Quaternary Science Reviews 19 (2000) 1723}1747
Ecosystem response to Lateglacial and early Holocene climate
oscillations in the Great Lakes region of North America
Zicheng Yu *
Department of Botany, University of Toronto, Toronto, Ont., Canada M5S 3B2
Centre for Biodiversity and Conservation Biology, Royal Ontario Museum, Toronto, Ont., Canada M5S 2C6
Abstract
Fossil pollen, plant macrofossils, gastropods, and elemental and stable-isotope geochemistry in a sediment core from Twiss Marl
Pond, southern Ontario, Canada, were used to document climate oscillations during the Last Glacial}Interglacial transition
(&13,000}8500 C BP) and understand their ecological e!ects. Chronology was provided by AMS C dating and regional pollen
correlation. Oxygen isotope (dO) results from mollusc shells, Chara-encrustations and bulk carbonates show a classic climate
sequence of a warm B+lling}Aller+d (BOA) at &12,500}10,920 C BP, a cold Younger Dryas (YD) at 10,920}10,000 C BP, the
Holocene warming at 10,000 C BP, a brief Preboreal Oscillation (PB) at 9650 C BP, and a possible Gerzensee/Killarney (G/K)
cooling shortly before 11,000 C BP.
Clay sediments at the base of the core contain high herb and shrub pollen and abundant arctic/alpine plant macrofossils, indicating
a treeless tundra with severe soil erosion in watershed. During the BOA warm period, authigenic marl began to be deposited, and
Picea woodland became established. The establishment of Picea woodland after peaks of dO and of carbonate accumulation
suggests a lagged response of upland vegetation to BOA warming. In contrast, the occurrence of warmth-loving aquatics Najas yexilis
and Typha latifolia at that time indicates sensitive responses of aquatic plants. The YD cooling is indicated by a &1.5 negative
excursion in dO, an increase in minerogenic matter and higher concentrations of erosion-derived elements (Al, Na, K, Ti and V).
Pollen data show no forest transformation in response to YD cooling, which is attributed to the insensitive nonecotonal vegetation at
that time. However, more openings in the forests and increased erosion in the watershed are indicated by a slight increase of herb
pollen, high concentrations of erosion elements and a Pediastrum peak. The onset of the Holocene was marked by an abrupt increase
of 2 in dO and the replacement of Picea woodland by Pinus-dominated forest. The Picea recurrence at 9650 C BP demonstrates
sensitive response of ecotonal vegetation to the PB climate oscillation, which is also indicated by 0.4 negative excursion of dO.
These new results suggest the importance of multiproxy records for reliable paleoclimate reconstruction.
Reevaluation and revised chronologies of previously published sites (Gage Street, and Nichols Brook) in the eastern Great Lakes
region show their major dO shifts correlative to the YD and PB oscillations as documented from Twiss Marl Pond and nearby
Crawford Lake. The sequence and magnitude of climatic oscillations from these sites match in detail with records from the Atlantic
Seaboard, suggesting that these oscillations are an expression of broad-scale, probably global, climate change rather than local
meltwater-induced climate cooling. 2000 Elsevier Science Ltd. All rights reserved.
1. Introduction
The Last Glacial}Interglacial transition (&13,000}8500
C BP) experienced several broad-scale climatic oscillations in both Europe and North America (e.g., Lotter et
al., 1992; Levesque et al., 1993; BjoK rck et al., 1996; Yu and
Eicher, 1998), including the Younger Dryas (YD), Preboreal Oscillation (PB) and pre-YD Gerzensee/Killarney
* Correspondence address: Canadian Forest Service, 5320 122 Street,
Edmonton, Alberta, Canada T6H 3S5. Tel.: (780) 435-7304; fax: (780)
435-7359.
E-mail address: [email protected] (Z. Yu).
Oscillations (G/K). The YD event is an abrupt and very
marked cooling episode between &11,000 and 10,000
C BP. In the Great Lakes region, there are two hypotheses to explain the YD-age climatic reversal, (1)
global climatic oscillation caused by change in oceanic
and atmospheric circulations (e.g., Rind et al., 1986;
Shane, 1987; Wright, 1989; Peteet et al., 1997; Yu and
Eicher, 1998) and (2) regional cooling caused by cold
proglacial lake e!ects (e.g., Lewis and Anderson, 1992;
Rea et al., 1994b). Thus, establishing the geographic
extent and relative magnitude of these climate oscillations related to di!erent altitudes and proximity to proglacial lakes and meltwater spillways is essential for
0277-3791/00/$ - see front matter 2000 Elsevier Science Ltd. All rights reserved.
PII: S 0 2 7 7 - 3 7 9 1 ( 0 0 ) 0 0 0 8 0 - 9
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Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
testing these hypotheses in particular and for understanding the mechanisms and causes of abrupt climatic
changes in general.
In this study, a multi-proxy data approach was employed to investigate changes in climate, landscape, vegetation, and lake biota during the Lateglacial and early
Holocene from a small pond at the edge of Ontario's
Niagara Escarpment. Oxygen isotopes from bulk carbonates (lake marl), Chara-encrustations and mollusc
(Valvata sincera and Pisidium ferrugineum) shells were
used as a proxy for paleotemperatures. Lithology and
elemental geochemistry were used to infer lake and
watershed conditions, which can provide insights into
soil development and soil}vegetation interactions within
lake watershed (Mackereth, 1966; Engstrom and Wright,
1984; MacDonald et al., 1993). Fossil pollen and plant
macrofossils provided evidence for regional vegetation
and also local/watershed #ora. Freshwater gastropods,
aquatic plant fossils and the green alga Pediastrum were
used as lake biota indicators. Chronology was controlled
by three accepted AMS C dates on terrestrial plant
macrofossils from this site and by dated regional pollen
correlation. Part of the data from this site, especially
oxygen isotopes from mollusc shells, has been presented
and brie#y discussed in a short report (Yu and Eicher,
1998). Here I provide a more detailed discussion of
additional proxy data and review other published records. The objectives of this paper are: (1) to present
a detailed multiproxy paleoecological record for the
Lateglacial and early Holocene; (2) to document expression of climatic oscillations, including YD, PB and G/K
events, in the eastern Great Lakes region; (3) to understand di!erential responses of upland vegetation and
aquatic biota to climate changes (as inferred from oxygen
isotopes); and (4) to review published paleoclimate records from the region and evaluate causes of these climate changes.
The Lateglacial is a prolonged period of slow and
hesitant amelioration of climate, characterised by extreme instability of climate between the Last Glacial
maximum (LGM) of 20,000}18,000 C BP and rapid
Holocene warming at 10,000 C BP. The Lateglacial
here refers to the time period of &13,000}10,000 C BP,
which includes classic European chronozones B+lling
(13,000}12,000 C BP), Aller+d (11,800}11,000 C BP)
and YD (11,000}10,000 C BP) (Mangerud et al., 1974).
Because evidence for the Older Dryas (OD) is not identi"ed in most records, a single warm episode (the B+llingAller+d; BOA) is used here as proposed by Broecker
(1992). Preboreal is the earliest period of the Holocene
(10,000}9000 C BP). The radiocarbon timescale will
be used whenever possible to avoid potential confusion from inconsistent use of the stratigraphical
nomenclature of Mangerud et al. (1974) and to facilitate
systematic regional correlation in the future (Lowe,
1994).
2. Study site
Twiss Marl Pond (informal name; 4332645N,
7935715W; 256 m asl altitude) is located about 60 km
southwest of Toronto at the edge of the Niagara Escarpment in southern Ontario (Fig. 1). The area was deglaciated about 13,000 C BP, and the site is mapped as
outwash gravel deposits (Karrow, 1987). The Quaternary
deposits consist of coarse-grained materials (boulder,
gravel and sand) of Wentworth Till and "ne-grained
materials (sand, silt and clay) of Halton Till, with extensive dolomite bedrock outcrops (Karrow, 1987).
The continental climate in southern Ontario is modi"ed by the Great Lakes and is characterised by short
warm summers, long cold winters, and evenly distributed
precipitation. There are no abrupt climate gradients in
the region, although mean annual temperatures decrease
from southwest to northeast, and precipitation patterns
re#ect topography and proximity to the Great Lakes.
The study area has a mean annual temperature of about
73C and annual precipitation of 850 mm (Environment
Canada, 1982). There is slightly more precipitation near
the Escarpment than on either side, possibly an orographic e!ect (Yu, 1997a). The nearest isotope station at
Simcoe has a weighted mean dO value of !9.27
relative to standard mean ocean water (SMOW; !10.26
for arithmetic mean) in precipitation (IAEA/WMO
GNIP database), and Simcoe has a higher annual temperature (8.63C) and higher precipitation (941 mm) because of its more southern and western location. The
study area is on the boundary between the Great
Lakes}St. Lawrence mixed forest and deciduous forest
regions (Rowe, 1972). Bedrock outcrops support forests
dominated by eastern white cedar (Thuja occidentalis);
the postglacial history of white cedar along the Niagara
Escarpment has been documented using detailed pollen,
stomata and macrofossils from nearby Crawford Lake
(Yu, 1997b).
Twiss Marl Pond (TMP) occurs in a swamp on a tributary of Bronte Creek (Fig. 1C). The original pond has
been "lled since about 7000 C BP by accumulated marl
(Yu, 1997a; Table 1). Presently, three man-made ponds
are connected by a small waterway in the northwest;
these ponds were excavated in the 1920s during mining of
marl deposits for manufacture of concrete sewer pipes
and culvert tiles (Guillet, 1969). The coring site was
located on the residual berm between two ponds in the
southwest end (Figs. 1C and 2).
3. Methods
3.1. Field sediment coring
Two sediment cores of 5 cm in diameter were taken
from the berm between two ponds in January and
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
1725
Fig. 1. Maps showing location of Twiss Marl Pond and other sites mentioned in the text. (A) Location of Twiss Marl Pond in North America, ice cores
(GISP2, GRIP) from Greenland Summit and Gerzensee in Switzerland. R, Rattle Lake (BjoK rck, 1985) and S, sites in the southern Great Lakes region
(Shane, 1987; Shane and Anderson, 1993). (B) Map showing Twiss Marl Pond (TMP) along Ontario's Niagara Escarpment and other sites. Crawford
Lake (Yu, 1997a; Yu and Eicher, 1998); Rostock site (McAndrews and Jackson, 1988); Gage St. site (Anderson, 1982; Schwert et al., 1985; Fritz et al.,
1987); Nichols Brook site in New York (Calkin and McAndrews, 1980; Fritz et al., 1987). (C) Detailed map showing location of Twiss Marl Pond in
relation to Crawford Lake along the Niagara Escarpment. Note that the two sites are not hydraulically connected.
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Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
Table 1
AMS radiocarbon (C) dates from Twiss Marl Pond, Ontario, Canada
C date (BP $1 SD) Calendar age (cal BP)
Sample depth (cm) Material dated
Carbon (mg)
Lab number
170}190
390}400
2.44
0.15
TO-5834
TO-5504
7900 ⴞ 70
8750$200
8630
0.80
1.27
0.63
0.14
0.41
0.24
TO-6106
TO-5505
TO-5506
TO-5835
TO-6107
TO-5836
9670 ⴞ 100
10,920 ⴞ 80
10,550$140
10,250$250
7810$140
10,190$160
10,924
12,845
400}415
480}490
510}520
560}570
570}590
600}610
30 Pinus strobus needles/fragments
3 Larix needles, 1 Betula bract and
unidenti"ed charcoal fragments
20 Picea and 40 Larix needle fragments
9 Picea needles and 1 Picea seed
4 Picea needles and 1 Larix needle
30 Dryas and 26 Salix leaf fragments
90 Dryas and 15 Salix leaf fragments
27 Dryas leaves, 60 Dryas leaf fragments
and 4 Salix leaves
TO: IsoTrace Laboratory at the University of Toronto, Canada.
All dates were corrected for natural and sputtering fractionation to a base of dC"!25.The dates are uncalibrated conventional radiocarbon
dates in years before present (BP), using Libby C meanlife of 8033 years.The errors represent 68.3% con"dence limits. SD: standard deviation.
Calendar age calibrated using the program CALIB Ver.3.0.3 (Stuiver and Reimer, 1993).
Rejected as too young. This sample was badly preserved and yielded very little datable carbon.Probably biased by charcoal fragments or by
contaminants.
Limited pretreatment of 4 h for TO-6106 and 2 h for TO-6107.The date of TO-6107 is rejected as too young.
Rejected as too young. Perhaps the Larix needle was intrusive.
Rejected as too young. These two samples were very small after pretreatments and produced little datable carbon.
November, 1995, respectively, with a modi"ed Livingstone piston corer (Wright, 1967). The sediments were
described, and the 1-m long core segments were wrapped
in plastic wrap and aluminum foil in the "eld. The two
cores correlated well by distinct sediment beds. The "rst
core was used for all analyses, with some supplementary
intervals from the second core for plant macrofossils,
mollusc shells and Chara-encrustations to "ll the gaps
between core segments in the "rst core.
3.2. AMS radiocarbon dating
Eight terrestrial plant macrofossil samples were submitted to the IsoTrace laboratory at the University of
Toronto for AMS C dating (Table 1). All ages used
in this paper are uncalibrated C ages, unless stated
otherwise.
3.3. Loss-on-ignition and elemental geochemistry
analysis
The organic matter and carbonate content of sediments were estimated with loss-on-ignition analysis
(Dean, 1974), as dry weight loss at 5503C for 1 h and
10003C for 1 h, respectively. The residue after 10003C was
called silicates, which include all mineral matter except
carbonates. Elemental concentrations were determined
with instrumental neutron activation analysis (INAA) of
bulk sediment (Muecke, 1980) in the SLOWPOKE Reactor Facility at the University of Toronto. All the dry bulk
samples (0.3}1 g) were irradiated for 5 min at a #ux of
1;10 neutrons/cm/s (2 kW). The irradiated samples
were counted for 5 min after a 10}20 min delay for short-
lived isotopes (Ca, Mg, Al, Na, Mn, Ti and V) and after
&10 h delay for long-lived isotope (K). Calibration of
elemental concentrations was based on SLOWPOKE
standards (R. Hancock, pers. comm., 1995).
3.4. Stable isotope analysis
Bulk carbonate (marl) samples were dried at 503C
overnight, and macroscopic plant remains and mollusc
shell fragments were picked out under a microscope
(U. Eicher, University of Bern, per. comm., 1996). Valvata
sincera and Pisidium ferrugineum shells and Charaencrustations were picked during macrofossil analysis.
Valvata sincera and P. ferrugineum are mollusc species
that are found in a non-marginal environment in lake
center and thus are likely to have representative isotopic
composition of lake water (Fritz, 1983). The analytical
procedure and methodology are similar to that used for
deep-sea sediment (McCrea, 1950). Because erroneous
results have been observed from pretreated samples
either by 4503C heating or chemical oxidation methods,
the bulk carbonate samples were prepared without pretreatment (U. Eicher, University of Bern, pers. comm.,
1996). Mollusc shells were pretreated with KClO solution. Each sample of 5}40 mg of carbonate was reacted
with 95% phosphoric acid (H PO ) at a constant
temperature of 503C for 1 h to produce CO . The CO
was then analyzed for its O/O and C/C ratios
with a mass spectrometer (Finnigan MAT 250) at the
University of Bern, Switzerland (see Siegenthaler and
Eicher, 1986). The results were presented as conventional
delta notation (d), which is de"ned as [(R
}R
)/
R
];1000 (where R is the absolute ratio of O/O
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
1727
number of Lycopodium spores (batch C710961,
mean"13,911$306 spores per tablet) was initially
added to each sample to estimate pollen concentration
(Benningho!, 1962; Maher, 1981). The concentrate
was stained with safranine and mounted in silicone oil.
Each pollen sample was counted under a light microscope at ;400 magni"cation and at ;1000 magni"cation for critical determinations in regularly spaced
transverses to mostly 300 terrestrial pollen grains; at
least "ve transverses were counted for pollen-rich samples and the whole slide counted for pollen-poor samples
(in all samples pollen sum greater than 200 grains).
Identi"cations followed McAndrews et al. (1973) aided
by the modern reference collection in J.H. McAndrews'
palynology laboratory at the Royal Ontario Museum
(ROM).
3.6. Plant and gastropod macrofossil analysis
During the subsampling for macrofossil analysis the
outer 1}2 mm of each sample was scraped o! to
remove possible contaminant macrofossils. All non-horizontally oriented plant macrofossils on or near the sediment surface were also removed for the same purpose.
Samples of 10-cm long sediment core segments (about
200 ml) were dispersed by gentle agitation in water and
then washed through 500-lm mesh to concentrate macrofossils. Identi"able materials were picked under
a stereomicroscope at ;8 magni"cation. Seeds and
leaves were identi"ed by comparison with a modern
reference collection at the ROM. For gastropod analysis,
the fraction '500 lm was air-dried and then all the
mollusc shells were picked out by #oating in water.
Gastropods were identi"ed using Clarke (1981) and La
Rocque (1970).
4. Results
4.1. Radiocarbon dates and chronology
Fig. 2. Ground photo (upper) and air photo (lower) showing the coring
site at Twiss Marl Pond, Ontario. Scale of air photo is approximately
1 : 13,000. Miller Lake is also an uno$cial name; a pollen diagram from
its sediments was presented in McAndrews (1994).
or C/C, and the Vienna-PDB (Pee Dee belemnite) is
the standard). The analytical precision is $0.02 for
both dO and dC.
3.5. Pollen analytical method
Pollen samples of 0.8 ml (or 1.6 ml for the basal clay
samples) were taken at mostly 10-cm intervals. Pollen
was concentrated with a modi"ed acetolysis procedure
(Fvgri and Iversen, 1989) and "ne sieving to remove
clay-sized particles (Cwynar et al., 1979). A known
Three of the eight AMS C dates on terrestrial plant
macrofossils (Table 1), together with ages estimated from
regional pollen correlation (Table 2), provided chronological control. The age-depth plot (Fig. 3) was used as
the age model for the subsequent discussion of
paleoecological records. Five AMS dates were rejected as
too young, mostly because the samples were small and
yielded very little datable carbon after pretreatments (see
notes in Table 1). As stated in the IsoTrace radiocarbon
analysis reports, small samples were likely biased by
contaminants, in addition to having a larger standard
deviation. The chronology is regarded as tentative and
may be in error, especially before 11,000 C BP.
The bottom of the core was assigned an age of 13,000
C BP based on the local deglaciation history (Barnett,
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Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
Table 2
Tentative chronology for stratigraphic records at Twiss Marl Pond,
Ontario, Canada
Paleoecological
event
C Age (BP) Remark
Basal age; local
deglaciation
&13,000
Estimated from local deglaciation
and Lake Whittlesey history
(12,920$400, W430; 12,900$
200, I-3175; 12,800$250, Y240;
Barnett, 1979); extrapolated from
other pollen sites (Winter Gulf:
12,610$200, I-8022; 12,730$
220, I-3665; Calkin and McAndrews, 1980; Rostock: 12,550$
120, WAT-1511; Pilny et al., 1987;
McAndrews and Jackson, 1988)
Low boundary
of Picea pollen
zone; a!orestation
&12,000
Estimated from Maplehurst Lake
(12,020$180, GSC-1156; Mott
and Farley-Gill, 1978), Brampton
site (11,970$150, BGS-550;
12,300$360, BGS-551; Terasmae and Matthews, 1980) and
Decoy Lake (11,770$90, TO1048; Szeicz and MacDonald,
1991)
Low boundary
of Younger Dryas
(YD) event
10,920
AMS date from Twiss Marl Pond
(10,920$80, TO-5505; Table 1)
Picea}Pinus
transition;
Holocene warming;
upper boundary
of YD event
10,000
Extrapolated/interpolated from
AMS
dates
of
Crawford
(9910 BP) and Twiss sites; pollen-dated (10,330 BP: Bennett,
1987; 10,070 BP: Szeicz and MacDonald, 1991)
Preboreal
Oscillation
9650
AMS dates of 9620$60 (345}347
cm; CAMS-16062) and 9670$70
(355}357 cm; CAMS-16061)
from Crawford Lake core D
Date on Picea needles and seed;
Dates on Larix needles (CAMS-16062) and Larix needles, wood
fragments, charcoal fragments and Compositae seed (CAMS-16061).
1979; Karrow, 1987) and on extrapolation from dated
pollen sites (Calkin and McAndrews, 1980; Pilny et al.,
1987; McAndrews and Jackson, 1988). The shrub/
herb-Picea transition (i.e., the lower boundary of
Picea pollen zone) was at an age of 12,000 C
BP, estimated from nearby dated sites (Mott and FarleyGill, 1978; Terasmae and Matthews, 1980; Calkin
and McAndrews, 1980; Szeicz and MacDonald, 1991;
see Table 2 for list of dates). The age of the PiceaPinus pollen transition was extrapolated as 9910 C
BP from Crawford Lake based on two AMS dates and
a pollen date of 4800 C BP for the mid-Holocene
hemlock decline (Yu, 1997a). The interpolated ages
for this transition are 10,330 and 10,070 C BP,
Fig. 3. The age-depth plot and age model for Twiss Marl Pond. Crosses
are AMS C dates (with sampling intervals and standard deviations),
open squares are pollen-inferred dates of 9650 BP for Picea recurrence
(AMS dates from Crawford Lake), 12,000 BP for the tundra-Picea
transition (estimated from Mott and Farley-Gill, 1978; Terasmae and
Matthews, 1980; Szeicz and MacDonald, 1991) and 13,000 BP for the
age of local deglaciation. Crossed squares are rejected AMS dates (see
Table 1 for details). PAZ: pollen assemblage zones.
respectively from Hams Lake (bulk sediment dates; Bennett, 1987) and from Decoy Lake (AMS dates; Szeicz
and MacDonald, 1991), both sites about 50 km west of
Twiss Marl Pond. Based on these AMS dates, the
Picea}Pinus transition in this region is very close to
10,000 C BP.
4.2. Sediment lithology and elemental geochemistry
The core is 630 cm long (Fig. 4). This paper presents
the results below 250 cm, which covered the Lateglacial
and early Holocene. The basal 60 cm (from 630 to
570 cm) is reddish and grey clay, and the lower reddish
clay could have been deposited by Glacial Lake Whittlesey (see Karrow, 1987). From 570 to 525 cm is
transitional clayey marl. The sediment from 525 to
250 cm is pond marl with 80}90% carbonate. Within the
marl unit there are silicate peaks at 470 cm and 430 cm.
Elemental geochemistry shows patterns similar to LOI
results but with more detail. The concentrations of
Al, Na, K, Mg, Ti, V and Mn were high in the clay and
all declined during the transitional clayey marl. The
two silicate peaks above the transition correspond to
high concentrations of these elements. A small peak
of Al and K at 500 cm was also evident. The increase
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
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Fig. 4. Lithology, sediment composition estimated from loss-on-ignition analysis and element concentrations from INAA for Twiss Marl Pond.
1 ka"1000 years; PAZ: pollen assemblage zones.
in Mn at 360 cm was the only elemental change in the
upper marl.
4.3. Oxygen isotopes
The dO and dC pro"les were divided into three
zones based on major shifts in oxygen isotopes (Figs. 5A
and B). The dO values are systematically &1}2
higher from mollusc shells (Pisidium ferrugineum and
Valvata sincera) than from Chara-encrustations and bulk
carbonate samples (Fig. 5A). dO ranges from !7.7 to
!11.4 for mollusc shells and from !7.8 to !12.5
for bulk samples.
Zone TI-1 (560}490 cm; &12,500}10,920 C BP).
The dO values are !9.5 to !10.3 from mollusc
shells and !10.3 to !11.8 from encrustations. The
encrustations provided the earliest reliable dO value at
550}560 cm for this site. The basal four bulk samples
from 630 to 540 cm have high dO values of !6.2 to
!8.7, which were from the combined signal of
authigenic calcite and allogenic detrital dolomite, as suggested by an o!set between the bulk and encrustation
samples and a high dO value of !6.46 analyzed on
nearby dolomite bedrock (average of two measurements
of !6.55 and !6.43). From 530 cm upward, bulk
samples provided unequivocal dO values of authigenic
calcite, indicated by close agreement between bulk and
encrustation samples. A slight negative shift (up to 0.3)
in dO from shells and marl occurred at 500}510 cm.
Zone TI-2 (490}390 cm; 10,920}9900 C BP). dO in
mollusc shells (mostly V. sincera) is from !11.4 to
!10.2, which declined abruptly to the lowest value at
the beginning of this zone and increased step-wise afterwards, with evident #uctuations. The close agreement
between V. sincera and P. ferrugineum at three levels
suggests that species vital e!ects are unimportant.
Chara-encrustations and bulk samples showed identical
dO values from !12.7 to !10.9.
Zone TI-3 (390}250 cm; 9900}8500 C BP). An
abrupt positive shift in dO occurs at the beginning of
this zone from !10.2 to !9.2 for mollusc samples,
and it gradually increased to !7.8 at the top of the
record. The dO values from bulk samples showed similar changes. A minor negative excursion of 0.4 was
evident from both shells and bulk samples at 370}380 cm
(9650 C BP).
4.4. Carbon isotopes
The dC values are &2.5 to 4 lower from mollusc
shells than from encrustations and marl samples
(Fig. 5B). dC in the pro"le ranges from!6.9 to!9.2
for shells, from 0 to !6.3 for encrustations, and from
#0.7 to !5.6 for marl samples.
Zone TI-1 (560}490 cm; &12,500}10,920 BP). The
dC values from Pisidium shells gradually decreased
from !6.9 to !8.3. Encrustations and bulk marl
samples also showed the decrease but have a larger
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Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
Fig. 5. (A) Oxygen isotope (dO) pro"les from Twiss Marl Pond. (B) Carbon isotope (dC) pro"les from Twiss Marl Pond. Di!erent curves and bars
are from measurements on various carbonate materials (two connected red bars: shells of two mollusc species; blue bars: Chara-encrustations, and
green curves: bulk marl carbonates). A few overlapping samples from two mollusc species show similar isotope values, suggesting that vital e!ects are
unimportant. Encrustation and bulk marl samples have almost identical dO values, but encrustation samples have slight negative dC values than
marl. PB, Preboreal Oscillation; G/K, Gerzensee/Kilarney Oscillations. V-PDB, Vienna Pee Dee belemnite.
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
change from 0 to !4.1 and from !0.9 to !3.0,
respectively. The basal four bulk samples from 630 to
540 cm have high dC values of #0.6 to !1, which
were from the combined signal of authigenic calcite and
allogenic detrital dolomite, as suggested by high dolomite bedrock dC value of #3.89 (average of two
measurements of #3.87 and #3.92).
Zone TI-2 (490}390 cm; 10,920}9900 BP). After an initial increase, dC in Valvata shells showed a gradual
decrease from !7.8 to !9.2. Similar decreases from
encrustations and marl samples occurred from !4.2 to
!6.3 and from !3.0 to !5.6, respectively.
Zone TI-3 (390}250 cm; 9900}8500 BP). Valvata shells
showed an abrupt positive shift in dC at the beginning
of this zone, and subsequently remained constant at
about !7.3. Marl samples also showed this increase
and declined to the top of the record.
4.5. Pollen stratigraphy
The Lateglacial and early-Holocene pollen sequence
was divided into three local pollen assemblage zones
(PAZ) with subzones, based on stratigraphically constrained cluster analysis (CONISS; Grimm, 1987). Only
selected pollen taxa were shown in the percentage pollen
diagram (Fig. 6).
Zone TP-1 includes two subzones. Subzone TP-1a:
Picea}Pinus}Cyperaceae}Alnus}Thuja/Juniperus}
Dryas (630}570 cm; &13,000}12,500 C BP). This basal subzone was dominated by Picea (30%) and Pinus
1731
(20%), shrubs Alnus, Salix and Thuja/Juniperus (likely
Juniperus) (collectively 10}12%) and Cyperaceae (10%)
and other herbs (collectively 10%). Populus (up to 5%)
might be locally present but other tree pollen types such
as Picea, Pinus, Ulmus, Ostrya/Carpinus and Fraxinus are
probably attributable to long-distance transport of pollen grains from the south and west. Pollen concentration
was very low at 5000}8000 grains/ml. Subzone TP-1b:
Salix-Cyperaceae-Artemisia-Thuja/Juniperus-Picea
(570}515 cm; &12,500}12,000 C BP). In this subzone
Picea and Pinus decreased, and Salix replaced Alnus as
the predominant shrub. Thuja/Juniperus peaks at 8%.
Cyperaceae and Artemisia increased sharply and remained the dominant herbs, together with Tubuli#orae
(undi!.) near the bottom. Pediastrum colonies appeared
in this subzone ((5%). Pollen concentration slightly
increased to 10,000}15,000 grains/ml.
Zone TP-2 includes two subzones. Subzone TP-2a:
Picea}Cyperaceae (515}415 cm; 12,000}10,100 C BP).
This subzone was dominated by Picea pollen (up to 85%)
with some Larix, Thuja/Juniperus and Cyperaceae. Herbs
and shrubs (Cyperaceae, Poaceae, Artemisia, Salix and
Alnus) all declined. The top of this subzone showed an
abrupt increase in Pediastrum colonies (up to 25%).
Pollen concentration increased to 30,000 grains/ml. Subzone TP-2b: Picea}Pinus transition (415}355 cm;
10,100}9500 C BP). In this subzone Picea decreased
from 50% to 5%, and Pinus increased to 40% with an
accompanied increase of deciduous tree pollen taxa such
as Betula, Fraxinus, Quercus, Ulmus and Ostrya/Carpinus,
Fig. 6. Summary percentage pollen diagram and inferred vegetation at Twiss Marl Pond. Zones were suggested by CONISS results.
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whereas Cyperaceae, Artemisia and Ambrosia decline.
A distinct feature of this transitional zone is the return of
Populus, Alnus, Salix and Corylus and persistence of Larix.
Pediastrum persisted at 10%. Pollen concentration declines to 20,000 grains/ml.
Zone TP-3: Pinus}Betula}Quercus (355}250 cm;
9500}8500 C BP). Pinus became dominant (up to 70%)
with increases in most deciduous tree pollen taxa}Betula,
Quercus, Ulmus and Ostrya/Carpinus. Virtually all shrubs,
including Thuja/Juniperus, herbs, and Pediastrum disappeared. Pollen concentration increased sharply to
60,000 grains/ml.
4.6. Plant-macrofossil diagram
The plant-macrofossil concentration diagram at Twiss
Marl Pond was divided into three local macrofossil assemblage zones, with subzones if necessary (Fig. 7), based
on CONISS (Grimm, 1987).
Zone TM-1: Dryas-Salix-Potamogeton (630-540 cm;
&13,000}12,200 C BP). This zone was dominated by
leaves of Dryas integrifolia and Salix herbacea and their
receptacles, seeds and buds, together with two Silene cf.
acaulis seeds. Aquatics include seeds of Potamogeton and
Najas yexilis, and oospores of Nitella (lower) and Chara
(upper). Also there were many unknown or unidenti"ed
macrofossils in this zone. A fossil fragment at 590}600 cm
appeared similar to a Picea seed, but it cannot be identi"ed with con"dence due to its poor preservation.
Zone TM-2: Picea (540}410 cm; 12,200}10,100
C BP). Picea needles and seeds are most abundant in
this zone, together with some Carex seeds and Ericaceae
leaves. An isolated Larix needle fragment at 510}520 cm
might be displaced from younger sediments, as also suggested by a too young C date with this needle fragment
(see note in Table 1).
Zone TM-3 includes two subzones. Subzone TM-3a:
Larix-Picea (410}340 cm; 10,100}9400 C BP). Larix
needles increased to peak values in this subzone, which
also contains a Larix seed. Picea declined but persisted.
Najas yexilis reappeared at the top of this subzone. Betula
bracts "rst appeared in this subzone. Subzone TM-3b: Betula-Larix-Najas-Pinus (340}250 cm; 9400}8500 C BP).
Betula, including B. papyrifera, shows continuous presence. Najas increased to peak values, whereas Larix declined to modest levels. A Pinus resinosa seed and a P.
strobus needle and pollen cone were found in this subzone.
4.7. Gastropod stratigraphy
The gastropod percentage diagram (Fig. 8) was divided
into four local zones based on major species composition
and abundance changes. No gastropods were found in
the basal clay sediments at 630}580 cm, although
pelecypods (bivalves) were present.
Zone TG-1: Gyraulus}Fossaria zone (580}510 cm).
Gyraulus parvus (30}100%) and Fossaria decampi (up to
70%) are the only species except for one terrestrial Vertigo modesta shell at 520}530 cm. Gyraulus parvus is the
"rst gastropod species to appear at this site. Gastropods
showed a sharp decline in abundance during the
transition to the next zone.
Fig. 7. Plant-macrofossil concentration diagram at Twiss Marl Pond. Zones were suggested by CONISS results.
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
1733
Fig. 8. Gastropod percentage diagram with total concentration at Twiss Marl Pond.
Zone TG-2: Valvata sincera}Fossaria zone (510}470 cm).
In this zone V. sincera appeared and dominated the assemblages (&70%) with some F. decampi and a few G. parvus
and Lymnaea stagnalis.
Zone TG-3: Gyraulus parvus}V. sincera}F. decampi}
Helisoma anceps zone (470}340 cm). Gyraulus parvus returned as the dominant species replacing F. decampi and V.
sincera. Helisoma anceps appeared at the beginning of this
zone and remained at &5%. A sharp decline in gastropod
concentration is evident in the upper part of this zone.
Zone TG-4: Valvata tricarinata}H. anceps}G. parvus
zone (340}250 cm). This zone has the most diverse gastropod assemblage, which includes all the major species.
Valvata tricarinata appeared in this zone and became codominant with G. parvus and H. anceps.
5. Discussion
5.1. Paleoclimatic interpretation of stable isotopes in
freshwater carbonates
The O/O ratios in freshwater carbonates have
been increasingly used in paleoclimatic studies as a proxy
record of continental climate (e.g., Stuiver, 1970; Eicher
and Siegenthaler, 1976; Siegenthaler et al., 1984; Fritz et
al., 1987; Ammann et al., 1994; Goslar et al., 1995; Ahlberg et al., 1996; von Grafenstein et al., 1998, 1999). In
isotopic equilibrium, the O/O ratios in carbonates
depend on the isotopic composition of lake water and on
water temperature (with a coe$cient of !0.24 per 3C).
Evaporative enrichment is the most important hydrological factor modifying the isotopic composition of
lake water. In high- and mid-latitudes, however, the isotopic signal of the water primarily re#ects the atmospheric temperature e!ect, with a coe$cient of 0.65 per
3C between air temperature and dO in precipitation
(Dansgaard, 1964; Siegenthaler and Eicher, 1986;
Rozanski et al., 1992). This is especially true in the Lateglacial and early Holocene, because a strong temperature
gradient prevailed in the study region immediately south
of the continental ice sheets (COHMAP, 1988; Webb et
al., 1993; Levesque et al., 1997). The variation of temperatures along this gradient might be sensitively re#ected in
precipitation dO values, which is unlike the climate
regime during the mid-Holocene when the Paci"c air
mass might come into play with a continental e!ect on
precipitation dO signals (Yu et al., 1997).
At Twiss Marl Pond, there is a close parallel of dO
from mollusc shells, encrustations and marl (Fig. 5A),
which suggests that they were all formed in isotopic
equilibrium with the water. The shells are systematically
enriched in O by &1}2 relative to marl and encrustation samples. This enrichment was due in part to the
mineralogy di!erence of molluscs (aragonite) and marl
(calcite), because aragonite shells are enriched by at least
0.6 over calcite under the same equilibrium conditions
(Tarutani et al., 1969). In addition, environmental di!erences between mollusc habitat and loci of marl precipitation can account for some of this enrichment (Fritz et al.,
1987). The di!erence in dO between carbonate shells
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Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
and marl at TMP is smaller than reported elsewhere (e.g.,
Fritz et al., 1987; Lister, 1988).
The dO records from mollusc shells, encrustations
and marl at TMP show a negative shift of 1.3 at 10,920
C BP (at 490 cm) and a positive shift of up to 2
around &10,000 C BP (at 390 cm) (Figs. 5A and 9).
The more negative intervening interval indicates a cold
period corresponding with the Younger Dryas event as
documented elsewhere (e.g., Eicher, 1980; Dansgaard et
al., 1993; see Fig. 10 for dO pro"les from Greenland
ice-core and Swiss lake sediments). Assuming that the
present carbonate dO}air temperature correlation of
0.41 per 3C (di!erence between 0.65 and !0.24 above)
held and there was no signi"cant change in evaporation
for this period, the 1.3 negative shift in dO at the
beginning of the YD event could be translated to a 33C
decrease in mean annual air temperature, and the 2
positive shift at the beginning of the Holocene to a 53C
increase. However, because these assumptions are not
necessarily valid for long-term climate change, these estimates of temperature changes are considered "rst approximations. The absolute temperature values are more
di$cult to estimate because we have no independent data
to evaluate the local hydrological e!ects, mostly evaporative enrichment (see Siegenthaler et al., 1984; Siegenthaler and Eicher, 1986).
During the YD interval, dO #uctuated, showing
minimum values at "rst and slight increase in the second
half. This #uctuation suggests either that the YD event
was not a homogeneous cold period, or that local factors
could play a role in isotope signals. The variations during
the YD interval were also indicated by elemental concentrations and sediment composition (see Figs. 4 and 9).
Although the strong detrital isotope signals have been
documented for the YD interval at a Danish lake (Hammarlund and Buchardt, 1996), the increased detrital input during this interval could not have contributed to the
dO #uctuations at TMP. This is because (1) mollusc
dO (which is not subject to detrital e!ect from particulate carbonate) shows similar #uctuations, (2) no signi"cant o!sets of dO between encrustations and bulk
carbonates were found, and (3) there was no correspondence between erosion episodes (silicate and erosion-element peaks) and positive (less negative) dO values
(eroded detrital carbonate has less negative dO values).
The similar #uctuations were also recorded in dO pro"les at nearby Crawford Lake (Fig. 10; Yu and Eicher,
1998), which has no hydrologic connection with TMP
(see Fig. 1C), in some European isotope records (Eicher,
1987) and in European paleoecological records (Isarin
and Bohncke, 1999) as well as in GRIP ice core (Fig. 10;
Dansgaard et al., 1993). Thus, the YD interval might
experience #uctuations in temperature and/or precipitation.
After the major positive shift in dO at the beginning
of the Holocene, a minor negative excursion of &0.4
was evident at 9600 C BP (370}380 cm), which correlates with the PB event recorded in nearby Crawford
Lake, the Greenland ice-cores, the North Atlantic Ocean
and central European lakes (Fig. 10; Eicher and Siegenthaler, 1976; Eicher, 1980; Lotter et al., 1992; Lehman
and Keigwin, 1992; Dansgaard et al., 1993; Yu and
Eicher, 1998). This minor oscillation was also indicated
by recurrence of Picea pollen at TMP (Fig. 9) and more
clearly at nearby Crawford Lake (Yu and Eicher, 1998).
Before the YD interval, another slight negative excursion
of dO at 500}510 cm was recorded in mollusc shells
Fig. 9. Summary diagram of multiproxy data for Twiss Marl Pond. Shaded bands show the Younger Dryas (YD), Preboreal Oscillation (PB) and
Gerzensee/Killarney Oscillations (G/K).
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
1735
Fig. 10. Correlation of oxygen isotopes (dO) from Twiss Marl Pond (mollusc shells), Crawford Lake (core SC; Yu and Eicher, 1998), Gerzensee (lake
carbonates; Eicher, 1980) and GRIP ice-core from Greenland (Dansgaard et al., 1993). All pro"les are based on depth scales (not shown except for
TMP). PB, Preboreal Oscillation (e.g., Lotter et al., 1992); IACP, intra-Aller+d cold period (Lehman and Keigwin, 1992); OD, Older Dryas (Dansgaard
et al., 1993); Gerzensee, Gerzensee Oscillation (Eicher, 1980; Lotter et al., 1992); Aegelsee, Aegelsee Oscillation (Lotter et al., 1992); G/K,
Gerzensee/Killarney Oscillations (Eicher, 1980; Lotter et al., 1992; Levesque et al., 1993). V-PDB, Vienna Pee Dee belemnite; V-SMOW, Vienna
standard mean ocean water.
and marl. This excursion may correlate with the intraAller+d cold period (IACP) in Greenland (Lehman and
Keigwin, 1992), Gerzensee Oscillation (G) in central
Europe (Eicher, 1980) and Killarney Oscillation (K) in
Atlantic Canada (Levesque et al., 1993), which was also
indicated by peaks of Al and K (Figs. 4 and 9). Both PB
and G/K events are weakly represented at TMP, probably because of lower temporal resolution of samples, but
the interpretation is supported by data from Crawford
Lake (Fig. 10).
The C/C ratio of mollusc shells and marl depends
mainly on local factors. The dC values in carbonates
are closely controlled by dC in dissolved inorganic
carbon (DIC) of lake water (Fritz and Poplawski, 1974).
The dC in DIC is in#uenced by (1) exchange with the
atmospheric CO reservoir, (2) biological productivity,
(3) decomposition of organic matter, (4) dissolved carbonate in in#owing groundwater, and (5) residence time
of lake water (evaporation e!ect) (Fritz and Poplawski,
1974; McKenzie, 1985; Siegenthaler and Eicher, 1986).
The dC results from mollusc shells, encrustations
and marl at TMP show parallel variations (Fig. 5B). The
mollusc shells are 2.6 to 7 depleted in C relative to
encrustations and marl, which is mostly within the
5}11 di!erence between calcite and aragonite and
comparable with other records (e.g., Stuiver, 1970; Fritz
et al., 1987). The dC values show a step-wise decrease
from &12,400 to 10,000 C BP before an abrupt increase of 2 at 10,000 BP. The Lateglacial trend of
C depletion may be caused by: (1) decreasing Cenriched detrital input from bedrock and groundwater,
(2) increasing supply of organic matter from progressively denser vegetation cover in the watershed and, presumably, increasing decomposition of organic matter in
the pond, and/or (3) decreasing aquatic productivity during the cold climate of the YD interval (Yu et al., 1997).
Vegetation development in the watershed has been emphasized as the major control of dC values in lake
water (Hammarlund, 1993; Hammarlund and Buchardt,
1996). A late-glacial dC decrease similar to TMP is
recorded from three cores at Crawford Lake (Yu, 1997a;
Yu et al., 1997) and at Nichols Brook and Gage Street
sites (Fritz et al., 1987; see Fig. 1 for location and see
below for detail). The positive shift in dC at the beginning of the Holocene was likely caused by increasing
aquatic productivity in response to climate warming, as
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Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
suggested by the proliferation of Najas and Potamogeton
(Fig. 11).
5.2. Regional vegetation, landscape and lake development
During the "rst several hundred years after local deglaciation, fossil pollen and plant macrofossils at TMP
indicate a pioneer near-treeless vegetation (Figs. 6 and 7).
This tundra vegetation was dominated by the shrubs
Alnus, Salix, Juniperus and Dryas and the herbs Cyperaceae and Artemisia, probably with scattered Populus
trees. Pollen concentration and, presumably, accumulation rate was extremely low. Pollen records indicate
a succession from an Alnus-Dryas-Cyperaceae sparse tundra or periglacial desert (TP-1a) to a Salix-JuniperusCyperaceae-Artemisia dense tundra (TP-1b). Increase in
vegetation cover was suggested by a slight increase of
pollen concentration from subzones TP-1a to TP-1b and
a gradual decline of major long-distance transported
Picea and Pinus pollen due mostly to increased pollen
contributions from denser local vegetation. The virtually
treeless vegetation was suggested by the absence of tree
macrofossils (except for a fragment of a possible Picea
seed), which were abundant in succeeding woodland/forest zones, and by abundant macrofossils of arctic/alpine
plants including Dryas integrifolia, Salix herbacea and
Silene cf. acaulis. These plants commonly grow in a treeless landscape.
The unstable landscape with severe slope erosion during the tundra zone (TP-1) produced clay and clayey
Fig. 11. Diagram showing aquatic plant fossils (colonies, pollen, seeds
and leaf spines) at Twiss Marl Pond. YD: Younger Dryas; BOA: B+lling
- Aller+d.
sediments with high silicates and a high concentration of
erosion-derived elements Ti, K, Mg, Al and V (Figs.
4 and 9). In dense tundra (subzone TP-1b), a slight
increase of organic matter during the transition from clay
to clayey marl indicates either increased terrestrial organic input from denser upland vegetation or increased
aquatic productivity. Macrofossils of the aquatic plants
Potamogeton and Najas yexilis and colonies of the green
alga Pediastrum in this subzone indicate the establishment of aquatic vegetation at that time. The presence of
Najas yexilis seeds and Typha latifolia pollen also suggests
that minimum mean July temperature was at least
12}143C based on the present distribution of these aquatics (Porsild and Cody, 1980; also see Kolstrup, 1979;
Isarin and Bohncke, 1999). No gastropods were found
during the sparse tundra/periglacial desert period (TP1a). The initial gastropod community that appeared during the dense tundra period (TP-1b) was dominated by
Gyraulus parvus and Fossaria decampi. Gyraulus parvus is
found in shallow ((1 m depth) and quiet water today;
Fossaria decampi is a shallow and cold water species
(Clarke, 1981). Both species suggest shallow water at the
site. All gastropod species recorded at TMP have broad
geographic ranges (Clarke, 1981) and usually are not
good indicators of climate, although their initial appearance might indicate availability of suitable habitats, indirectly related to climate.
Picea woodland became established at &12,000
C BP, as indicated by predominant Picea pollen and
abundant Picea needles and seeds. Herb pollen decreased
abruptly but persisted at low percentages, suggesting
a woodland rather than a closed forest. The abrupt
decrease of erosion-derived elements and increased organic matter and authigenic carbonate suggest slope
stabilisation after Picea woodland establishment in the
watershed. In the upper Picea subzone (TP-2a) at
10,920}10,000 BP, increased openings were suggested by
a slight increase of herb pollen, while peaks of silicate and
erosion elements suggest activated erosion episodes under cold climate during the YD period. The abrupt decline in abundance of gastropods at the tundrawoodland transition (TP-1b/TP-2a) at &12,000 C BP
(510 cm) suggests the dramatic response of lake biota to
changing environments brought by this major vegetation
shift. Valvata sincera replaced G. parvus as the dominant
gastropod at the beginning of the Picea period, and then
dominated the gastropod communities together with G.
parvus and F. decampi. Valvata sincera lives in lakes on
muddy substrate and submerged aquatics, likely Myriophyllum verticillatum as indicated by its macrofossils.
Pediastrum increased abruptly in the upper Picea
subzone (TP-2a) during the YD interval, which suggests
more silty water due to accelerated soil erosion under
cold climate (cf. Lotter and Holzer, 1994).
At 10,100 C BP, Pinus together with Betula, Quercus,
Fraxinus and Ulmus began to replace Picea in the forests.
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
However, Picea pollen shows a hesitant decline with
a recurrence at 9650 C BP, and its macrofossils persisted until &9400 BP. This transitional forest (TP-2b) was
also characterised by high Larix and Abies pollen and
macrofossils, reappearance of Alnus and Salix pollen, and
a decrease in pollen concentration. The gastropod
abundance shows another sharp decline at the onset of
the Holocene warming, which is also attributed to the
response of lake biota to changing environments caused
by this major climate shift. The Picea recurrence at
9650 BP corresponds with the PB cooling event. This
early Holocene climate oscillation has also been
documented, mostly from palynological evidence, along
the Great Lakes}St. Lawrence waterway (Richard
et al., 1992; Anderson and Lewis, 1992; Richard, 1994)
and in Newfoundland (Anderson and Macpherson,
1994). At TMP, there are no indications of change in
soil erosion from lithology and elemental geochemistry
during this period, probably indicating existence of a
stabilized forest-covered landscape in the watershed.
In the pond, there was no signi"cant change in gastropod species composition; gastropod abundance
gradually recovered from the 10,000 BP decline and soon
peaked. Pediastrum declined to and maintained at modest levels.
After a transitional period that lasted &500 years,
Pinus forest became established in the watershed with
Betula, Quercus, Ostrya/Carpinus and Ulmus. This closed
Pinus-dominated forest is indicated by negligible herb
1737
pollen and high tree pollen concentrations. The landscape was stable with little soil erosion and little input of
minerogenic matter into the pond, suggested by low
concentrations of erosion elements and disappearance of
Pediastrum. Proliferation of Najas yexilis indicates a rise
in temperature and increased aquatic productivity. Valvata tricarinata appeared and became co-dominant together with Helisoma anceps and Gyraulus parvus. Valvata
tricarinata lives on aquatic plants or submerged objects.
Abundant gastropods also indicate high biological productivity in the pond.
5.3. Climatic oscillations and biological responses
The BOA warm period from &12,700 to 10,920
C BP is indicated by high dO values, authigenic marl
accumulation, occurrence of warmth-demanding aquatics, and a!orestation. dO was among the "rst proxies
studied to show the warming (Figs. 9 and 12), which
indicates rapid response of dO in precipitation to climate. The high dO values during the BOA occurred at
the beginning of the gradual transition from clay to marl
and before the vegetational change from dense tundra to
spruce woodland (Fig. 9). High carbonate content in
authigenic marl results from increased lake productivity
owing to warm climate and from reduced minerogenic
input from watershed soil erosion linked to forest development. The occurrence of Picea peaks after peaks of
dO and carbonates suggests that upland vegetation
Fig. 12. Correlation of di!erent proxies at Twiss Marl Pond. The tentative temperature estimates are based on the present geographic distribution of
aquatic plants and modern dO - temperature relations.
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Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
lagged behind climate warming, probably due to slow
development of suitable soils (cf. Pennington, 1986), especially in the bedrock outcrop settings and very shallow
soils around this site. The Najas yexilis and Typha latifolia
fossils during the tundra period suggest a more rapid
response of aquatic plants to the BOA warming than
upland vegetation (cf. Iversen, 1954). An alternative explanation for di!erential response of aquatic plants and
upland vegetation to the BOA warm period would be the
individualistic behaviour of plant taxa because of their
di!erent climatic requirements. In northern latitudes,
aquatic plants grow in locations that freeze in the winter
and therefore are not subjected to the low temperatures
to which trees are subjected on the uplands. The presence
of aquatic plants may indicate suitable summer temperatures (Iversen, 1954) but may have little indication
about limiting winter temperatures for trees (T. Webb III,
1997, person. comm.). This was likely the case during the
Lateglacial, when there was strong seasonality with
relatively warmer summer and colder winter caused by
enhanced summer solar radiation (Berger, 1978;
COHMAP, 1988).
The YD climate reversal is indicated by negative excursions of dO, silicate peaks, high concentrations of
erosion elements and high Pediastrum. However, this
event occurred in the upper portion of the Picea pollen
subzone (TP-2a), indicating that no forest transformation
responded to the YD cooling. This stability re#ects insensitivity of nonecotonal vegetation in the study area at
that time (Yu, 1997a; Yu and Eicher, 1998). In the study
area the Picea dominance started at &12,000 C BP
and lasted for about 2000 years (see above and Table 2);
at the onset of the YD cooling at &11,000 BP the Picea
woodland belt might move latitudinally in response to
this climate #uctuation. However, TMP was in the
middle of this broad Picea belt at that time (see isopoll
maps in Jacobson et al., 1987), and thus no forest succession was recorded by pollen analysis. A similar case
was documented in Maine, USA, where Cwynar and
Levesque (1995) provide clear evidence from aquatic
chironomids for the YD climate reversal, but demonstrate
the insensitive response of nonecotonal upland vegetation.
Changes in upland vegetation during the YD have been
documented in the southern and northern boundaries of
the Picea belt at 11,000 BP, e.g., Picea recurrence in Ohio
and Indiana (Shane, 1987; Shane and Anderson, 1993) and
succession from a park-tundra with Picea and Fraxinus to
a dwarf-shrub tundra in northwestern Ontario (BjoK rck,
1985). Because of the individualistic behaviour of di!erent
plant taxa, the nonecotonal location for Picea at the
beginning of the YD does not mean that other upland
taxa were insensitive to the YD cooling. However, the
predominance of Picea may prevent the detection of
slight changes in other upland taxa by pollen analysis.
The Holocene warming at 10,000 BP was simultaneously recorded by all proxy data at the same
stratigraphic level, indicating no time-lag. The abrupt
positive shift of dO indicated rising air temperature.
Vegetation changed from open Picea woodland to
Pinus-dominated closed forest. The PB cooling at
9600 BP caused the recurrence of Picea because the site
was near the Picea-Pinus ecotone at that time. The slight
increase in Salix, Alnus and herbs during the gradual
transition from Picea to Pinus from &10,100 to 9500 BP
indicates changes in upland vegetation brought by this
forest reorganization.
There are no persistent and characteristic patterns in
gastropods species composition during either warm or
cold climate intervals. However, these species changes
together with dramatic declines in overall abundance
during the major vegetation and climate shifts re#ect
rapid and sensitive response and adjustment of aquatic
fauna to changes in lake and watershed conditions. Similarly, the Pediastrum peak during the YD and its persistence during the Picea-Pinus transition indicate sensitive
response of aquatic #ora to lake and watershed conditions, and indirectly to climate changes. Multiple proxy
investigations on more aquatic biota, such as
chironomids, ostracodes and diatoms, may provide further insights and understanding of responses of aquatic
system to climate changes.
5.4. Expression of Younger Dryas in eastern North
America
In eastern North America, there are regional variations of lithological and palynological evidence for the
YD event. In Atlantic Canada, the YD is indicated by
succession from woodland/forest to tundra or from shrub
tundra to herb tundra. Relatively low organic matter (as
measured by LOI) indicates renewed soli#uction, increased minerogenic inwash or reduced lake productivity
due to cooling (e.g., Mott et al., 1986; Mayle et al., 1993a).
Recently chironomids have been successfully used to
reconstruct summer surface-water temperatures for the
YD and G/K events in Atlantic Canada and northern
New England (Walker et al., 1991; Levesque et al., 1993,
1994; Cwynar and Levesque, 1995). In southern Quebec,
slight changes in pollen assemblages within herb tundra
and decreased pollen concentration suggest the YD climate reversal (Marcoux and Richard, 1995). In southern
New England, the YD is indicated by a decrease in the
pollen percentage of thermophilous temperate deciduous
trees such as Quercus and Fraxinus and increase of boreal
species such as Abies, Larix, Picea, Betula and Alnus (e.g.,
Peteet et al., 1990, 1993; Maenza-Gmelch, 1997). In the
southern Great Lakes region (western Ohio, Indiana and
Illinois), the indication is a recurrence of Picea after an
interval with relatively abundant Fraxinus and Ostrya/Carpinus pollen (Shane, 1987; Shane and Anderson,
1993). Recently, Mayle et al. (1993b) have summarized
the pollen evidence for the YD event in eastern North
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
America and propose the peak or persistence of the alder
(Alnus) pollen as an indicator. However, this is clearly not
the case at Twiss Marl Pond and Crawford Lake (Yu and
Eicher, 1998) in the eastern Great Lakes region.
In the eastern Great Lakes region, several studies have
explicitly examined climate and vegetation changes
around the Last Glacial}Interglacial transition. Two coring sites from Lake Erie documented major shifts in
dO (Fritz et al., 1975; Lewis and Anderson, 1992),
which were initially interpreted as representing temperature changes (Fritz et al., 1975) and later as caused by
changes in meltwater dynamics of the Great Lakes (Fritz,
1983; Fritz et al., 1987; Lewis and Anderson, 1992; also
see Rea et al., 1994a, b). Multiproxy data from several
small lake sites, however, failed to show evidence for
Lateglacial climate oscillations and even climate warming at the onset of the Holocene (Fritz et al., 1987; but see
below). Fossil insects at several sites from this region
indicate a stable climate condition without a major climate amelioration at the onset of Holocene (from 11,000
to 9000 C BP across the Picea}Pinus transition) (Miller
and Morgan, 1982; Schwert et al., 1985). In a recent
synthesis and mutual climatic range (MCR) analysis of 40
fossil beetle assemblages from 24 sites in northeastern
North America, Elias et al. (1996) concluded that `our
results show fewer changes during the Lateglacial, with
almost a plateau of summer temperatures from 13 to
10 kaa (p. 420).
At the Gage Street site in the city of Kitchener (see
Fig. 1B for location), a multidisciplinary investigation did
not identify Lateglacial climatic oscillations (Anderson,
1982; Schwert et al., 1985; Fritz et al., 1987). Two dates on
organic debris from the same stratigraphic level gave
signi"cantly di!erent ages (8600$100, BIRM-896;
10,700$900, WAT-157), all of which are much younger
than the age of '12,000 BP inferred from pollen zonation (Schwert et al., 1985). This dating inconsistency may
prevent correct identi"cation of critical intervals, although the dO pro"le from the Gage Street site does
show #uctuations in the Lateglacial and early Holocene.
Their dO records showed a negative excursion of
&1.2 in the upper Picea zone, especially from marl
samples (Fig. 13A) but also a slight excursion from Valvata tricarinata shell samples (Fritz et al., 1987). Their
chronology puts this interval older than 11,000 C BP.
However, the Picea-Pinus transition at 360 cm (the
revised zonation is slightly di!erent from theirs) has
recently been dated at about 10,000 C BP for
southwestern Ontario (see above and Table 2), which is
younger than the previously-accepted age of 10,600 BP
(Karrow et al., 1975). The herb}Picea transition at
420 cm was estimated around 12,000 BP, based on dates
from nearby sites (Table 2). These revised ages make this
negative excursion of dO at 400}360 cm from the Gage
Street site close to the Younger Dryas interval. The
depletion trend of dC at 400}360 cm (Fig. 13A) is also
1739
similar to the dC records at TMP and Crawford Lake
(Fig. 5B; Yu, 1997a). In addition, a slight negative excursion in dO at 350 cm corresponds to the recurrence or
persistence of Picea pollen at the Gage Street site (Fritz et
al., 1987), which may correlate with the Preboreal Oscillation at 9650 C BP as demonstrated at TMP as well as
Crawford Lake (Fig. 14). The negative excursion in dO
at 140}110 cm (at &7500 C BP) may correlate to the
8200 cal BP cooling event (HE5) as documented at Crawford Lake as well as in Greenland ice cores (Fig. 14)
among other sites.
At the Nichols Brook site in northwestern New York
(see Fig. 1B for site location), based on lithology, pollen,
fossil insects and detailed stable isotope records from
mollusc shells, Chara-encrustations and bulk carbonates,
Fritz et al. (1987) concluded that there was no evidence
for Lateglacial climatic oscillations. In fact, oxygen isotopes do show several shifts. The dO values from Valvata tricarinata shells show a negative shift of &1 from
!9.3 at 310}270 cm to !10.3 at 240}170 cm, and
a positive shift of &1 at 170 cm (Fig. 13B; Fritz et al.,
1987). Their chronology places the level at 250 cm to
12,100 C BP and the 170 cm level to 11,330 C BP.
The overall negative values (less than !11) of dO
from bulk marl indicate short residence time of water and
less evaporative enrichment, suggesting slowly #owing
water. Fossil insects also indicate #owing water from 250
to &125 cm, which suggest that wood from this interval
used for C dating might have been redeposited from
older sediments. If we consider three dates (10,850$240,
WAT-908; 11,330$270, WAT-913; and 11,720$390,
WAT-915) from the #owing water interval as too old,
then the reinterpreted chronology based on the other
four dates (9050$150, WAT-914; 9540$190, WAT891; 9690$370, WAT-912; and 12,320$240, WAT844) from wood in open or stagnant water may provide
alternative ages for these isotopic oscillations. This reinterpretation would put the negative dO shift at 250 cm
to about 11,600 BP and the positive shift at 170 cm to
10,200 BP. Then the negative excursion of &1 in dO
at 250}170 cm could correspond with the YD event. The
PB event is also suggested at this site by Picea pollen
recurrence (Fritz et al., 1987) and a slight negative excursion of dO from mollusc shells at 150 cm (9600 BP?).
In Lake Erie at coring site 13194, two large increases
(&5) at 12,700 C BP and 10,000 C BP in dO
from mollusc and ostracode shells are originally interpreted as re#ecting changes of dO in precipitation and
air temperature (Fritz et al., 1975). However, Fritz (1983)
and Fritz et al. (1987) realize that a reduction in glacial
meltwater in the Great Lakes drainage system may be
more important and likely accounted for these positive
isotopic shifts. During deglaciation, the Great Lakes contained meltwater (see Rea et al., 1994a, b), so that
a change in water budget may complicate climatic signals
in these isotopic records. Such a meltwater complication
1740
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
Fig. 13. The dO and dC pro"les, and revised pollen zones (dashed lines) and revised C chronology from (A) Gage Street site in southwestern
Ontario (Fritz et al., 1987) and (B) Nichols Brook site in northwestern New York (Fritz et al., 1987). See Fritz et al. (1987) for detailed description of
their chronology and pollen zonation. YD, Younger Dryas; PB, Preboreal Oscillation; HE5, Holocene Event 5, i.e., the 8200 cal BP cooling event as
documented in Greenland ice-cores (Alley et al., 1997) and Crawford Lake (Yu and Eicher, 1998).
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
Fig. 14. Event correlation of dO pro"les from carbonates at Crawford Lake (core SC: heavy curve; core BC: dashed curve) in the Great
Lakes region and from Greenland ice-core GISP2 (left bar-curve) on
C and calendar year time scales, respectively (Grootes et al., 1993; Yu
and Eicher, 1998). At Crawford Lake the sequence of climatic oscillations includes the BOA warming, G/K, YD, PB and HE-5 [Holocene
Event 5, i.e., the 8200 cal. BP cooling event], and possibly OD (Older
Dryas). All these oscillations recorded at Crawford Lake are with about
a third to half the amplitude of the Greenland record, and PB and HE-5
are half the amplitude of the YD and G/K/IACP at both locations.
Figure was modi"ed from Yu and Eicher (1998).
may also apply to the isotope record at the Long Point
site in the Lake Erie basin (Lewis and Anderson, 1992),
which also shows a large increase (&8) in dO at
&10,000 BP, and for sites from Lake Huron and Georgian Bay (Rea et al., 1994a, b). In fact, these authors
interpret their isotopic records as caused by meltwater
dynamics and lake level history, rather than climate
change.
At Twiss Marl Pond and Crawford Lake (Figs. 9, 10
and 14), the YD cooling is indicated by a negative excursion in dO, together with a persistence and slight increase of shrub and herb pollen, and decreased carbonate
and increased erosion-derived elements, suggesting more
openings in forests and accelerated soil erosion under
cold climate. A decrease in carbonate and increase in
minerogenic matter during the YD cold interval has been
recorded at some marl lakes in Switzerland (e.g., Lotter
and Zbinden, 1989; Lotter et al., 1992). Peak Pediastrum,
a planktonic green alga of silty water, also suggests
increased erosion episodes at TMP; Pediastrum peaks
during the YD interval have been documented in some
North American (BjoK rck, 1985) and European lakes (Lotter and Holzer, 1994).
There is a lack of independent dating before 11,000
C BP at TMP and Crawford Lake, although the onset
of the YD event was dated at 10,920$80 (TO-5505) at
TMP (Table 1) and its end was constrained by two dates
around 9650 BP just above it at Crawford Lake (Table 2;
1741
Yu and Eicher, 1998). The reinterpretation of isotopic
records at Gage Street and Nichols Brook sites relies
mostly on indirect dating. This lack of direct dates prevents reliable discussion on rates of climatic changes and
on synchroneity of climatic events, especially of centuryscale events. On the other hand, there is still a signi"cant
discrepancy on timing of the YD event from the bestdated records in the world, caused by dating uncertainty
and poor understanding of regional variations of the YD
event. The ages for the beginning of the YD from two
recent Summit ice cores (GRIP and GISP2), 30 km apart,
in Greenland are about 200 calendar years di!erent; both
chronologies are based on counting of annual ice layers
(Johnsen et al., 1992; Taylor et al., 1993). Also, this ice-core
age of the YD is more than 500 calendar years di!erent
from the best-dated European records based on varved lake
sediments (e.g., at Soppensee and Gosciaz; Hajdas et al.,
1993; Goslar et al., 1995). This dating uncertainty demonstrates how di$cult it is to prove or disprove the synchroneity of this climate event in these intensively studied regions,
not to mention other parts of the world. Thus, the best
approach is to make tentative correlation with these
well-documented records from the North Atlantic region
without necessarily assuming synchroneity (see Figs. 10
and 14), until we have more reliable independent dating
for each site with improved or new dating techniques.
The regional variations in evidence for the YD event as
summarized above suggest that the YD likely has di!erent expressions in paleoecological records depending on
geographic location and character of a particular site. In
some regions, the YD event may not be expressed as
a cold interval (see An et al., 1993; Roberts et al., 1993;
Kneller and Peteet, 1999). The YD climatic reversal may
also appear to be time-transgressive if su$cient sites are
available with reliable dating control (Cwynar and
Levesque, 1995). The multiproxy records from the eastern
Great Lakes region may correlate with records from the
Atlantic Seaboard but may imply di!erent nature of climate changes, such as changes in seasonality and balance
of thermal and moisture conditions, rather than simply
temperature and precipitation changes. In this case, the
usage of the term Younger Dryas may only have chronological sense, which means a climatic (not necessarily
cooling) event occurred from about 11,000 to 10,000
C BP, although the YD-age climate events in di!erent
regions may be connected in mechanisms and causes.
5.5. Causes of Lateglacial climatic oscillations in the Great
Lakes region
There are two hypotheses to explain the YD age climate reversal in the Great Lakes region. One hypothesis
is that the YD-age cooling was a locally induced climate
change, caused by enhanced cold glacial meltwater
(Lewis and Anderson, 1992) or simply by highstands
of proglacial lakes (Rea et al., 1994b). Another hypothesis suggests that it was a regional expression of the
1742
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
broad-scale climatic oscillation, connected with changes
in oceanic and atmospheric circulations (Broecker et al.,
1985; Rind et al., 1986; Shane, 1987; Wright, 1989;
Peteet et al., 1997; Yu and Eicher, 1998), which
links it with climatic oscillations from other regions in
the world.
Lewis and Anderson (1992) attribute a gradual decrease of 3 in dO from&11,000 to 10,500 C BP in
the Lake Erie basin to the isotopically light meltwater
in#ow from the highstand of proglacial Lake Algonquin
in the Lake Huron basin. Using as analogue the cooling
e!ect of lowlands west of modern Hudson Bay (Rouse,
1991), they propose that the cooling from cold
meltwater-rich proglacial lakes is responsible for the
Picea recurrence in the southern Great Lakes region
(Shane, 1987; Shane and Anderson, 1993) and for pollen
anomalies at other sites as summarized in Lewis and
Anderson (1992). Their hypothesis predicts that the
strongest evidence should be found on the lowlands and
at sites adjacent to the proglacial lakes, and they stated
`The cooling is absent or of limited amplitude east of
Lake Erie and on higher inland sites in southern Ontarioa (Lewis and Anderson, 1992; p. 248). However, the
evidence for the YD cooling at Twiss Marl Pond and
Crawford Lake contradicts this prediction. These two
sites are atop the Niagara Escarpment on highlands, and
both are distant from the proglacial lakes at that time. If
we accept the revised interpretation of isotopic records
from the Gage Street and Nichols Brook sites (see above),
then these two sites also do not support their prediction
because they are both on highlands (see Fig. 1 in Lewis
and Anderson, 1992) and are distant from proglacial
lakes. Shane and Anderson (1993) also pointed out
that the strongest YD signal is found at sites distant
from the lakes, which is opposite to what the Lewis
and Anderson hypothesis predicts. On the basis of the
stable isotopes and pollen record from a small lake near
Glacial Lake Agassiz during its second expansion, Hu
et al. (1997) showed that lake-induced cooling in the
early Holocene did not cause signi"cant change in
surrounding vegetation. In the Great Lakes region,
the available data suggest that the detectability of
pollen anomalies during the YD interval relies more on
site location relative to ecotonal vegetation, rather
than on the distance from proglacial lakes or downwind/upwind directions (Saarnisto, 1974; Shane, 1987;
Wright, 1989).
On the basis of isotope and lake-level data from the
sediments of Lake Huron and Georgian Bay, Rea et al.
(1994a, b) found that the highstands, including Lake
Algonquin at 11,200}10,200 C BP, were characterised
by isotopically heavy water, indicating a much smaller
meltwater component relative to input from local precipitation and run-o!. In contrast, the intervening lowstands were characterised by cold, dilute, isotopically
very light water, as indicated by their isotope and ostra-
code data (Rea et al., 1994b). These new results do not
appear to support the Lewis and Anderson's (1989, 1992)
hypothesis, although Rea et al. (1994b) maintain that
lake-e!ect cooling is probably related more to surface
area than the meltwater volume. Nevertheless, the relatively warm lake water during highstands of the Great
Lakes certainly would have a less pronounced cooling
e!ect. In fact, the lake areas were smaller than (for Lakes
Ontario, Erie and Michigan) or similar to (for Lake
Huron) the present lake areas of the Great Lakes at
10,800 BP during the main Algonquin phase (see Fig. 2b
in Lewis and Anderson, 1992).
The cooling e!ect of large lakes on adjacent land does
occur. However, the main di$culty with the Lewis and
Anderson hypothesis is how this cooling e!ect links to
the YD climate reversal, which requires more enhanced
lake and ice-sheet cooling during the YD interval than
before and after. The extensive ice sheet itself might very
well have overprinted the localized cooling caused by
proglacial lake water by modifying atmospheric temperature (Clark et al., 1999), which is di!erent from the
present Hudson Bay situation. At present, the extensive
and continuous Hudson Bay generates a distinct air mass
o! its coast, whereas the proglacial lakes would be less
important than the extensive ice sheet to the north in
generating an air mass at that time.
Anderson and Lewis (1992) applied a similar mechanism of a meltwater-induced cooling e!ect to account for
early Holocene pollen anomalies at 10,000}8000 C BP
as summarised in their paper. However, their hypothesis
does not predict evidence for the PB cooling from sites
such as TMP and Crawford Lake, and perhaps the Gage
and Nichols sites (see above), because these sites are
distant from any proglacial lakes at 9500 BP (see Fig. 1 in
Anderson and Lewis, 1992). In fact, the areas of the Great
Lakes at 9500 BP, except Lake Superior, were much
smaller than today, and water levels in the Huron and
Michigan basins were near their lowest levels ever (see
Fig. 20, in Teller, 1987).
The broad-scale cooling hypothesis for explaining the
abrupt climate changes, including the YD event, relies on
some kind of threshold or trigger in the regional or global
ocean and atmosphere climate systems. A prevalent hypothesis is that the routing change of the Laurentide ice
sheet meltwater from the Mississippi drainage system to
the St. Lawrence river at &11,000 BP triggered the YD
cooling event (Broecker et al., 1989; Broecker, 1994). This
hypothesis maintains that a sudden dilution of North
Atlantic surface water in response to this meltwater diversion caused a reduction in the production rate of
North Atlantic Deep Water (NADW) and a shutdown of
thermohaline circulation. This shutdown ceased the
northward movement of subtropical warm surface
waters and reduced delivery of ocean heat to subpolar
region, causing cooling of the high latitudes around the
North Atlantic, especially downwind, i.e., Europe. This
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
cooling e!ect extended upwind in North America as far
as Ohio (Wright, 1989). Using a coupled ocean-atmospheric general circulation model, Mikolajewicz et al.
(1997) found that a cold North Atlantic could cool
the entire North Hemisphere downstream through the
atmosphere.
The widespread evidence for a YD age event in localities far from the North Atlantic Ocean suggests that the
trigger for the YD cooling may lie in the atmosphere
rather than in the North Atlantic (Denton and Hendy,
1994; Lowell et al., 1995; Thompson et al., 1995). Water
vapour is far more abundant and more potent at trapping heat than carbon dioxide, and reductions in atmospheric water vapour could lead to a reverse greenhouse
e!ect. However, recent data indicate that abrupt warming at the end of the last glaciation (B+lling warming) and
that at the end of the Younger Dryas (Holocene warming) in the North Atlantic occurred several decades before the tropical warming indicated by an increase in
atmospheric methane concentration, suggesting that the
trigger of deglacial climatic oscillations was related to
North Atlantic Ocean rather than to changes in the
tropics (Severinghaus et al., 1998; Severinghaus and
Brook, 1999). Although at a global scale the exact causes
of rapid climate oscillations, including the YD event, are
still incompletely understood, accumulating evidence
1743
from di!erent regions and climate modelling will ultimately shed light on their causes and mechanisms and
the Earth's climate dynamics as a whole.
In summary, the hypothesis of local meltwater-induced
cooling cannot be applied to the YD and PB climate
oscillations recorded in the Great Lakes region. At TMP
and Crawford Lake (Figs. 10 and 14), the sequence
and relative magnitude of climatic changes match in
detail the records from the North Atlantic region and
indicate that these oscillations are likely an expression
of broad scale, probably global, climate changes.
This similarity cannot be satisfactorily explained by any
local or regional processes, and it must relate to a
more fundamental climate trigger in the coupled ocean}
atmosphere}ice sheet climate system, which would
explain these widespread, abrupt climate events. In addition to the YD event, other minor oscillations (PB,
G/K) may also have occurred beyond the North Atlantic
region. The close match of the dO records at sites in
the Great Lakes region and at other sites in the North
Atlantic region (Fig. 15) indicates that the climatic
forcing was coeval and acted in the same manner in
both the North Atlantic region and interior North
America, and that the climatic signals were likely
carried through the atmosphere over the Northern
Hemisphere.
Fig. 15. The geographic distribution of Lateglacial and early Holocene climate oscillations around the North Atlantic Ocean (from central Europe to
central North America) (Lehman and Keigwin, 1992; Dansgaard et al., 1993; Yu and Eicher, 1998; and many regional syntheses in Lowe, 1994). The
positions of the North Atlantic polar front during the last deglaciation were from Ruddiman and McIntyre (1981), which indicate that the polar front
shifted southward during the cold YD interval. The big open circle in the eastern Great Lakes region shows location of Twiss Marl Pond and Crawford
Lake, where multiple proxy evidence has been found for three out of four recognized Lateglacial and early Holocene climate oscillations.
1744
Z. Yu / Quaternary Science Reviews 19 (2000) 1723}1747
6. Summary and conclusions
(1) Multiple proxy data from Twiss Marl Pond provide a comprehensive reconstruction of lake, vegetation
and climate changes around the Pleistocene}Holocene
boundary in the eastern Great Lakes region. The combined pollen and plant}macrofossil records indicate treeless tundra for several hundred years after deglaciation.
This pioneer vegetation consisted of arctic/alpine plants
Dryas integrifolia, Salix herbacea, Silene cf. acaulis, Alnus
cf. crispa and Juniperus cf. communis.
(2) At Twiss Marl Pond, the Younger Dryas (YD)
cooling event at 10,920}10,000 C BP is indicated by
a negative excursion (&1.5) in dO, an increase in
minerogenic matter and higher concentrations of erosion-derived elements (Al, Na, K, Ti and V). The Preboreal Oscillation (PB) at 9650 C BP is indicated by
Picea recurrence and a negative excursion (&0.4) of
dO. A slight increase of Al and a negative excursion of
dO suggest a pre-YD Gerzensee/Killarney Oscillation.
(3) The dO was among the "rst proxies studied
to show the B+lling-Aller+d (BOA) warming, indicating rapid response of dO in precipitation. The establishment of Picea woodland after peaks of dO and
carbonate suggests a lagged response by upland vegetation. Occurrence of warmth-loving aquatic fossils indicates sensitive response of aquatic plants to BOA
warming.
(4) Increased openings in forest and accelerated soil
erosion were in response to YD cooling, but there was no
forest transformation, due to insensitive Picea-predominated nonecotonal vegetation. Picea-Pinus ecotonal vegetation in the early Holocene shows sensitive response to
PB cooling. Proliferation of Pediastrum during YD interval was caused either directly by temperature decline or
by more silty water due to accelerated soil erosion.
Abrupt declines in abundance of freshwater gastropods
during the tundra-woodland transition and Holocene
warming suggest that aquatic biota were sensitive to
changing environments. These new results con"rm the
notion that terrestrial and aquatic systems respond differently to climate changes (Wright, 1984) and suggest the
importance of multiproxy records for reliable paleoclimate reconstruction.
(5) The regional variations in evidence for the YD
event in eastern North America suggest that climate
oscillations likely have di!erent expressions in paleorecords depending on geographic location and character
of a particular site. Revised chronology of previously
published sites (Gage Street, and Nichols Brook) in the
eastern Great Lakes region places their major dO shifts
similar to the YD and PB events as shown from TMP
and Crawford Lake. The sequence and relative magnitude of climatic oscillations from these sites match in
detail with records from the Atlantic Seaboard, suggesting that these oscillations are an expression of broad-
scale, probably global, climate change rather than local
meltwater-induced climate cooling. The fact that these
sites are further within the interior of North America
implies that climate signals were likely carried over the
Northern Hemisphere through the atmosphere.
Acknowledgements
I thank U. Eicher for his excellent stable isotope
measurements at the University of Bern, Switzerland and
for providing numerical data from Gerzensee; J. Denhart
and J. H. McAndrews for "eld assistance with sediment
coring; D. Siddiqi, J.H. McAndrews, R. Farvacque, R.
Hancock, S. Aufreiter, and J. Yang for laboratory assistance; J.H. McAndrews, T. Webb III, and H.E. Wright, Jr.
for helpful comments and/or discussions; and F.S. Hu
and D.M. Peteet for constructive reviews. This work was
supported by the Natural Sciences and Engineering Research Council of Canada, Royal Ontario Museum and
University of Toronto (Open Fellowships and Department of Botany).
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