Bermuda Rise in the path of
deep-water flow (Keigwin et al. 1991).
The rapid increase in northwardheat flux caused
by this fundamental change in surface and deep circulation must have contributed to the dramatic
warming in Europe near the Oldest Dryas/B611ing
transition through the demonstrated linkage between North Atlantic sea-surface temperature and
European climate (Rind et al. 1986; Lehman and
126
Keigwin 1992a). The fact that this is the earliest
such dramatic warming over Europe is consistent
with the notion that LNADW, as it now forms, was
cut off during the LGM and Oldest Dryas time.
Thus the flooding of warm and salty surface water
into the Nordic Seas and the consequent initiation
of the Nordic Heat Pump of Imbrie et al. (1992),
with its extensive import of ocean heat (Lehman
and Keigwin 1992b), was probably the most importantevent in the switch to an interglacial climate
mode.
A less dramatic earlier warming pulse occurred
in Europe within Oldest Dryas time. The primary
evidence comes from using records of the prominent late-glacial climate signal that swept across
Europe near the Oldest Dryas/B611ingtransition to
unravel the deglacial chronology of the European
Alps. This prominent Oldest Dryas/B611ing signal
is coherent from Greenland to Switzerland. The
southeastern end of this transect is anchored by the
paleoclimate record at Gerzensee in Switzerland
that shows the characteristic 6180 and vegetation
shifts at the biozone Ia/Ib (Oldest Dryas/B611ing)
boundary (Eicher and Siegenthaler 1976) (Fig. 6).
For us, the importantpoint is that not only Gerzensee but numerous other lakes and bogs with similar
evidence for the Oldest Dryas/B61iing warming
event are located well within the boundaries of the
expanded glacier system of the European Alps at
the LGM, a situation previously pointed out by
Schltichter (1988). It is particularly pertinent that
many such sites occur in major valleys and passes.
At each site, the Oldest Dryas/Bolling transition is
located above the base of the core, and hence occurred subsequent to deglaciation. The major point
is that extensive deglaciation of the European Alps
had already occurred before the Oldest Dryas/
Bolling warming that is so prominent in paleoclimate records across the North Atlantic Ocean and
Europe. In fact, the areal distribution of the core
sites in Switzerland indicates that, by the time of
this Oldest Dryas/B61iing transition, mountain glaciers were already confined to upper reaches of
deep alpine valleys or to inter-valley mountain
massifs. The radiocarbon age of the Oldest Dryas/
Bolling transition can only be placed roughly at
12,700-13,000 14Cyr BP, because of an atmospheric 14C plateau of several hundred years duration
(Zbinden et al. 1989). If it is assumed that this transition was simultaneous within the Swiss and Austrian Alps, as it seems to have been across Europe,
then this means that considerable mountain deglaciation antedated 12,700-13,000 14Cyr BP.
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
We interpret this early deglaciation of the Alps
as evidence for an Oldest Dryas rise in mountain
snowlines in southern Europe within the Oldest
Dryas chron. We attributethis snowline rise to climate warming. A contrary suggestion has been
made for mountain glaciers of the French Vosges,
near the site of the Grande Pile bog (Seret et al.
1990, 1992). It is inferred from indirect evidence
that these glaciers were likely to have achieved
their maximum in middle Pleniglacial time when
slightly warmer climate was presumably accompanied by increased snowfall; glaciers were supposedly contracted during cold late Pleniglacial conditions because they were starved of precipitation.
That such a scenario could not apply to the Austrian
and Swiss Alps is demonstrated by stratigraphic
sections near Innsbruck and Ztirich, respectively
(Patzelt and Resch 1986; Schltichter et al. 1987).
At both sites till representing glacier advance to the
maximum position of the last glaciation is bracketed by radiocarbon ages of about 29,000 and
14,000 14C yr BP; therefore the maximum extent of
glaciers coincided with late Pleniglacial cold and
dry conditions, not with the more humid conditions
of middle Pleniglacial time. Supporting this contention is the lack of radiocarbon evidence in the
Swiss and Austrian Alps for ice recession during
late Pleniglacial time.
What was the timing of the Oldest Dryas deglaciation of the Swiss and Austrian Alps, and what
was its duration? The initiation of this recession is
defined at Lake Zurich. Here deglaciation from the
Zurich Stadial moraines, which are situated only
20 km behind the maximum ice-margin position of
the LGM (Schltichter et al. 1987; Schliichter and
Rothlisberger 1995), occurred close to 14,600 14C
yr BP (Lister 1988). Applying to a core from Lake
Zuirich, this date comes from a twig in non-disturbed deposits immediately above ice-rafted and
tectonized sediments deposited as glacial ice
cleared the lake. Hence, we place the beginning of
snowline rise and climate amelioration at shortly
before 14,600 14Cyr BP. Widespread deglaciation
deep into the mountains driven by this snowline
rise was complete before the onset of warming near
the Oldest Dryas/B6lling transition.
From the amount of mountain snowline rise, we
can estimate the fraction of the total full glacialinterglacial change encompassed in this Oldest
Dryas warming of the European Alps. Full-glacial
mountain snowline depression relative to the AD
1850 position (Maisch 1992) was about 1100 m
(Furrer 1991 and references therein). Snowline
Geografiska Annaler * 81 A (1999) · 2
lowering during the Zurich Stadial was slightly
less. By the end of the widespread Oldest Dryas deglaciation, the remaining alpine ice was confined to
the upper reaches of the deep valleys and to the inter-valley massifs. Glacier termini most probably
stood near the Gschnitz or Clavadel moraines, as
shown by the position of these moraines relative to
radiocarbon-dated pollen cores. The Clavadel moraines register a snowline lowering of about 430 m
relative to the AD 1850 position, and the Gschnitz
about 670 m (Furreret al. 1987). Hence the snowline rise during Oldest Dryas time was at least 330
m and perhaps as much as 670 m. This amounts to
about one-third (and perhaps more) of the full glacial-interglacial snowline rise.
By comparison with pollen records on the north
flank of the Alps at Gerzensee and Faulenseemoos
(Eicher and Siegenthaler 1976), Tourbiere de Chirens (Eicher et al. 1981), Rotsee (Lotter 1991), and
Soppensee (Lotter et al. 1992), the climate amelioration that caused the snowline rise during Oldest
Dryas time was not sufficient to allow reforestation
north of the Alps. Instead, relatively severe conditions persisted, and as a result the vegetation remained open. In the French and Swiss Alps this
open Oldest Dryas landscape was dominated by a
grass assemblage and some shrubs, along with alpine and steppe herbs. However, all of the abovementioned sites lie within the limits of expanded
alpine ice at the LGM. Therefore, we must look to
other sites outside the last glacial limits for evidence of initiation of the Oldest Dryas environment. In the Pyrenees for example, palynologic
data indicate two phases of vegetation development prior to the Oldest Dryas/B61iing transition
(Jalut et al. 1992). The first phase reflected cold,
open steppe and semi-desert conditions that persisted through the LGM until about 15,000 14Cyr
BP. The second phase began about 15,000 14Cyr BP
with a large increase in pollen concentration, a
spread of juniper, and perhaps an initial establishment of birch and pine in some places. This second
phase was succeeded by the typical rise of birch at
the Oldest Dryas/Bolling transition. Hence, in the
Pyrenees the Oldest Dryas climate warming episode began about 15,000 14C yr BP. In southern
France at both Les Echets and Le Bouchet, the vegetation record indicates that the arid, cold late
Pleniglacial ended about 15,000 14Cyr BP with the
rise of Artemisia, Chenopodiaceae, and Caryophyllaceae accompanied by the spread of steppe,
which marks the first warming. These botanical
events signal the late Pleniglacial/Oldest Dryas
127
G.H. DENTON ET AL.
transition. Within the sparse available radiocarbon
control, then, this transition is equivalent to the beginning of ice recession from Lake Zurich. Farther
north in the Netherlands, a sharp rise in Artemisia
about 14,000 14Cyr BP during the Oldest Dryas was
followed by the immigration of large birch trees in
the Blling (Hammen and Vogel 1966).
Recession of the Scandinavian-Barents Sea icesheet complex suggests that Oldest Dryas snowline rise and associated climatic amelioration was
not restricted to the Alps but was widespread
throughout Europe (Andersen 1981). By the time
of the Oldest Dryas/B611ing transition, the southern margin of the Scandinavian Ice Sheet stood at
the Luga (dated between 12,650 and 13,200 14C yr
BP)-North Lithuanian-Wolin-Halland ice-marginal position, and thus had retreated about 250 km
from its maximum LGM position (Andersen 1981;
Lundquist and Saarnisto 1995). Almost 175 km of
this recession was from the Vepsovo/Pomeranian
and Krestay/Kalinin ice-marginal positions, which
have estimated ages of 15,000 14Cyr BPand 14,500
14Cyr BP, respectively (Andersen 1981). This recession took place during the Rauniss interstade of
still-severe climate, radiocarbon-dated between
13,250 14Cyr BP and 14,300 14Cyr BP (Serebrjannyj et al. 1970; Raukas 1976; Velichko and Faustova 1986). Hence European alpine glaciers and
the Scandinavian Ice Sheet both showed considerable recession in the 1600 14C-yrinterval preceding the Oldest Dryas/B61iing warming. In the case
of the Scandinavian Ice Sheet, this recession ended
with readvance to the Luga moraine; in the case of
the alpine system, the recession ended with ice
margins close to the Gschnitz or Clavadel moraines in the inner Alps. The Oldest Dryas climate
amelioration devastated the glacier system of the
European Alps because of its susceptibility to a
moderate snowline rise, while the massive Scandinavian Ice Sheet lost marginal ice but still remained largely intact.
The marine-based, western margin of the coalesced Barents Sea and Scandinavian Ice Sheets
also showed Oldest Dryas recession. One piece of
evidence comes from a low-b180 spike in cores
from off the continental margin (Jones 1991;
Weinelt et al. 1991). The first age estimate of this
spike was 14,500 14Cyr BP from low-sedimentation-rate core PS-21295 in the Fram Strait off the
Barents Sea continental shelf (Jones and Keigwin
1988). But younger ages for this spike of 13,60014,000 14Cyr BPcome from cores with higher sedimentation rates farthersouth in the Norwegian Sea
128
(Sarnthein et al. 1992). This first pronounced 6b80
decrease in the Norwegian Sea is taken to represent
a great iceberg outburst caused by collapse of marine-based ice sheets from the Norwegian and Barents continental shelves (Jones 1991; Sarnthein et
al. 1992; Landvik et al. 1998). Likewise, a pronounced 6180 depletion is recorded at 13,10014,300 14Cyr BP just northwest of the Faeroe Islands during the Heinrich 1 event (Rasmussen et al.
1997); the magnitude of this spike implies significant iceberg contribution from the Faeroe and
Shetland Islands, Scotland, and Scandinavia. This
is consistent with the conclusion of Fronval et al.
(1996) from marine-sediment cores off the western
margin of the Scandinavian Ice Sheet. A near-basal
radiocarbon date from the Troll 3.1 core is also
consistent with this premise, because it suggests
ice recession from the outer Norwegian continental
shelf at 14,700-15,000 14C yr BP (Lehman et al.
1991). These early conclusions are supported by
new radiocarbondates from marine sediment cores
that show recession of grounded portions of ice
sheets from the Barents and the Norwegian continental shelves at 14,500-14,800 14Cyr BP (Bischof
1994; Haflidason et al. 1995; Svendsen et al.
1996). This recession of marine margins corresponds with retreat of terrestrial Scandinavian ice
during the Rauniss interstade, as well as with Oldest Dryas recession of mountain glaciers in the
Swiss Alps. Taken together, these data imply that
the Oldest Dryas amelioration affected the European land mass over a wide latitude range and
marked the onset of the last termination.
It should be noted that the Oldest Dryas climate
amelioration in Europe was not nearly as dramatic
as the subsequent warming near the Oldest Dryas/
Bolling transition. Paleoclimate records from the
Rauniss interstade of Oldest Dryas age in northern
Europe and from correlative Oldest Dryas pollenbearing sediments in France and Switzerland indicate that the climate still remained severe enough
to preclude reforestation north of the Alps. Even at
the end of Oldest Dryas time, Europe was still
marked by a cold tundra and steppe environment,
reflecting cold North Atlantic sea-surface temperatures. This is consistent with the fact that NADW
of the current mode, with a strong LNADW component produced in the Nordic Seas from the inflow of warm, salty water, was not evident in Oldest
Dryas time (Sarnthein et al. 1992). In contrast, the
dramatic warming near the Oldest Dryas/B611ing
transition was tied to a major thermohaline switch,
with the initiation of strong LNADW production
Geografiska Annaler · 81 A (1999) - 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
Drafted by R. D. Kelly Jr. 1998
Fig. 7. Southern Chilean Lake District paleoclimate record (from Fig. 2) transferredto a calendar-year time scale and compared with
Greenland and Antarctic ice-core stable isotope records. The ice-core records are from Sowers and Bender (1995) and Steig et al. (1998).
The ice cores are placed in a common calendar-year time scale by linking the Antarctic 6 80 record of ice to the Greenland chronology
by matching the signatures of trapped gasses in the ice cores (after correcting for the difference in ice and gas ages).
and greatly increased poleward heat flux in the
North Atlantic Ocean (Lehman and Keigwin
1992a, b; Sarnthein et al. 1994).
Global implications of Chilean Andes and New
Zealand data
Figure 5 shows a detailed 6180 record of benthonic
foraminifers from core TR163-31B from the eastern equatorial Pacific Ocean, taken to be representative largely of Northern Hemisphere ice volume.
Fig. 6 compares the paleoclimate records from the
Chilean Andes and Southern Alps of New Zealand
with those from the North Atlantic region. Fig. 7
displays the Southern Hemisphere middle-latitude
paleoclimate records compared with the Antarctic
and Greenland stable isotope signals in the Taylor
Dome, Byrd, and GISP2 ice cores. The following
implications are derived from these figures and
from the text.
Relative to present-day values, snowline and/or
treeline depression at the LGM was about the
same in the Southern Alps (875 m) and the ChilGeografiska Annaler . 81 A (1999) · 2
ean Andes (1000 m) as in many mountain areas
in the Northern Hemisphere. Moreover, the timing of the LGM was similar in both polar hemispheres (with the possible exception of Antarctica). The implication is of planetary cooling at
the LGM that was about equivalent in both polar
hemispheres.
* The last glacial-interglacial transition began
abruptly in both hemispheres (at least outside of
Antarctica) with a warming pulse at close to
14,600 14Cyr BP (17,300 cal. yr BP) within Oldest Dryas time. In Chile the last glacier expansion of the LGM culminated at 14,550-14,805
14Cyr BP, and was followed by massive recession. Note that, because of new radiocarbon
dates given in Denton et al. (1999b), this age is
slightly older than reported in Lowell et al.
(1995). Nothofagus increased significantly at
the Canal de la Puntilla and Huelmo sites
(Moreno 1998; Moreno et al. 1999) about
14,600 14Cyr BP. At a number of sites in Valle
Central over nearly 2° of latitude, pollen diagrams show at about 14,000 14Cyr BP the first
strong influx of thermophilic elements of the
129
G.H. DENTON ET AL.
North Patagonian Evergreen Forest since before
49,892 14C yr BP (Heusser et al. 1999). In the
Southern Alps of New Zealand, significant glacier recession had occurred prior to 13,500 14C
yr BP after an advance to near-maximum positions between 14,000 and 15,000 14Cyr BP. In
northern New Zealand reforestation of the
Waikato lowlands began shortly after 14,700
14Cyr BP (Newnham et al. 1989) and much of
the open glacial landscape of North Island was
rapidly reforested by podocarp and hardwood
trees by 14,000 14C-yrBP. The 6180 curve of benthonic foraminifers from core TR163-3 lB in the
eastern tropical Pacific Ocean shows that a
steady increase in benthonic 6180 gave way to a
unilateral decrease about 14,500 14Cyr BP (this
6180 shift is not adjusted for unknown mixing
time of the ocean at that date) (Fig. 5). We take
this change to mark the beginning of the termination of glaciation in the Northern Hemisphere, because it implies a fundamental change
from volume increase to volume decline of
Northern Hemisphere ice sheets. In core SU8118 in the eastern North Atlantic Ocean off Portugal, the termination also began at about 14,500
14C yr BP as defined by the planktonic 6180
record (Fig. 6). As detailed above, the southern
margin of the Scandinavian Ice Sheet retreated
significantly during Oldest Dryas time and
grounded portions of ice sheets receded from
both the Barents and the Norwegian continental
shelves between 14,500 and 14,800 14Cyr BP.
Also as mentioned above, the alpine glacier system in the Swiss Alps contracted from the forelands to the inner valleys between 14,600 14Cyr
BP and 12,700-13,000 14Cyr BP. In addition, by
the end of Oldest Dryas time, the southern margin of the Laurentide Ice Sheet in North America
had already retreated to the position of the Port
Huron moraine system (Mayewski et al. 1981)
and the southern Cordilleran Ice Sheet had already undergone considerable recession (Booth
1987).
Another major step of the last glacial/interglacial transition was a decisive warming pulse at
12,700-13,000 14Cyr BP. In the Chilean Andes
this step is marked by the rapid spread of a
closed-canopy North Patagonian Evergreen
Forest throughout the lowlands of the Chilean
Lake District and Isle Grande de Chiloe. Development of this forest culminated at 12,00012,200 14C yr BP, when climate conditions were
close to full interglacial values. Plant species in-
130
dicative of wet moorland environments disappeared, possibly from a southward shift of the
westerlies storm belt (Moreno et al. 1999). In
the North Atlantic region this second step is
marked by the abrupt warming near the Oldest
Dryas/B611ing transition shown by isotope,
methane, and dust records in the Greenland ice
cores, as well as by isotope and vegetation
records in Swiss lacustrine sediments. The modem mode of North Atlantic thermohaline circulation resumed, with strong production of
LNADW and pronounced warming in the Nordic Seas from the inflow of warm, salty water.
The insect records from Great Britain and the
Swiss Plateau show that this warming near the
Oldest Dryas/B11ling transition culminated in
temperatures nearly as high as those of today
(Atkinson et al. 1987; Amman 1989a). At the
height of B611ingwarming, snowline on Swiss
alpine glaciers probably rose to the position it
occupied during the AD 1850 highstand of the
Little Ice Age (Maisch 1982).
A Younger-Dryas-ageclimate reversalcharacterized the North Atlantic and perhaps the middlelatitude Southern Hemisphere paleoclimate
records. The Greenland ice-core 6180 signal in
Figs 6 and 7 shows the typical form of the North
Atlantic climatic deterioration,which began after
the peak B6iling warmth and culminated in the
Younger Dryas cold reversal. In the Chilean Andes, pollen diagrams show climate reversal (expansion of Podocarpus nubigena followed later
by decline of mesic North Patagonian taxa) beginning at 12,000-12,200 14Cyr BP after peak
late-glacial warmth(Heusser et al. 1999; Moreno
et al. 1999). Unfortunately,fire disturbanceof the
vegetation is indicated at many sites during
Younger Dryas time, complicating paleoclimate
interpretation. However, the Taiquemo pollen
record escaped the influence of fire and shows
episodic cooling between 11,360 and 10,355 14C
yr BP (Heusser et al. 1999). In this regard,the sediment record from proglacial Lago Mascardi in
Argentina, located near Mt. Tronador only 115
km east of our key pollen sites at Canal de la Puntilla and Fundo Llanquihue, shows evidence not
only of rapid ice recession beginning at 13,000
14C yr BP and peaking at 12,400 14C yr BP, but also
of a subsequent reversal of trend that culminated
in a Younger-Dryas-age glacier readvance between 11,400 14C yr BP and 10,200 14C yr BP
(Ariztegui et al. 1997) (Figs 2,6). In the Southern
Alps of New Zealand, an advance of Franz Josef
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
Glacier occurred at the beginning of Younger
Dryas time at 11,050 14C yr BP (Denton and
Hendy 1994) (the beginning of Younger Dryas
time in the NorthAtlantic region is most accurately dated to 11,030 14Cyr BP in Switzerland; Hajdas et al. 1995). AnotherYounger-Dryas-ageglacier advance is dated to 10,100 14Cyr BPin the upper Cropp River Valley (Basher and McSaveney
1989; Lowell et al. 1995). Exposure dates indicate that the prominent Lake Misery moraine
complex in Arthur'sPass is YoungerDryas in age
(Ivy-Ochs et al. 1999). These lattermoraines represent snowline lowering of about 260-360 m below the Little Ice Age position, compared with a
similar value of 200-300 m for the Egesen moraines of Younger Dryas age in the Swiss Alps
(Kerschner 1978, 1985; Maisch 1982, 1987).
Thus there is mounting evidence for YoungerDryas-age readvanceof SouthernAlps glaciers in
New Zealand, although caution in accepting such
a conclusion is urged by Mabin (1995), McGlone
(1995), and Singer et al. (1998).
The Chilean Andes glacial geologic record features pronounced maxima during the LGM at
29,400 14C yr BP, 26,700 14C yr BP, 22,29522,570 14Cyr BP, and 14,550-14,805 14Cyr BP
(Fig. 2 and Table 1). The vegetation record from
Taiquem6 shows a series of striking maxima of
grass pollen during middle and late Llanquihue
time. The duration of these Gramineae events
are listed in Table 1 (Heusser et al. 1999) as
44,520-47,110 14C yr BP for T-11; 32,10535,764 14Cyr BP for T-9; 24,895-26,019 14Cyr
BP forT-7; 21,430-22,774 14Cyr BP for T-5; and
13,040-15,200 14Cyr BP for T-3. Also shown in
Table 1 are the durations (not the peaks) of the
Heinrich lithic events over 20° of latitude in the
North Atlantic Ocean (Elliot et al. 1998). Heinrich events H- 1, H-2, H-3, H-4, and H-5 show a
reasonably close match to Gramineae events in
Chile. With the exception of the situation at
29,400 14Cyr BP, the Chilean glacier maxima
fall within or close to the Heinrich-Gramineae
events. This reinforces our earlier conclusion
that there is a broad correspondence between
North Atlantic Heinrich events and Chilean Andes glacier maxima (Lowell et al. 1995).
However, the far greater chronologic and
stratigraphicdetail now available (Denton et al.
1999; Moreno et al. 1999; Heusser et al. 1999)
shows an importantnew insight. The last glacier
maximum (14,550-14,805 14C yr BP) in the
Chilean Andes occurs at the beginning of the
Geografiska Annaler * 81 A (1999) · 2
long H-1 event in the North Atlantic Ocean (Table 1). It is notable that much of the H-1 lithic
event correlates with glacier recession in the
Andes. This relationship is reinforced by detailed pollen records showing that a marked rise
of Nothofagus and then the invasion of thermophilous tree species into the Chilean Lake
District (Moreno et al. 1999; Heusser et al.
1999) was also coeval with much of the H- 1 lithic event in the North Atlantic Ocean. Moreover,
recession of terrestrial Scandinavian ice and
mountain glaciers in the European Alps was coincident with the younger part of the H-1 lithic
layer. The implication is that the H- 1 lithic event
encompasses not only ice-sheet expansion onto
North Atlantic continental shelves during global
cooling, but also the subsequent unstable collapse of marine-based ice. That this advancecollapse sequence may be characteristic of
Heinrich events is suggested by the age of the
major Chilean glacier advance of 22,29522,570 14C yr BP near the beginning of the H-2
lithic event.
The New Zealand Southern Alps record features
the last two of the major glacier maxima in the
Chilean Andes. In addition, New Zealand glaciers show maxima at 17,500 14C yr BP and
16,200 14Cyr BP. Maxima may also have been
achieved just before and just after deposition of
the Raupo interstadial bed, now dated between
19,740 and 18,780 14Cyr BP, as well as during
Younger Dryas time.
A potential weakness of our trans-Pacific comparison is that some of the glacier fluctuations are not
yet recorded in both data sets. One could therefore
argue that the climate signals do not match during
these parts of the record. But we think it more likely
that these differences merely reflect the fact that
glacial moraine sequences are inherently incomplete.
Questions
What caused the atmospheric cooling at the LGM
evident in the snowline and/or treeline records
from the Chilean Andes and Southern Alps? We argue that this cooling is approximately synchronous
and of the same magnitude in both hemispheres,
thus pointing to the primary role of overall atmospheric cooling (from greenhouse gas content or reflectivity) rather than simply redistribution of heat
on the planet from changes in ocean circulation.
131
G.H. DENTON ET AL.
What caused the atmospheric warming evident in
the first two major steps (Oldest Dryas and Oldest
Dryas/Billing) of the last termination? These two
steps, registered in both hemispheres, abruptly terminated the LGM, together bringing the atmosphere from full-glacial to nearly full-interglacial
temperature in about 1600 14Cyr. The synchrony
and magnitude of these two steps in the two hemispheres, at least outside of Antarctica, implicates
greenhouse gas as the direct cause of atmospheric
warming ratherthan simply switches in ocean heat
transfer. The Byrd ice-core record in Antarctica,
calibrated with the Greenland ice-core chronology
by trapped gas, shows only a minor change in atmospheric CO2 during these steps (Neftel et al.
1988; Staffelbach et al. 1991; Sowers and Bender
1995; Blunier et al. 1997). If the new ice-core chronologies are correct, the implication is that jumps
in atmospheric CO2 were not responsible for the
abrupt warming steps that terminated the last ice
age simultaneously in both hemispheres. This
leaves changes in the inventory of atmospheric water vapor as the most likely factor to have caused
the two major atmospheric warming steps during
the last termination. It follows that a decreased inventory of atmospheric water vapor was the most
important cause of LGM planetary cooling
(Broecker 1994).
An alternate explanation is that Northern Hemisphere ice sheets responded to seasonality forcing
and, in turn, drove the last termination through
their global thermal impact. The radiocarbon chronology of the two important initial steps of the termination makes this explanation highly unlikely.
These two steps, which encompassed 1600 14Cyr,
culminated in B6iling time with atmospheric temperatures approaching interglacial values from
Greenland to the Swiss Plateau, as well as in the
Chilean Andes. Reference to Fig. 5 shows that ice
volume was then still near the maximum value of
the LGM. From this phasing of events it thus seems
most probable that, ratherthan causing this crucial
warming step of the termination, the ice sheets responded to it by a greatly increased ablation rate
registered in the sea-level record as meltwater
pulse IA (Fairbanks 1989).
What caused the changes in water-vapor production implied by paleoclimate data during the initial
Oldest Dryas step of the termination? It is unlikely
that the thermal impact of a decrease in ice-sheet
albedo or elevation is implicated, because glacier
retreat started simultaneously in both hemispheres
132
(see also Broecker and Denton 1990). A major
North Atlantic thermohaline switch is also an unlikely trigger, as the Heinrich 1 ice-rafting event of
Oldest Dryas age actually suppressed overturning
in the North Atlantic Ocean to its lowest level
(Sarnthein et al. 1994). Likewise, our data suggest
that any southward shift of the Southern Hemisphere westerlies that might implicate important
sea-ice changes in the Southern Ocean was delayed
until near the Oldest Dryas/ B611ing transition
(Moreno et al. 1999).
The effect of rising summer insolation in the
Northern Hemisphere is commonly cited as a trigger for the ocean-atmosphere reorganization of the
last termination (Imbrie et al. 1992, 1993). For example, Broecker and Denton (1990) pointed out
that terminations commonly occur during rises toward maxima in seasonality at middle and high latitudes in the Northern Hemisphere. But our data
imply that any such Northern Hemisphere effect
was translated to middle latitudes of the Southern
Hemisphere by an atmospheric process. In the absence of major changes in Northern Hemisphere
ice sheets or of major switches in North Atlantic
thermohaline circulation, it is difficult to identify a
high-latitude mechanism that was triggered by rising summer insolation during Oldest Dryas time.
This difficulty is compounded by the fact that during Oldest Dryas time the Greenland ice-core 680
record is dissimilar to paleoclimate records elsewhere in the Northern Hemisphere and in Chile,
probably because of a regional climate signal induced by the Heinrich 1 iceberg influx into the
North Atlantic Ocean. This suggests to us that increased water-vaporproduction in the tropics is the
most likely source of the initial Oldest Dryas step
of the last termination, perhaps from forcing by
half-precession insolation, whose amplitude is
controlled by eccentricity (McIntyre and Molfino
1996; Berger and Loutre 1997).
What caused the changes in water-vapor production inferred to have produced decisive warming
near the Oldest Dryas/Billing transition? A key
point is that near-interglacial warmth was achieved
in early Blling time in both hemispheres (from
Greenland to the Swiss Plateau in the North Atlantic region, as well as in the Chilean Andes). It is
very unlikely that this event was driven by the direct thermal impact of the Northern Hemisphere
ice sheets, because reference to Fig. 5 shows that
they were then still close to their maximum LGM
volume. Instead the chronologies reviewed here afGeografiska Annaler · 81 A (1999) - 2
INTERHEMISPHERIC
LINKAGEOF PALEOCLIMATE
DURINGTHELASTGLACIATION
ford strong circumstantial evidence that this decisive warming was caused by an abrupt switch of
thermohaline circulation to the modern mode of
operation, with renewed downwelling in the Nordic Seas. Because near-interglacial warmth was
also achieved in middle latitudes of the Southern
Hemisphere, the decisive switch of thermohaline
circulation must also have triggered increased production of atmospheric water vapor to near-interglacial values. As well as having a strong influence
on high-latitude Northern Hemisphere climate
(Lehman and Keigwin 1992b), renewed NADW
production may also have caused reduction of
Southern Ocean sea-ice extent and the southward
shift of the westerlies by the standard explanation
for coupling the hemispheres by thermohaline circulation (Weyl 1968). Such speculation is consistent with the Taylor Dome but not with the Byrd and
Vostok ice-core records from Antarctica near the
Oldest Dryas/B611ingtransition, as discussed in the
next section.
What caused the resumption of vigorous North Atlantic thermohaline circulation with a strong
LNADWcomponent near the Oldest Dryas/B6lling
transition? The radiocarbon chronologies given
here are consistent with the idea that the initial Oldest Dryas warming step triggered this decisive
event of the last termination. But how? The fact that
the termination began when ice sheets achieved
their maximum volume as recorded by the 6180
signal in benthonic foraminifers (Fig. 5) (Shackleton et al. 1988) suggests that the existence of large
ice sheets was a necessary condition for the initiation of the last termination. The Heinrich 1 event is
also implicated during the most recent termination,
because it occurs right at the time of the break in the
fundamental benthonic 8180 trends in core T16331B in Fig. 5. To explore this last implication, we
turn to a discussion of Heinrich events in the North
Atlantic Ocean.
Massive, short-lived discharges of icebergs into
the North Atlantic Ocean occurred each 700010,000 years during the gradual buildup phase of
the last 100,000-yr cycle (Heinrich 1988). These
outbursts left prominent Heinrich layers of icerafted debris with sharp lower boundaries deposited rapidly on the sea floor along the southern margin of the glacial-age North Atlantic Ocean
(Broecker et al. 1992; Bond et al. 1992, 1993;
Bond and Lotti 1995; Manighetti and McCave
1995; Manighetti et al. 1995). Heinrich deposits
are marked by high percentages of ice-rafted detriGeografiska Annaler · 81 A (1999) · 2
tus and low concentrations of foraminifers. They
become thinner from west to east across the Atlantic Ocean, and commonly are rich in detrital carbonate. These characteristics point to a source in
eastern Canada or western Greenland, with Hudson Strait a prime candidate, for most of the debris
in the main iceberg track. Ice-rafted grains from the
penultimate Heinrich layer deposited about 20,500
14C yr BP show a lead isotopic composition also
consistent with derivation of ice-rafted debris from
eastern Canada (Gwiazda et al. 1996). Revel et al.
(1996) argued that Scandinavian, British, and Icelandic ice caps contributed to the flux along the
eastern part of the main iceberg track, as well as to
the north of the main track. They thus suggested
that all ice caps surroundingthe North Atlantic, not
just the Laurentide Ice Sheet, experienced major
discharge of icebergs at the time of Heinrich
events. This is consistent with the conclusions
drawn from the 6180 record of a core taken northwest of the Faeroe Islands (Rasmussen etal. 1997).
From two important cores at the eastern end of the
maximum iceberg track, Bond and Lotti (1995)
likewise found evidence of synchronous discharges of icebergs from several ice sheets during Heinrich events. From a high-resolution marine-sediment record from the Irminger Basin, which registers ice-rafted debris from the Greenland and Norwegian Seas, Elliot et al. (1998) concluded that
lithic layers corresponding to Heinrich events were
deposited by enhanced calving from Nordic ice
caps, as well as the Laurentide Ice Sheet, over 20°
of latitude in the North Atlantic Ocean.
A comparison of Greenland and North Atlantic
paleoclimate records reveals that Heinrich events
occurred at or near the culmination of Bond cooling hemicycles superimposed on the overall buildup phase of the last 100,000-yr cycle (Bond et al.
1993; Rasmussen et al. 1997). Following each
Heinrich event, the sea-surface temperatures
warmed as the North Atlantic Ocean made a short
but failed excursion toward the interglacial mode
of circulation. The situation is even more complex,
because millennial-scale Dansgaard-Oeschger oscillations are superimposed on each cooling hemicycle (Bond et al. 1993).
The cause of the huge discharges of icebergs
during Heinrich events is not clear. Thus such a discharge could be a trigger for climate change, a response to climate forcing, or both (Broecker et al.
1993; Broecker 1994, 1995a, b). There are at least
three possible explanations for these massive iceberg discharges into the North Atlantic Ocean. One
133
G.H. DENTON ET AL.
is from internally triggered surges of the Laurentide Ice Sheet (Broecker et al. 1992; MacAyeal
1993). This surge mechanism has the advantage of
explaining rapid deposition of Heinrich layers.
One apparent disadvantage is that clear geologic
evidence of down-trough flow of an ice stream
draining the heart of the Laurentide Ice Sheet has
not been found in Hudson Strait (Miller et al.
1993). Another disadvantage is that several ice
sheets, not just Laurentide ice, apparently had an
increased iceberg flux at the time of Heinrich discharge events (Bond and Lotti 1995; Fronval et al.
1996; Revel etal. 1996; Rasmussen etal. 1997; Elliot et al. 1998).
A second potential origin for the Heinrich icebergs is simply increased calving from expanded
ice sheets on North Atlantic continental shelves
during maxima of Bond cycles. The main disadvantage of this mechanism is that it cannot easily
account for the catastrophic outbursts of icebergs
implied by the sharp lower boundaries of the Heinrich ice-rafted debris that buried organic horizons
on the ocean floor within a few hundred years and
accumulated so rapidly that it is only little disturbed by bioturbation (Manighetti and McCave
1995; Manighetti et al. 1995).
A third potential source for a massive iceberg
outburst is from collapse of unstable grounded
marine ice on North Atlantic continental shelves,
with associated ice-sheet downdraw by accelerated discharge of feeder ice streams. For example,
Fronval et al. (1996) recently argued that those
portions of ice sheets grounded on extensive shelf
areas would be unstable, subject to collapse by
marine downdraw. This suggestion is very similar
to the mechanism previously proposed by Ruddiman and McIntyre (1981) to explain in many
North Atlantic marine sediment cores an interval
nearly barren of planktonic foraminifers. This interval, now known to be equivalent to the Heinrich
1 event, was interpreted to result from the most
massive influx of icebergs and meltwater to the
North Atlantic Ocean during the entire 10,000-yr
deglaciation. From a consideration of the mechanics of the last deglaciation by Denton and Hughes
(1981), this massive early iceberg influx was attributed to the collapse of unstable marine-based
ice sheets on continental shelves, with interior
downdraw of feeder ice streams (Ruddiman and
McIntyre 1981). An important focus of marine
downdraw was Hudson Strait (Denton and Hughes
1981), a potential source of detrital carbonate.
However, a difficulty again involves the lack of
134
evidence for ice flow eastward through Hudson
Strait (Miller et al. 1993).
By the scenario of unstable marine collapse,
each cooling Bond hemicycle superimposed on the
long buildup phase of the 100,000-yr cycle involved expansion of marine-based ice onto continental shelves to reach a maximum at the beginning of a Heinrich event. Each cooling hemicycle
then ended with collapse of destabilized ice from
these shelves during a Heinrich event. Although
marine ice-sheet collapses could be triggered internally, it is unlikely that all ice caps would then collapse simultaneously. Rather, as pointed out by
Revel et al. (1996), the fact that not only the Laurentide Ice Sheet, but also the Greenland Ice Sheet,
the Icelandic ice cap, the Scandinavian Ice Sheet,
and the Irish-Scottish ice cap, were all sources of
enhanced calving during the time of Heinrich
events points to common external forcing. A warming pulse and/or sea-level rise could initiate a near
simultaneous collapse of extensive marine-based
sectors of ice sheets on North Atlantic continental
shelves, thus explaining iceberg outbursts from all
ice caps during the time of Heinrich events.
We prefer the thirdoption because it is consistent
with several aspects of our Chilean paleoclimate
record. The chronologic data in Table 1 suggest that
the last five Heinrich lithic events correspond with
Chilean grass maxima. The implication is that the
Bond hemicycles represent global cooling that not
only drives large ice sheets onto North Atlantic
continental shelves, where they are potentially unstable, but also causes Chilean mountain-glacier
advance. The last mountain-glacier maximum in
Chile occurs just at the beginning of the H-1 lithic
event. The ensuing collapse of mountain glaciers,
along with the invasion of the southern Chilean
Lake District by Nothofagus and thermophilic tree
species, is coeval with deposition of much of the H1 lithic layer. This situation suggests that major recession of Chilean mountain glaciers was coeval
with the unstable collapse of grounded ice from
North Atlantic continental shelves that produced
most of the icebergs responsible for the H-1 lithic
layer. The pollen record shows that climate warming was the trigger for the Chilean glacier collapse.
Retreat of terrestrialScandinavian ice and of European alpine glaciers suggests that warming may
also have triggered the marine ice-sheet collapse
from North Atlantic continental shelves.
We also prefer the third option because it features an external trigger of marine collapse mechanisms that would act nearly simultaneously on all
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCI MATE DURING THE ILASTGILACIATION
expanded marine margins. It has long been suggested that marine-based sectors of large ice sheets
could be susceptible to very rapid grounding-line
recession from hydrostatic instability triggered by
sea-level rise or climate warming (Weertman 1957,
1961, 1974; Denton and Hughes 1981; Denton et
al. 1986; Hughes 1987). A potential climate trigger
for the youngest Heinrich collapse is documented
in the Oldest Dryas paleoclimate record in both
hemispheres. As just mentioned, the terrestrial
margin of the Scandinavian Ice Sheet, as well as the
glacier system of the European Alps, receded extensively during Oldest Dryas time, coincident
with the youngest Heinrich collapse. As also just
mentioned, our South American data show that
Oldest Dryas climate warming coincided with
much of the youngest Heinrich lithic layer. This
warming in southern South America is the most extensive since sometime prior to MIS 4 (Heusser et
al. 1999). It is mirroredby similar warming in New
Zealand (Newnham et al. 1989). Thus the Oldest
Dryas climate warming that caused recession of
terrestrial glaciers in both hemispheres could also
have triggered unstable collapse of marine icesheet sectors on North Atlantic continental shelves.
Such a collapse would have kept the North Atlantic
cold, both from the influx of icebergs and from the
consequent suppression of thermohaline downwelling and hence northwardocean heat transport.
This could explain a very curious aspect of the
Greenland ice-core 18sOrecord, namely the striking absence of any abrupt warming signal until
near the Oldest Dryas/Boiling transition. In this
way, persistence of cold conditions in Greenland
during Oldest Dryas time could reflect a massive
influx of icebergs into the North Atlantic Ocean. In
contrast to the stable isotope record, the first rise of
methane in the Greenland ice cores, probably indicative of climate change in the tropics, occurred
about 17,000 cal. yr BP (Chappellaz et al. 1993) and
hence was correlative with the initial Oldest Dryas
warming pulse recognized in the southern Chilean
Lake District, New Zealand, and Europe.
The differences in two kinds of paleoclimate
signals from core SU81-18 from the eastern North
Atlantic Ocean off Portugal support the interpretation that regional cooling from the Heinrich 1 ice
collapse dominated the Greenland isotope record
in Oldest Dryas time. The 5180 signal of the planktonic species Globigerina bulloides in SU81-18
shows a marked depletion beginning at 14,500 14C
yr BP (Fig. 6) (Bard et al. 1987), compatible with
the timing of the Heinrich 1 ice collapse. And yet
Geografiska Annaler · 81 A (1999) · 2
sea-surface temperatures in SU81-18 reconstructed from micropaleontological transfer functions
show a significant drop in Oldest Dryas time.
Again, this is consistent with decisive cooling of
the North Atlantic Ocean from the flood of Heinrich 1 icebergs and the consequent collapse or nearcollapse of the glacial mode of thermohaline overturn (Sarnthein et al. 1994). Yet another reason for
preferring the third option is that deglaciation of
marine-based sectors of both the Scandinavian
(Lehman et al. 1991; Haflidason et al. 1995) and
Barents Sea (Bischof 1994; Svendsen et al. 1996)
Ice Sheets coincided with the Heinrich 1 ice-rafting event in the North Atlantic Ocean. In core V-2381 in the North Atlantic, Heinrich ice rafting began
about 14,800 14C yr BP, culminated at close to
14,100-14,300 14Cyr BP, and was in rapid decline
after 13,700 14Cyr BP (Bond and Lotti 1995). All
this is consistent with the initial reduction of Northern Hemisphere ice volume as deduced from the
benthonic 6180 record in eastern Pacific core
TR163-31B (Fig. 5) (Shackleton et al. 1988).
The consequence of selecting this third option is
that the magnitude of each Heinrich collapse could
well be linked to a combination of ice-sheet size
and of the intensity of the warming trigger. The final Heinrich collapse that peaked at 14,100 14Cyr
BP could then be the largest, because by this time
the ice sheets had reached their maximum volume
of the last 100,000-yr cycle (Fig. 5).
The key to the last termination is that ice sheets
had to grow sufficiently large to produce a Heinrich
collapse massive enough to cripple the glacial
mode of the Atlantic salinity conveyor. The conveyor circulation then reorganized into its interglacial mode of operation. The probable reason for
this reorganization is that the massive collapse during the time of the Heinrich 1 event flushed out so
much marine-based ice peripheral to the North Atlantic Ocean that, at the end of the collapse, the iceberg flux dropped dramatically. Through the long
buildup phase, the ice-sheet behavior during Bond
cycles could be an important regulator of North Atlantic near-surface salinity. The steady background
iceberg influx during buildup could keep North Atlantic salinity low. A Heinrich collapse could produce the lowest salinity at the peak of each cycle.
North Atlantic salinity could then jump after each
Heinrich collapse, thus triggering thermohaline
switches and a rise of North Atlantic sea-surface
temperatures. The largest salinity jump would follow the massive Heinrich 1 event, because it
flushed out the most ice. At the same time the sum135
G.H. DENTON ET AL.
mer radiation field from orbital forcing was approaching an interglacial value, which might favor
evaporation, and hence net export of water vapor,
from the northern North Atlantic Ocean. Perhaps
these two factors together trip the North Atlantic
into its interglacial circulation mode, with a strong
LNADW component.
By this scenario, the resumption of the interglacial mode of North Atlantic thermohaline circulation occurs because of, not in spite of, the growth
of huge ice sheets peripheral to the North Atlantic
Ocean. The unstable behavior of marine shelf portions of these ice sheets is the key ingredient. The
reason that ice sheets have to grow to such huge dimensions is to produce an unstable collapse large
enough to deplete the grounded marine ice reservoir and its feeder ice streams. Once the resulting
massive iceberg influx ceases, the North Atlantic
becomes largely free of icebergs for the first time
since early in the long buildup phase of the
100,000-yr asymmetric cycle. North Atlantic salinity can then rapidly increase and trigger renewed
vigorous downwelling in the Nordic Seas. This explains the curious observation that terminations occur just when ice sheets achieve their largest volume. It also is consistent with the observation that
a prerequisite for a sharp, complete termination
seems to be excessive ice volume in huge Northern
Hemisphere ice sheets (Raymo 1997).
What caused the asymmetric 100,000-yr glacial
cycles of late Quaternary time? Terminations are a
dominant feature of these cycles. From radiocarbon-dated paleoclimate sequences, we argued
above that renewal of the modern mode of NADW
formation, with a strong LNADW component, was
the key event of the last termination. From this we
assume that the current mode of thermohaline circulation is the essential feature of late Quaternary
interglaciations. We now discuss the implications
of this assumption.
The modern mode of thermohaline overturn in
the northern North Atlantic Ocean switched on
abruptly near the Oldest Dryas/B61iing transition.
This event was accompanied by rapid reforestation
of Europe, including the northward spread of thermophilic trees, presumably because of the demonstratedlink between North Atlantic heat import and
European climate. Even more-extensive forest
cover occurred in northern Europe during Eemian
(MIS 5e) time (Behre 1989), consistent with vigorous formation of LNADW from the inflow of
warm, salty surface water. This is in accord with a
136
paleoceanographic reconstruction of the Norwegian and Greenland Seas (Kellogg 1980).
Deep-water temperatures of MIS 5e were similar to those of the Holocene (Labeyrie et al. 1987).
Deep-water temperature in the Pacific (Chappell
and Shackleton 1986; Labeyrie et al. 1987) and
southern Indian (Labeyrie et al. 1987) Oceans
dropped to glacial values about 115,000 years ago
during the MIS 5e/5d transition. Deep Atlantic waters reached glacial values in two major steps, one
at about 115,000 years ago and the other at about
75,000 years ago during the MIS 5/4 transition (Labeyrie et al. 1987). After these cooling steps, the
deep ocean waters remained at glacial temperatures throughout MIS 3 and MIS 2 (Chappell and
Shackleton 1986; Labeyrie et al. 1987).
The paleoenvironment of northern Europe
(Beaulieu and Reille 1984, 1989, 1990, 1992; Behre 1989; Guiot 1990; Seret et al. 1992) suggests
that these cooling steps of deep-water temperature
could well reflect the progressive shutdown of
warm LNADW early in the last glacial cycle. At the
close of the Eemian, a cooling episode that left an
environment markedonly by shrubtundrain northern Europe during the Herning stadial (Behre
1989) implies sharp curtailment of warm LNADW
formation as the primary cause of the deep-water
temperaturedecline about 115,000 years ago. During the subsequent Br6rup (MIS 5c) and Odderade
(MIS 5a) European interstades, each lasting as
much as 10,000 years, northern Europe was reforested, with coniferous forests in the north giving
way to deciduous forests in the south (Behre 1989).
The inferred north-to-south climate gradient was
distinctly steeper than that of the Eemian (Behre
1989). During MIS 5a the Norwegian and Greenland Seas were seasonally ice covered, with the formation of LNADW in smaller amounts than today
(Kellogg 1980), in accord with the Odderade paleovegetation reconstruction.
The onset of cold Pleniglacial conditions in
northern Europe immediately after the Odderade
interstade (Behre 1989) is consistent with the shutdown of warm LNADW as the cause of the second
major cooling step in North Atlantic deep water at
the MIS 5/4 transition about 75,000 years ago.
Warming sufficient to cause reforestation in northern Europe did not occur during the cold Pleniglacial climate conditions of MIS 3 and MIS 2, even
at the peaks of Greenland ice-core interstades 8
(Denekamp in Europe), 12 (Hengelo in Europe),
14 (Glinde in Europe), and 16 (Oerel in Europe)
(Kolstrup and Wijmstra 1977; Behre 1989; Guiot
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATEDURING THE LAST GLACIATION
1990; Pons et al. 1992; Schulz et al. 1998). Rather,
these interstades were marked only by reversion
from tundra to open, treeless shrub tundra (Behre
1989). There probably was some thermohaline
downwelling near the Faeroe Islands in these interstades (Rasmussen et al. 1997), but the lack of reforestation in northernEurope suggests that it must
have been very weak. Continuous cold, treeless
conditions in Europe during MIS 3 and MIS 2 are
consistent with a lack of LNADW formation from
warm, salty water and, consequently, with continuous cold deep-water temperatures. These considerations suggest that the modem mode of North Atlantic thermohaline circulation switched off in two
steps early in the last glacial cycle when ice sheets
were still small. It did not switch back on again until near the Oldest Dryas/B61iing transition, when
ice sheets were near their maximum LGM volume.
The result is that early in the last glacial cycle the
climate system became largely detached from orbital forcing except for oscillations superimposed
on the long buildup to the LGM (see also Imbrie et
al. 1993) (Fig. 5).
The formation of LNADW could have been
switched off early in the last cycle by orbital forcing of ice-sheet growth and the consequent iceberg
influx into the critical areas of downwelling in the
Nordic Seas. But subsequent orbital forcing in the
opposite sense was not sufficient to switch it back
on. Instead, an extraordinary mechanism must
have operated in the climate system in order for the
North Atlantic salinity conveyor to switch abruptly
into a vigorous interglacial mode of operation at
what seems to be a very unlikely time, namely, just
when Northern Hemisphere ice sheets were close
to their maximum LGM volume. We argue above
that, rather than precluding a circulation switch to
an interglacial mode, the growth of these ice sheets
to their maximum volume was actually a prerequisite for the decisive ocean-atmosphere reorganization near the Oldest Dryas/B611ing transition, because it set up the necessary conditions for the huge
Heinrich 1 ice-sheet collapse into the North Atlantic Ocean, which we think is the key precursor
event.
The marine-based portions of ice sheets grounded on North Atlantic continental shelves exhibited
unstable collapse behavior at the culmination of
each of the shorter Bond cycles superimposed on
the long growth phase of the last asymmetric
100,000-yr cycle. This long and fluctuating growth
phase ended abruptly when ice sheets became large
enough to produce an unstable collapse sufficiently
Geografiska Annaler
81 A (1999)
2
massive not only to cripple the glacial mode of
North Atlantic thermohaline circulation, but also to
flush out grounded continental-shelf ice and feeder
ice streams adjacent to it. As a result, the influx of
icebergs into the North Atlantic declined greatly
once the collapse was over. In the absence of an iceberg influx and of conveyor circulation, the salinity
of the northern North Atlantic Ocean increased
rapidly, probably from net export of water vapor,
until a vigorous interglacial mode of LNADW formation was initiated.
By this scenario, the length of an individual
100,000-yr cycle is set by two main factors that
work in combination to produce the massive marine ice-sheet collapse that resets the North Atlantic salinity conveyor. The first factor is the length of
time required to accumulate the excess ice volume
necessary for a large collapse. The second is the
timing of the warming trigger for this collapse, presumably set by the effect of eccentricity on the amplitude of precession and half-precession insolation. The timing of the warming trigger, and hence
of the collapse, need not be unique. For example,
if the Heinrich 1 collapse had been too small to reset the salinity conveyor, then a later collapse could
have triggered the termination at about 12,000
years ago (Northern Hemisphere summer insolation maximum). In this way terminations can occur
at any of several times within intervals of high eccentricity. Thus the role of orbital forcing of terminations is to cause the climate warming that triggers a massive ice-sheet collapse and then, from
high summer insolation, to impose a precipitation/
evaporation ratio that promotes this interglacial
circulation mode by the net export of water vapor
from the northern North Atlantic Ocean. This is
compatible with the argument given earlier that,
during terminations, insolation forcing simply acts
as a trigger for the fundamental reorganization of
a non-linear system, but does not control the amplitude of the resulting reorganization.
From the paleoclimate record, we argue above
that the critical step of the last glacial/interglacial
transition was the resumption of the modern mode
of NADW formation, with a vigorous LNADW
component. We also argue that it took the entire
buildup phase of each asymmetric glacial cycle to
produce ice sheets large enough to reset the modern
mode of global thermohaline circulation through a
massive Heinrich marine-ice-sheet collapse. The
notion that it is so difficult to turnon the interglacial
mode of thermohaline circulation suggests that
100,000-yr glacial cycles emerged because the
137
G.H.DENTONETAL.
modem interglacial mode of thermohaline circulation became easier to switch off and harder to
switch on than it was prior to the 100,000-yr regime. But what caused the development of such an
asymmetric behavior of the interglacial mode of
thermohaline circulation? For a possible answer
we turn to the topography of the Greenland-Scotland submarine ridge across the North Atlantic
Ocean. Wright and Miller (1996) postulated that
the changes in the height of the Greenland-Scotland ridge (from varying activity of the Icelandic
mantle plume) regulated the Neogene flow of
northward-flowing warm surface water into the
Nordic Seas and hence controlled NADW circulation. Rise of this ridge from mantle-plume activity
is postulated to have brought mid-Pliocene Arctic
warmth to a close (Wright and Miller 1996). A further rise of this ridge between 950,000 and 600,000
years ago could have caused the asymmetric behavior of North Atlantic thermohaline circulation
that we suggest was characteristic of late Quaternary time. This would have initiated asymmetric
100,000-yr climate cycles because of the necessity
for the buildup of large ice sheets to reset the interglacial mode of thermohaline circulation by unstable collapse from continental shelves.
What caused the climate cooling reversal thatfollowed peak Bolling warmth and culminated in the
YoungerDryas cold pulse? Any explanation must
address the 1000 14C-yr length of the Younger
Dryas cold pulse and its occurrence at very near the
peak of maximum Northern Hemisphere summer
insolation. Also, it should clarify why sea level
continued to rise throughout this climate deterioration and even during the culminating Younger
Dryas cold pulse (Fairbanks 1989; Bard et al.
1996).
If our paleoclimate reconstructions from Chilean pollen records are correct, then peak Bollingage warmth was short-lived, and was followed by
a climate reversal that persisted through much of
late-B61ling and Aller6d time in both hemispheres.
Also, if our reconstructions from the Chilean Andes and Southern Alps, along with the reconstruction of Ariztegui et al. (1997) from the Argentine
Andes, are correct, then a Younger-Dryas-age cold
event at the end of this long deterioration may have
marked both hemispheres. The radiocarbon dates
of the Younger Dryas readvance of Franz Josef Glacier in the Southern Alps of New Zealand (11,050
14C yr BP; Denton and Hendy 1995) and of the
Younger Dryas isotope shift in Switzerland
138
(11,030 14C yr BP; Hajdas et al. 1995) imply rapid
transmittal of an atmospheric cooling signal at the
beginning of Younger Dryas time. The nearly
equivalent amount of snowline lowering in the European Alps and the Southern Alps suggests that
the Younger Dryas event represented a reversal of
about 35% toward full-glacial conditions (IvyOchs et al. 1999). This suggestion must be considered tentative, however, until more Southern Hemisphere data on Younger-Dryas-age climate events
are available (for alternative view see Markgraf
1991,1993; Mabin 1995; McGlone 1995; Singer et
al. 1998).
A leading explanation for the basic shape of the
late-glacial climate signal is that the North Atlantic
salinity conveyor turned on at maximum strength
near the Oldest Dryas/B611ing transition and subsequently weakened by dilution of salt from melting of continental ice sheets caused by the peak
Boiling warmth. The increasingly sluggish thermohaline overturn culminated in Younger Dryas
shutdown-or near shutdown-of LNADW production, probably triggered by an additional pulse of
icebergs and/or meltwater (Broecker 1990, 1992;
Broecker et al. 1990; Fanning and Weaver 1997).
New modeling studies of the stability of North
Atlantic thermohaline circulation show that cooling and wind-stress feedbacks destabilize those
modes of circulation without deepwater formation
(Schiller et al. 1997). In the modeling experiments
deep-water formation resumes soon after meltwater input ceases. The implication is that only a meltwater pulse lasting 1000 14C yr could have produced the North Atlantic Younger Dryas event.
Two sources have been suggested. One is diversion
of North American glacial lakes from the Mississippi drainage into the Gulf of St Lawrence drainage (Rooth 1982; Broecker et al. 1988, 1989), or
the Baltic Ice Lake into the North Atlantic (Bj6rck
et al. 1996). Potential problems are that freshening
of the Champlain Sea apparentlydid not occur until
about 500 14Cyr after the beginning of the Younger
Dryas (Rodrigues and Villas 1994) and that a distinct 6180 minimum has not been found off Nova
Scotia (Keigwin and Jones 1995). A second potential source of freshwater is an iceberg influx from
the Arctic due to the collapse of floating ice shelves
and of ice sheets grounded on high-latitude continental shelves (Mercer 1969). The available uplift
curves from Svalbard, Franz Josef Land, and the
Queen Elizabeth Islands are at least permissive of
this hypothesis (Blake 1975, 1993; Forman et al.
1996, 1997). The collapse of marine ice sheets
Geografiska Annaler * 81 A (1999) - 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
couldeasily encompass1000 14Cyr andcouldexplain continued sea-level rise during Younger
Dryastime.
In any case, our New Zealanddata suggest a
SouthPacificatmosphericsignalrightatthebeginning of YoungerDryas time. If a NorthAtlantic
thermohalineswitch is the ultimatecause of the
YoungerDryas reversal,then the atmosphericeffects of this switchmusthavebeentransmittedrapidly,bothto the tropics(Robertset al. 1993;Hughen et al. 1996; Clappertonet al. 1997) and to the
South Pacific region (Denton and Hendy 1994).
The new trapped-gaschronologiesof theByrdicecore recordin Antarcticasuggest that changes of
atmosphericCO2contentwerenottheoriginof this
atmosphericeffect (SowersandBender1995;Blunier et al. 1997). This again leaves a linkage between thermohalinecirculationand the tropical
productionof watervaporas themostlikely source
of this atmosphericsignal.
Did fluctuations in the strength of NADW cause
out-of-phase climate variations in the Northern
and Southern Hemispheres during the last glacia-
tion? Crowley (1992) suggested that changes in
oceanicheattransferacrossthe equatorcausedby
variationsin the productionof NADW would alternatelyfavornetheatgainin one andthentheother hemisphere.This is becausethe amountof heat
transportedto Antarcticaby NADW would be
morethancompensatedfor by the heattransferred
across the equatorby northward-flowing
shallow
currentsin the Atlantic Ocean. During times of
strong NADW production,the SouthernHemisphere would lose heat. The reverse situation
wouldoccurduringtimesof weakproduction.The
middle-latitudepaleoclimaterecords given here
suggest thatthis mechanismdid not dominateatmosphericclimateoscillationsduringthe last glacial/interglacialtransition.Middle-latitudeChilean climatewarmedin OldestDryastime when the
productionof NADWreachedits lowest value,but
What is the role of millennial-scale oscillations in glaciersalso recededin the NorthernHemisphere
producing abrupt climate changes? A pervasive at the same time. Moreover,middle-latitudeChilmillennial-scaleoscillation that persists through eanclimatewarmedalmostto full-interglacialvalmajorclimate transitionshas been recognizedin ues coincidentwitha switchto themodernmodeof
Holocene,late-glacial,andglacialrecords(Denton NADW formationnear the Oldest Dryas/B611ing
and Karlen 1973; Karlen and Denton 1976; transition.
O'Brienetal. 1995;Bondetal. 1997).Infact,Bond
et al. (1997) suggestedthatmillennial-scaleoscillationsarethe pacemakerof rapidclimatechange. Southern Ocean sediment cores and Antarctic
In this regard,the timingof abruptinitialwarming ice cores
in theChileanrecordmayhavebeenpacedandam- Ourradiocarbonchronologyof pollenprofilesand
plifiedby themillennial-scaleoscillationssuperim- morainesshows that the estimatesof 875 m for
posedon theLGM.Forexample,thelastmajormil- mountainsnowlineloweringin the SouthernAlps
lennial-scaleglacierpulse reacheda maximumat of New Zealand(Porter1975)andof about1000m
14,550-14,805 14Cyr BP, just priorto the abrupt for snowline (Porter1981) and treeline lowering
warmingin OldestDryastimethattriggeredthelast (Morenoet al. 1999; Heusseret al. 1999) in the
termination.Althoughnot documentedin Chile, a southernChileanLake Districtpertainto glacier
similar situation occurredlater in the Northern maximawithinthe long, cold intervalof the LGM.
Hemisphere,where the PortHuron(GreatLakes) Thesevaluesof snowlineloweringareverysimilar
and Luga (north-centralEurope) ice-sheet read- to those in many mountainranges elsewhere on
vances immediatelypreceded the abruptOldest Earth(BroeckerandDenton 1990). Such uniform
Dryas/B611ing
warming(e.g. Denton and Hughes loweringof mountainsnowlinesstronglysuggests
1981).These relationshipssuggestthatmillennial- globalcoolingof the samemagnitudein bothhemscale pulsesmay be importantin settingthe timing ispheresduringthe LGM. From our middle-latiof abruptwarmingsteps. For example, the cold tude SouthernHemispheredata,we arguethatthe
pulse at 14,550-14,805 14Cyr BP could have de- LGM was terminatedby several abruptclimate
layed the onset of the last termination.Then the changes that are registeredin both hemispheres.
subsequentwarm phase of the millennial-scale Most importantarethose duringthe OldestDryas
transition.Topulse could have reinforcedabruptOldest Dryas and nearthe Oldest Dryas/B611ing
climatewarming.Thus the timingand abruptness getherthese changesbroughtthe planetfromfullof the Oldest Dryas/B11lingwarmingcould have glacial to near-interglacialtemperaturesin about
been affectedby a millennial-scaleoscillation.
1600 14Cyr.
Geografiska Annaler * 81 A (1999) · 2
139
G.H.DENTONETAL.
We argue that neither the global cooling of the
LGM, nor the abruptglobal warming steps that terminated the last glaciation, can be explained simply by redistribution of heat on the planet from
ocean circulation changes. Rather,either the reflectivity of the planet or the greenhouse gas content of
the atmosphere had to change. Because of the basic
similarities of the Chilean and New Zealand paleoclimate records with the classic North Atlantic/
European record, we stress changes in atmospheric
greenhouse gas content (probably predominantly
the water vapor inventory, somehow linked to thermohaline switches), as the major mechanism of interhemispheric climate linkage. During the decisive ocean-atmosphere reorganization near the
Oldest Dryas/B611ing transition, a switch in thermohaline circulation must have been connected
with a change in the production of water vapor.
During the precursor Oldest Dryas warming a
change in atmospheric water-vapor content was
not obviously connected with a thermohaline circulation switch but was probably linked to precession effects in the tropics.
But there is a potential problem with our approach. Namely, it has been suggested for more
than two decades that sea-surface temperature
change in the Southern Ocean leads ice-volume
change in the Northern Hemisphere into and out of
interglaciations (Hays et al. 1976; Hays 1978).
This inference comes from offsets in individual
marine sediment cores of the 6180 values of benthonic and planktonic foraminifers and the sea-surface temperature estimated from diatoms, foraminifers, and radiolaria (Howard and Prell 1984,
1992; Labeyrie etal. 1986,1996; Labracherie etal.
1989; Pichon etal. 1992). For example, sea-surface
temperature is taken to lead ice-volume change by
several thousand years during the transition into
the current interglaciation. In the absence of benthonic 6180 change, the early Holocene peak in
Southern Ocean sea-surface temperatures is taken
to indicate a lead of Southern compared to the
Northern Hemisphere climate in the transition out
of the current interglaciation (Hays 1978).
To complicate the situation, the stable isotope
record calibrated with new trapped-gas chronologies from the Byrd ice core in Antarctica does not
show the abrupt changes during the last termination so evident in the North Atlantic (Fig. 7). Instead, a gradual change in the stable-isotope signal
in the Byrd ice core began several thousand years
before the first abrupt warming registered in
Greenland ice cores at the Oldest Dryas/Bolling
140
transition (Sowers and Bender 1995; Blunier et al.
1998). Also, late-glacial stable isotope oscillations
recorded in the Byrd and Vostok ice cores during
Bolling/Aller6d and Younger Dryas times were out
of phase with their Greenland counterparts(Fig. 7)
(Sowers and Bender 1995; Blunier et al. 1997,
1998). These results are taken to be consistent with
Southern Ocean marine cores in showing a Southern Hemisphere lead through important climate
transitions. This Antarctic lead seemingly precludes any hypothesis that climate changes are a response to Northern Hemisphere events (Blunier et
al. 1998). The asynchronism strongly suggests,
moreover, that the interhemispheric connection
cannot be through the atmosphere. Rather, it favors
models that call for overall ocean heat extraction
from the Southern Hemisphere during times of vigorous North Atlantic thermohaline circulation
(Crowley 1992; Stocker et al. 1992).
Because they suggest that any Antarctic asynchrony is not hemisphere-wide, our paleoclimate
data from Chile and New Zealand are not in accord
with an ocean heat-transfermechanism that applies
to the entire hemisphere. Moreover, a heat-transfer
mechanism by itself cannot explain the overall
warming of both hemispheres during the last deglaciation. One way around this dilemma is to postulate that the asynchrony is confined to regions
south of the Antarctic Circumpolar Current, thus
pointing to a mechanism that applies only to Antarctica rather than to the Southern Hemisphere as
a whole. In this regard, Broecker (1998) attributed
out-of-phase late-glacial climate oscillations between Greenland and Antarctica to a bipolar seesaw behavior of thermohaline circulation. By this
mechanism, decreased (increased) downwelling of
NADW in the North Atlantic Ocean was replaced
by increased (decreased) downwelling in the
Southern Ocean. The resulting out-of-phase import
of ocean heat into the North Atlantic and Southern
Oceans is what would have caused out-of-phase
late-glacial climate signals in the Greenland and
Antarctic ice cores. Broecker (1998) and Broecker
and Henderson (1999) suggested an early trigger of
thermohaline overturn in the Southern Ocean from
local summer insolation during at least the last two
terminations. A possible greaterAntarctic lead during the penultimate than during the last termination
may reflect stronger early insolation forcing
(Broecker and Henderson 1999).
But there are also complications with this approach. One is that almost all the important marine
sediment cores showing a lead of Southern Ocean
Geografiska Annaler * 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
Fig. 8. Index map of the Southern
Ocean showing the location of major oceanic boundaries, including
the Subtropical (STF), Subantarctic (SAF), Polar (PF), and southern
Antarctic Circumpolar Current
(SACCF) Fronts. Deep-sea sediment cores MD88-770 (46°01'S,
(41°36'S,
96°28'E), RC11-83
9°48'E), MD84-527
(43°49'S,
51°19'E), and DSDP 594 (45°32'S,
174°57'E) as well as the Taylor
Dome (77°47'S, 2400 m), Byrd
(80°01S, 1530 m), and Vostok
(78°28'S, 3488 m) ice cores, are
also shown, as well as many of the
pertinent Southern Ocean cores
discussed in the literature. Frontal
boundaries and terminology are after Orsi et al. (1995).
lie northof the PolarFrontin Subantemperatures
tarctic waters (Fig. 8). Hence these cores come
from nearly the same SouthernHemispherelatitudes as our terrestrialdata,which we interpretas
showingpaleoclimatechanges similarto those in
the North Atlantic/Europeansector of the planet
transition.A secduringthelastglacial/interglacial
ond is thatthe newTaylorDome ice core (77°47'S,
2400 m) (Mayewskiet al. 1996;Steig et al. 1998),
situatedin Antarcticaabouthalfwaybetween the
Byrd(80°0I'S, 1530m) andVostok(78°28'S,3488
m) ice cores (Fig. 8), showsa late-glacialpaleoclimaterecordsimilarto thatin Greenlandandin the
middlelatitudesof the SouthernHemisphere.For
example,unlike the Byrd or Vostokrecords,that
from the TaylorDome shows an abruptwarming
near the Oldest Dryas/B611ingtransitionand a
YoungerDryas reversal (Mayewski et al. 1996;
White and Steig 1998; Lehman1998; Steig et al.
1998).
Thustherearenow two fundamentallydifferent
paleoclimatesignalsfromAntarcticice cores, one
in phaseandthe otherout of phasewith NorthAtlantic and middle-latitudeSouthernHemisphere
Geografiska Annaler · 81 A (1999) · 2
records.This has led to yet a thirdexplanationof
Antarcticice-corepaleoclimaterecords.Whiteand
Steig (1998) andSteig andWhite(1998) suggested
thatthe differencesin ice-corerecordsreflectheterogeneityin the upwellingpatternsof NADW in
the SouthernOcean.
Antarctic CircumpolarTransect
A resolution of the differing interpretationsof
SouthernHemispherepaleoclimaterecordsis not
at hand.A majorproblemis thewidespreadlackof
theextensivedatingbasenecessaryto placethedetails of deglaciationon a hemisphericand interhemisphericscale.Weproposethattheissuecanbe
clarifiedby obtainingdetailedmarineand terrestrial paleoclimaterecordsalong a proposedAntarctic CircumpolarTransect(ACT) between 42°
and 54°S (Fig. 8). The essential feature of this
transectalong the outer marginof the Southern
Ocean aroundAntarcticawould be detailedAMS
radiocarbon chronologies. Thus the proposed
transectis designed to lie largely within or just
northof Subantarcticwaters(where foraminifers
141
G.H. DENTON ET AL.
are common for radiocarbon dating) and to pass
across mountainous areas of South America, New
Zealand, Tasmania, and South Georgia (where
wood and gyttja are common for radiocarbon dating). Comparison of the resulting vegetation,
mountain-glacier, benthonic 8'80, planktonic
b180, and sea-surface-temperature records should
clarify the issue of leads and lags of Southern Hemisphere paleoclimate along the outer margin of the
Southern Ocean relative to both high-latitude Antarctica and the Northern Hemisphere. The results
should also clarify the issue of a bipolar seesaw of
thermohaline circulation, as well as the amplitude
of paleoclimate changes in different sectors of the
Southern Ocean.
The currentsituation along the marine portion of
ACT is that three marine sediment cores have an
AMS radiocarbon chronology. Fig. 9 shows the
dated isotope record for RC11-83 (41°36'S,
9°481E) (Charles et al. 1996). This particular core
does not show an obvious lead of planktonic over
benthonic 6180 at the initiation of the last deglaciation, placed before 15,000 14Cyr BP largely on the
basis of a single date in this interval. The transitory
oscillation or plateau of planktonic 6180 during deglaciation is thought to be correlative with the Antarctic Cold Reversal (Billing in age, Blunier et al.
1997) in the Vostok ice-core record (Charles et al.
1996). Such a correspondence could be expected
because the subantarctic waters near RC11-83 are
the ultimate source of most precipitation in the region of Vostok (Koster et al. 1992). However, the
6180 minimum at about 12,360 14Cyr BP in RC1183 is coincident with the full development of a
closed-canopy North Patagonian Evergreen Forest
at 12,200 14Cyr BP in the southern Lake District of
Chile. Also the subsequent isotope signal is compatible with the reversal in climate trend evident in
pollen records in the southern Lake District
(Heusser et al. 1999; Moreno et al. 1999). Thus it
is difficult to understand how the late-glacial
planktonic 5180 signal from RC 11-83 is taken to be
correlative with the Vostok record and to show a
1000-yr lead of Southern relative to Northern
Hemisphere climate.
Marine sediment core MD88-770 (46°01'S,
96°28'E), calibrated with 16 acceptable AMS
radiocarbon dates, shows a small lead of planktonic 5180 and summer sea-surface temperatures over
benthonic 6180 at the initiation of the last deglaciation (Fig. 9) (Labeyrie et al. 1996). Bracketing radiocarbon dates suggests that surface water began
to warm at sometime between 16,780 and 13,500
142
14Cyr BP, whereas benthonic 6180 began to change
shortly before 12,150 14C yr BP. The reason for the
large delay in response of benthonic b180 relative
to the change in marine sediment core TR 163-31 B
(Fig. 7) is not known. In any case, interglacial values of sea-surface temperature were reached at
MD88-770 by 12,150 14Cyr BP.
Marine sediment core MD84-527 (43°49'S,
51 ° 19'E), calibrated by 15 AMS radiocarbondates,
together with core MD84-551 (55°00'S, 73° 17'E),
calibrated by three AMS dates, shows that the deglacial 6180 decline in the Southern Ocean began
after 16,510 and prior to 13,000 14Cyr BP (Fig. 9)
(Bard etal. 1990). Between 13,000 and 12,000 14C
yr BP, the sea-surface temperatures reached interglacial values (Labracherie et al. 1989; Bard et al.
1990). There followed transitory oscillations, but
none is recorded in all indicators and so it is not
known if they are global or regional in origin (Bard
etal. 1990).
The potential value of ACT can be illustrated by
examples from the few paleoclimate records now
available in the pertinent latitudinal band of the
Southern Hemisphere.
* The Byrd and Vostok stable isotope records
from Antarctica are not proxies for hemispherewide paleotemperature changes during the last
glacial-interglacial transition, because they differ significantly from the Taylor Dome, New
Zealand, and southern Chilean Lake District
paleoclimate signals.
* Although Antarctic warming inferred from the
stable isotope record at the Byrd ice core seems
to have preceded the first abrupt warming in
Greenland ice-core records (near the Oldest
Dryas/B61ling transition), it does not necessarily follow that Southern Hemisphere climate as a
whole led Northern Hemisphere climate into the
last glacial/interglacial transition. This is because the Chilean and New Zealand paleoclimate records both show abrupt warming at
about 14,600 14Cyr BP that correlates with similar abruptwarming in the Northern Hemisphere
and is quite different in timing and character
from the gradual early warming documented in
the Byrd ice core. Moreover, the abrupt warming in the Greenland ice cores near the Oldest
Dryas/B61oingtransition occurred at least 1600
14Cyr after deglaciation began elsewhere in the
Northern Hemisphere.
* Although the chronology of the last deglaciation
in the Southern Ocean remains confusing, all
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
a.
MD88-770
SummerSST (°C)
2
-
4
6
8
10I
5
Foraminiferal 6180 (%o)
4
3
2
5700
-<~v
g
cm-
0.
40
0.
,
oO-C
Diatoms --
~-\
8~~~~~~--2
;/:=
~->
Foraminifers
Foraminifers
~<~*~~
-
'f>|-
% -,._._
,'
2150
1,530
13500
216,780
Foraminifers
v-"
.
<
-^
-<
Diatoms
~
Benthonic
Y
~32,790
23
21,610
20,400
1,290
25,160
°25,010
2730,500
(
-
-'-37,220
44,620
C.
MD84-527
N. pachyderma 1. 6180 (%0)
b.
RC11-83
E
-O
0.
*0
0
Draftedby R. D.KellyJr. 1997
-
-
-0
28,620
-
-
30,580
Fig. 9. Foraminiferal records from the three well-dated deep-sea cores in the Southern Ocean. Isotope records from core MD88-770
show a small lead of planktonic 6180 and summer SST over benthonic )180. Interglacial SST values were reached by 12,150 14Cyr
BP (Labeyrie et al. 1996). Oxygen-isotope records from RC11-83 do not show an obvious lead of planktonic over benthonic 6180 at
the beginning of deglaciation; however, a plateau in planktonic 6180 during deglaciation is thought by Charles etal. (1996) to correlate
with the Antarctic Cold Reversal seen in the Vostok Ice Core. Based on gas chronologic comparisons of Byrd and GISP2 ice cores,
the Antarctic Cold Reversal is thought to have occurred during Boiling time (Sowers and Bender 1995). In any case, 6180 values typical
of interglacial temperatures were reached by 12,360 14Cyr BP (Charles et al. 1996). The third well-dated record comes from MD84527 where i180 values of Neogloboquadrina pachyderma indicate that full interglacial conditions were reached by 12,130 14Cyr BP
(Bard et al. 1990; Labracherie et al. 1989).
Geografiska Annaler * 81 A (1999) · 2
143
G.H. DENTON ET AL
three of the extensively AMS-dated marine sediment cores show near-Holocene values of seasurface temperatures, or else significant depletion of planktonic 6180, at close to 12,00012,500 14Cyr BP.This is in accord with the situation in the southern Chilean Lake District terrestrial pollen records (Fig. 5) (Heusser et al.
1999; Moreno et al. 1999) and the Argentine
Andes glacial record (Ariztegui et al. 1997).
This warm interval corresponds with peak
Bolling warmth in the North Atlantic region
(Fig. 5), and it is out of phase with the climate
oscillations in the Byrd and Vostok, but not in
the Taylor Dome, ice cores in Antarctica (Sowers and Bender 1995; Blunier et al. 1997; Steig
etal. 1998).
Figure 10 suggests that the Southern Hemisphere summer insolation peak at about 23,000
cal. yr BP(which is nearly out of phase with the
precession-forced summer minimum in the
Northern Hemisphere) did not cause an early
warming response in the New Zealand and
Chilean paleoclimate records during Termination I. In contrast, detailed pollen analysis of
cores and sections from the Canal de la Puntilla,
Llanquihue, Fundo Llanquihue, Bella Vista
Bluff, Dalcahue, and Mayol sites together show
that the cold, wet conditions of the LGM continued through the interval of high southern insolation right up to the time of the youngest glacial
peak at 14,550-14,805 14Cyr BP(Denton et al.
1999; Heusser et al. 1999; Moreno et al. 1999).
LGM conditions came to a close only after the
culmination of this youngest advance. The same
situation holds for New Zealand. In like fashion,
reference to Fig. 4 shows that significant early
warming is not obvious in the New Zealand
paleoclimate record during Termination II.
Summary
Combined glacial geologic and palynologic data
from the Chilean Andes and Southern Alps highlight the following features of middle latitude
Southern Hemisphere paleoclimate through the
LGM and the last termination.
The moraine chronology suggests that full-glacial or near-full-glacial climate conditions prevailed from about 29,400 to 14,550 14Cyr BPin
the southern Lake District-Isla Grande de Chiloe region. However, pollen records from Isla
Grande de Chilo6 suggest that full-glacial con144
Southern Lake District, Chile
(40°30'- 42°25'S; 72o25'- 73045'W)
Mean Summer Temperature
0
0)
a)
(U
10
Insolation (W/m2)
-- ·- ------ - ----
Drafted by R. D. Kelly Jr. 1998
Fig. 10. Comparison of our new paleoclimate record for the Chilean Lake District and the New Zealand Southern Alps derived
from glacial geologic and palynological data with summer solar
insolation at 45°N and 45°S over the last 40,000 years (Berger
1978). A rise in Southern Hemisphere summer insolation beginning 30,000 cal. years ago did not produce signs of early warming
in New Zealand and Chilean paleoclimate records.
ditions did not set in until approximately 26,000
14C yr BP.
* Within this long interval of the LGM, snowline
was depressed about 1000 m, relative to presentday values, during glacier advances into the outer Llanquihue-age moraine belt at 29,400,
26,760, 22,295-22,570, and 14,550-14,805 14C
yr BP. Additional glacial maxima may have been
achieved shortly before 17,800 14C yr BP and
again shortly before 15,730 14C yr BP. The coeval lowering of treeline suggests that mean summer temperature was depressed 6-8°C, compared to modern values, during the LGM.
* In the Southern Alps, the known glacial advances during the LGM culminated at 22,400,
17,700, 16,200, and 14,000-15,000 14C yr BP.
Geografiska Annaler
81 A (1999) · 2
LINKAGEOF PALEOCLIMATE
INTERHEMISPHERIC
DURINGTHELASTGLACIATION
Additional maxima may have occurred shortly
before 19,740 14C yr BP, shortly after 18,600 14C
yr BP, and in Younger Dryas time. Relative to
present-day values, snowline depression during
the most extensive advances was about 875 m
(Porter 1975).
The last glacial/interglacial transition began
with a decisive warming at 14,600-14,700 14C
yr BP in both New Zealand and the Chilean Andes. Subsequent details are not yet known in
New Zealand, except for late-glacial readvances. But a second and decisive warming pulse in
the region of the southern Lake District-Isla
Grande de Chiloe occurred at 12,700-13,000
14C yr BP. After an interval of near-interglacial
warmth, a late-glacial climate reversal set in
close to 12,200 14C yr BP and continued at least
into the early part of Younger Dryas time. Although strong contrary opinions have been expressed, this reversal of trend may well have culminated in a Younger Dryas event in both the
Chilean Andes and the Southern Alps.
The similarity in timing and magnitude of middlelatitude snowline lowering with that in the Northern Hemisphere suggests global atmospheric cooling during the LGM. When compared with paleoclimate records from the North Atlantic region, the
middle-latitude Southern Hemisphere terrestrial
data imply interhemispheric symmetry of the
structure and timing of the last glacial/interglacial
transition. Particularly prominent are warming
pulses beginning at 14,600 14C yr BP (Oldest
Dryas) and 12,700-13,00014C yr BP (near the Oldest Dryas-Boiling transition) that together terminated the LGM and brought large regions in both
hemispheres to near-interglacial warmth in about
1600 14Cyr.
This synchrony of atmospheric temperature
changes during both the LGM and the glacialinterglacial transition suggests that shifts of ocean
heat across the equator from changing North Atlantic ventilation power (Crowley 1992) was not
the major control on interhemispheric atmospheric
signals. Rather it implicates changes in greenhouse
gas, most likely in the inventory of atmospheric
water vapor, as the dominant climate linkage between the hemispheres.
A conspicuous anomaly in the Greenland icecore records may afford the clue for the cause of the
two climate warming pulses that dominated the last
termination. Namely, there is little or no indication
in these ice-core records of the Oldest Dryas warmGeografiska Annaler - 81 A (1999) · 2
ing pulse so evident elsewhere. We suggest that this
was because of a massive collapse of marine icesheet sectors from North Atlantic continental
shelves during the Heinrich 1 event of Oldest Dryas
age. We speculate that this massive collapse was
the key connection between these two (Oldest
Dryas and Oldest Dryas/B61iing) warming pulses,
suggested by our Southern Hemisphere data to
have been interhemispheric events. The timing and
magnitude of the Oldest Dryas warming pulse
make it a prime candidate for triggering the massive Heinrich 1 ice collapse, which so decreased
North Atlantic salinity that downwelling was suppressed to its lowest level of the last glacial cycle
(Sarnthein et al. 1994), thereby depressing North
Atlantic sea-surface temperatures and maintaining
regional cooling of Greenland. We speculate further that a strong rise in the salinity of the northern
North Atlantic Ocean occurred at the end of the collapse, because the reservoir of glacial ice on continental shelves, as well as in feeder ice streams,
was so depleted that the influx of icebergs dropped
precipitously. Together with rising summer insolation in the Northern Hemisphere, such a strong salinity jump could have switched North Atlantic
thermohaline circulation into its interglacial mode,
with a vigorous LNDW component formed from
inflow of warm and salty water into the Nordic
Seas. The paleoclimate record is the prime clue to
the fundamental role of this thermohaline circulation switch in the last termination, because it shows
coeval warming to near-interglacial values not only
in Chile but even in Europe, despite the fact that
Northern Hemisphere ice sheets were still close to
maximum LGM volumes.
In this view, the decisive thermohaline mode
switch near the Oldest Dryas/B61iing boundary
was the culmination of a series of cycles in which
marine-based ice built up on North Atlantic continental shelves and then collapsed during Heinrich
events. Such buildup-and-collapse cycles of
grounded continental shelf ice could have been a
powerful regulator of North Atlantic near-surface
salinity. Low salinity from a steady influx of icebergs occurred during each buildup phase, the lowest salinity accompanied each Heinrich iceberg
pulse, and a salinity jump occurred after each Heinrich collapse depleted reservoirs of grounded ice
marginal to the North Atlantic Ocean. The salinity
jump after each Heinrich collapse caused increased
conveyor overturn and ocean heat input into the
North Atlantic Ocean. But it is postulated that only
the Heinrich 1 collapse was massive enough to
145
G.H. DENTON ETAL.
drain peripheral ice completely (and hence to cause
a salinity jump big enough to initiate the interglacial mode of North Atlantic thermohaline overturn), because by this time the circum-Atlantic ice
sheets had reached maximum size (and hence vulnerability to unstable collapse from continental
shelves) and also because the Oldest Dryas warming trigger was the most significant, at least as recorded in the Chilean Lake District, since prior to
49,892 14Cyr BP.
Because it is registered so strongly in our Southern Hemisphere data, as well as in some European
paleoclimate repositories, at a time when Greenland ice-core records preclude a major North Atlantic thermohaline switch, the Oldest Dryas
warming pulse is inferred to have originated in the
tropics, perhaps from direct forcing of atmospheric
water-vapor production by half-precession effect.
In contrast, the decisive warming near the Oldest
Dryas/Bolling transition was most likely caused by
the major thermohaline switch recorded so prominently in North Atlantic paleoclimate repositories.
This switch was the most fundamental change in
deep circulation of the world ocean during the last
glacial cycle. The fact that atmospheric warming
was registered as far south as the Chilean Lake District and Isla Grande de Chiloe, and even at Taylor
Dome in peripheral East Antarctica (Steig et al.
1998), strongly suggests a coincident change in the
greenhouse gas content of the atmosphere, most
likely water vapor. Thus our paleoclimate record
from the southern Chilean Lake District suggests
that the major thermohaline switch near the Oldest
Dryas/B611ing transition affected tropical atmosphere/ocean dynamics in such a way that the production of water vapor was reset to near-interglacial values.
Because the middle-latitude Southern Hemisphere and the North Atlantic regional paleoclimate records show key similarities, we stress tropical and North Atlantic triggers for major climate
changes during the last glacial-interglacial transition. But there is a major problem with this approach. Namely, the Byrd ice core in central West
Antarctica, along with some Southern Ocean marine cores, suggest that Southern Hemisphere
warming preceded that in the Northern Hemisphere by several thousand years at the beginning
of the last termination (Hays et al. 1976; Sowers
and Bender 1995). Also the Vostok and Byrd icecore records imply that the B611ing and Younger
Dryas oscillations were out of phase between the
two hemispheres (Blunier et al. 1997). The South146
em Hemisphere lead at climate transitions has been
attributed to varying ocean heat transport across
the equator linked to changing strength of North
Atlantic Deep Water production (Crowley 1992;
Blunier et al. 1998), to local insolation forcing of
sea ice on the Southern Ocean (Kim et al. 1998), or
to a bipolar seesaw of thermohaline circulation
(Broecker 1998).
There are two unsolved stratigraphic problems
in identifying interhemispheric climate linkages.
The first is that a new ice-core record from the peripheral Taylor Dome of the East Antarctic Ice
Sheet shows a climate signal very close to that of
the Greenland ice cores, implying interhemispheric synchrony in atmospheric signals (Steig et al.
1998), probably facilitated by rapid thermohaline
linkage of polar seas (Weyl 1968). The reason for
such fundamental differences between the Taylor
Dome and Byrd isotope records is unknown. The
second problem is that the key subantarcticmarine
cores that suggest a Southern Hemisphere lead of
sea-surface temperaturesare at nearly the same latitudes as the terrestrialrecords in the southern Chilean Lake District and New Zealand discussed here.
Again, the reason for the perceived differences between the two data sets is unknown.
A possible resolution of these problems can
come from detailed AMS radiocarbon chronologies of marine and terrestrial records from a proposed Antarctic CircumpolarTransect (ACT) at the
latitude of subantarctic waters (where abundantforaminifers allow extensive dating). This transect
would cross key land masses, including southern
South America and New Zealand, to permit detailed comparison of marine and terrestrialrecords.
The results from Pacific, Atlantic, and Indian
Ocean sectors should show if abrupt climate
changes during the last glacial/interglacial transition were synchronous between the hemispheres, if
there was a lead in Southern Hemisphere climate,
or if there was a mosaic of NADW outcrops at the
surface of the Southern Ocean that caused mixed
climate signals in Antarctic ice cores.
Acknowledgements
The research was supported by the Office of Climate Dynamics of the National Science Foundation, the Lamont-Scripps Consortium for Climate
Research of the National Oceanic andAtmospheric
Administration, the National Geographic Society,
the Norwegian Research Council, and the Swiss
National Science Fund. P.I. Moreno was supported
Geografiska Annaler · 81 A (1999) · 2
INTERHEMISPHERIC LINKAGE OF PALEOCLIMATE DURING THE LAST GLACIATION
by the EPSCoRProgramof the NationalScience
Foundation.The overallprojectwas carriedout in
co-operationwith the Servicio Nacional de Geologia y Mineria, Santiago, Chile. We are very
gratefulto W.S.Broeckerforhis continuedsupport
andfor manydiscussionson problemsof paleoclimate;he initiatedthis workin 1989 by invitingG.
Denton,C. Heusser,andL. Heusseron a field trip
to southernSouthAmericafundedby Exxon. We
thankGerardBondfor manydiscussionsaboutthe
NorthAtlanticpaleoclimaterecord.ArturoHauser
Y., JorgeMufiozB., and Hugo MorenoR. helped
enormouslywithgeologicaladviceandwithlogistics in the southernChileanLakeDistrict.G. Dentonis gratefulto Nick Shackletonforprovidingthe
basic isotope data for TR163-31B reproducedin
Fig. 5. WethankW.Beck, G. Bonani,C. Eastoe,S.
Gulliksen,I. Hajdas,A. Hogg, T. Jull,R. Kalin,R.
Nydal, and M. Stuiverfor providingradiocarbon
dates. RichardKelly draftedthe figures, and D.
Seymourtyped the manuscript.George Jacobson
aidedgreatlyin scientificdiscussions,andalso edited the manuscript.We thankJohnSplettstoesser
for editingthe manuscript.The reviewsof Chalmers Clappertonand David Sugden improvedthe
manuscriptextensively.This paper is funded in
part by a grants/cooperativeagreementfrom the
National Oceanic and AtmosphericAdministration. The views expressedhereinare those of the
author(s)anddo notnecessarilyreflecttheviews of
NOAA or any of its sub-agencies.
George H. Denton, Department of Geological Sciences and Institutefor Quaternary Studies, Bryand
Global Sciences Center, University of Maine, Orono, Maine 04469, USA.
Thomas V. Lowell, Department of Geology, University of Cincinnati, Cincinnati, Ohio 45221,
USA.
Calvin J. Heusser, 100 Clinton Road, Tuxedo, New
York10987, USA.
Patricio I. Moreno, Institutefor Quaternary Studies, Bryand Global Sciences Center, University of
Maine, Orono, Maine 04469, USA.
Bj0rn G. Andersen, Institute for Geology, University of Oslo, NO-0316, Blindern, Oslo 3, Norway.
Linda E. Heusser, Lamont-Doherty Earth Observatory, Palisades, New York10964, USA.
Geografiska Annaler * 81 A (1999) · 2
Christian Schliichter, Institute of Geology, University of Bern, 3012 Bern, Switzerland.
David R. Marchant, Department of Geology, Boston University, 675 Commonwealth Avenue, Boston, Massachusetts 02215, USA.
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