ARTICLE IN PRESS
Quaternary Science Reviews 23 (2004) 1901–1924
Pollen-based summer temperature reconstructions for the eastern
Canadian boreal forest, subarctic, and Arctic$
Michael W. Kerwina,*, Jonathan T. Overpeckb, Robert S. Webbc, Katherine H. Andersond
b
a
Department of Geography, University of Denver, 2050 E Iliff Ave, Denver, CO 80208-0183, USA
Department of Geosciences, Institute for the Study of Planet Earth, University of Arizona, Tucson, AZ 85721, USA
c
NOAA Climate Diagnostics Center, Boulder, CO 80305, USA
d
Institute of Arctic and Alpine Research, University of Colorado, Boulder, CO 80309, USA
Received 16 September 2003; accepted 11 March 2004
Abstract
Pollen-based paleoclimate reconstructions using response surface and modern analog methods reveal an 8000-year record of July
temperature fluctuations for 25 eastern Canadian lake sites located from the forest–tundra to the high Canadian Arctic. Postglacial
conditions, characterized by warmer than present summer temperatures, prevailed in Baffin Island and NE Labrador beginning
about 7500 and 7000 14C yr BP, respectively, resulting in warmer than present conditions throughout the region by 6000 14C yr BP
(+0.5 C to 1 C). Further south, in Quebec and W Labrador, July temperatures were 1–2 C colder than present until after 6000
14
C yr BP, and only reached modern values after all residual Laurentide ice had melted. Increased summertime insolation and the
final disappearance of Laurentide Ice during the middle Holocene probably caused July temperatures throughout eastern Canada to
peak between approximately 5000 and 3500 14C yr BP. Mid-Holocene warming relative to today was more pronounced in Baffin
Island and NE Labrador (+1 C to 2 C) compared to the boreal and subarctic regions of Quebec and W Labrador (o+1 C). Over
the past 4000 years, decreasing summertime insolation and colder sea surface temperatures in the Davis Strait and Labrador Sea
contributed to a decline in July temperatures of 1–2 C throughout Baffin Island, and the tundra regions of N Quebec and Labrador.
The absence of similar cooling in the records from the boreal forest may support the notion that the mean position of the
summertime polar front blocked the colder Arctic air during the late Holocene.
r 2004 Elsevier Ltd. All rights reserved.
1. Introduction
High latitude regions of eastern North America have
experienced large environmental and climate fluctuations since deglaciation (e.g., Bradley, 1990; Fisher et al.,
1995; Overpeck et al., 1997; Barber et al., 1999; CAPE,
2001), and are expected to be among the first regions to
respond significantly to global warming over the next
century (Houghton et al., 1996). Even subtle Arctic
climate variations can be amplified greatly through the
melting of snow and ice (including sea-ice) which
decreases surface albedo, the prevalence of wintertime
$
Supplementary data associated with this article can be found in the
online version at doi: 10.1016/j.quascirev.2004.03.013
*Corresponding author. Tel.:+1-303-871-3998; fax:+1-303-8712201.
E-mail address: [email protected] (M.W. Kerwin).
0277-3791/$ - see front matter r 2004 Elsevier Ltd. All rights reserved.
doi:10.1016/j.quascirev.2004.03.013
temperature inversions that trap relatively warm air at
the surface, and through cloud feedbacks (Walsh and
Crane, 1992; Curry et al., 1996; Houghton et al., 1996;
Moritz et al., 2002). In turn, local Arctic climate
perturbations can exert hemispheric and global influence
via changes in atmospheric circulation, CO2 and CH4
concentrations, and surface water runoff that affects
thermohaline circulation (Rind, 1987; Broecker and
Denton, 1989; Weller et al., 1995; Weller, 1998).
Documenting how the climate system has behaved in
the past can provide important insights into how the
climate system may behave over the next few centuries
(Overpeck, 1995). In particular, quantitative reconstructions of past high latitude climate change (e.g., CAPE,
2001) can provide constraints on the magnitude of
future Arctic change.
Lacustrine pollen records are among the most widely
available and reliable proxy indicators of past high
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latitude environmental, vegetation, and climate changes
(e.g., Ritchie, 1987; Webb et al., 1987, 1993b, 1998;
Anderson et al., 1989, 2002). Numerous pollen records
have been generated in Arctic and subarctic regions of
eastern Canada over the past four decades. The
palynological interpretation of these records document
the large-scale changes in vegetation and climate that
took place during the late Quaternary (e.g., Bartley and
Matthews, 1969; Nichols, 1970, 1974, 1975; Boulton
et al., 1976; Andrews et al., 1980a; Davis, 1980; Lamb,
1980; Mode, 1980, 1996; Short et al., 1985; Gajewski,
1995; Gajewski et al., 1995; Williams et al., 1996;
Gajewski and Frappier, 2001). Whereas the absolute age
and duration of these changes are becoming increasingly
better constrained through the use of annually laminated lake sediments (Lamoureux and Bradley, 1996;
Hughen et al., 2000; Moore et al., 2001) and AMS
radiocarbon dating, the quantitative interpretation of
these records, especially in the Arctic, has been
hampered by the presence of exotic pollen from distant
regions (see Section 3) and limited modern pollen data
needed to quantify the modern relationship between
pollen and climate.
From 1994–1998, we collected pollen data from
surface lake sediments in the calibration data sparse
NE Canada. We focused specifically on Baffin Island
and N Labrador, where less than 10 modern pollen
samples from lake sediments existed prior to our efforts.
Ultimately, we generated new modern pollen data from
76 lakes in Labrador and Baffin Island, and combined
these results with existing lake-based pollen data from
199 boreal and subarctic sites in eastern Canada. The
resulting high-quality modern pollen dataset has greatly
improved our understanding of pollen–vegetation and
pollen–climate relationships in the Arctic and subarctic
regions of eastern Canada (Kerwin, 2000; Kerwin et al.,
in preparation).
In this study, we use this new modern pollen database
to produce improved quantitative reconstructions of
past climate variations from Arctic and subarctic pollen
records. We apply the response surface and modern
analog techniques to 25 previously published Holocene
pollen records from the boreal forest to the high
Canadian Arctic to reconstruct past climate conditions.
We reconstructed July temperatures because the link
between summer temperatures and boreal-subarctic to
Arctic plant growth and productivity is well documented
(e.g., Birks, 1981; Bennett, 1988; Payette et al., 1989;
Huntley, 1991; MacDonald et al., 1993). In addition,
our reconstructions can be directly compared to other
summer temperature reconstructions. Our reconstructions beyond 8000 14C yr BP are significantly less robust
because only a few records extended beyond that date,
and most assemblages in the few records prior to 8000
14
C yr BP had no modern analogs (e.g., Overpeck et al.,
1992). Specific attention is focused on July temperature
anomalies at 6000 14C yr BP, and the onset of the
Neoglaciation during the late Holocene. The production
of coherent, millennia-long records of Arctic and
subarctic temperature change will not only bolster our
understanding of natural climate variability, but can
be used to evaluate how well the latest generation of
Earth System Models (e.g., Crucifx et al., 2002) simulate
the full range of past environmental variability and
change.
2. Regional setting
We focus on NE North America from Quebec,
Labrador, and Newfoundland in the south to all of
Baffin Island in the north (50–73 N latitude; 46–85 W
longitude). Mean July temperatures vary from a high of
17.5 C ( 22.5 C January;
1.5 C annual) in the
interior portions of S. Quebec to a low of 4.5 C
( 30 C January; 16.5 C annual) along the NE coast
of Baffin Island (Fig. 1). Annual precipitation is highest
along the east coast of Labrador and Newfoundland
(1200–1600 mm/year), and lowest in the polar desert
regions of NW Baffin Island (125–200 mm/year). Steep
temperature and precipitation gradients characterize
boreal and subarctic regions of Quebec and Labrador.
The climate gradients across Baffin Island, where most
of the new pollen samples were collected, are not as
steep, especially with respect to July temperature (Fig. 1).
Average July temperatures across all of Baffin Island
range between 5 C and 7 C. The east–west trending
July temperature gradient illustrates the impact of sea
surface temperatures in Baffin Bay and the Davis Strait
on the summer climate of Baffin Island. Cold (o2 C)
summer ocean temperatures along the east coast of
Baffin Island suppress the coastal air temperatures
relative to warming inland. The maritime effect is
diminished when the ocean surrounding Baffin Island
is covered by sea ice, resulting in a strong north–south
gradient in January temperature.
3. Data
3.1. Pollen data
Our pollen data analyses used 275 new and previously
published modern and 25 previously produced fossil
records from lakes in eastern Canada (Fig. 2 and
Supplementary Tables 1 and 2 Blake, 1972). Our
modern pollen database spans five vegetation zones
from the Picea mariana dominated boreal forest in the
south to the rock and herb barrens of the northern higharctic tundra (Polunin, 1951; Porsild, 1964; Young,
1971). Our study region was selected to encompass the
range of Holocene vegetation and climate regimes in
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1903
Fig. 1. Mean (1951–1980) July and January temperature and annual precipitation values on the 25-km grid created by Bartlein and others (Bartlein
et al., 1994; Thompson et al., 1999). Large lakes and ice sheets on Baffin Island lack climate data and show up as white spaces on these maps.
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Fig. 2. Map of eastern Canada showing the five vegetation zones in our study area (Porsild, 1964; Polunin, 1951) and 275 modern (top) and 25 fossil
(bottom) pollen sites used in our analyses. The 76 sites represented by yellow circles in the modern sites panel were collected by us from 1994 to 1997.
The sites portrayed by blue squares represent existing data from the North American Pollen Database or the Brown University modern pollen
database (Avizinis and Webb, 1983). See Supplementary Tables 1 and 2 for details about each site.
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eastern Canada. We incorporated only the most
abundant, climatically sensitive pollen taxa for the
response surface analyses and paleoclimate reconstructions (Supplementary Table 3). Any taxon that did not
exceed 2% in at least 2 of the modern pollen samples or
3% in at least 5 fossil samples was excluded from the
sum (Webb et al., 1993a). Pinus was removed from the
sum because it is persistent in sediments from central to
northern Baffin Island more than 1400 km from the
nearest Pinus banksiana tree. Changes in Pinus pollen,
therefore, are not directly related to the modern
vegetation or climate within the vast majority of our
study area. Whereas Pinus is likely to be an important
taxon for reconstructions in the SW portion of the study
area, it had to be removed from the sum in order to
attempt reconstructions in the mid- to high arctic.
Ambrosia was removed from the sum because of
potential human disturbance issues in the southern
portion of our Canadian study area and because
Ambrosia plants are not native to Arctic and subarctic
Canada (Porsild, 1964). Several herbaceous taxa including Caryophyllaceae, Fabaceae, Ranunculus, and Rosaceae were combined into an ‘‘Arctic herb’’ group
because their distribution in geographic and climatic
space was similar, but the abundance of individual
pollen types were not significant.
A challenge in this study was how to handle the
abundant, but potentially exotic pollen types such as
Alnus and Betula. Alnus crispa is one of the co-dominant
shrubs in the subarctic vegetation zone of Quebec and
Labrador (Rousseau, 1968). Variations in Alnus pollen
certainly reflect differences in the vegetation and climate
of Labrador and Quebec (Lamb, 1984), but not in Baffin
Island (where Alnus sp. are absent from the local flora)
despite the presence of large quantities of Alnus pollen in
Baffin Island lakes (Short et al., 1985; Kerwin, 2000;
Kerwin et al., in preparation). Likewise Betula glandulosa is considered the best indicator of the low-arctic
vegetation zone on Baffin Island and hence the climate
of that region (Andrews et al., 1980b; Jacobs et al.,
1985). However, the relationship between Betula pollen
and either climate or vegetation differences in the midand high arctic (where Betula sp. are absent from the
local flora) is elusive. Complicating the matter further is
the presence of tree species such as Betula papyrifera and
Alnus rugosa in the boreal forest on the SE edge of our
study area. The pollen from these tree species is typically
not differentiated from shrub species of the same genus.
Hence, non-monotonic relationships between pollen and
climate can arise simply from the mixing of tree and
shrub species and not differences in climate.
We used two different pollen sums to address these
issues: one with Alnus in the sum for the boreal and
subarctic regions of Quebec and Labrador where Alnus
sp. is part of the flora, and one without Alnus pollen for
Baffin Island where modern Alnus pollen is entirely
1905
exotic. Betula was included in both sums because it is the
dominant shrub in the low-arctic vegetation zone
(Porsild, 1964; Polunin, 1951), and crucial to the
analysis of climate in that region. The pollen percentages
used here are based on a sum of 15 (or 14 when Alnus is
excluded in Baffin Island) pollen taxa, selected from the
50 most abundant terrestrial pollen types within a
combined dataset of 275 modern pollen samples and
1255 fossil pollen samples from 33 lake sites (Supplementary Table 3).
3.2. Climate data
Temperature and precipitation estimates throughout
our Eastern Canadian study area were derived from a
25-km climate grid for North America generated by
Bartlein and others (Bartlein et al., 1994). Bartlein et al.
(1994) gridded January, July, and annual temperature
and precipitation using modern climate observations
(1951–1980) from over 45 meteorological stations in
eastern Canada. In most cases, singular value decomposition least-squares techniques provided regression
equations to estimate climate at each grid point as a
function of location and elevation from the station
(Thompson et al., 1999). In parts of the Arctic where
meteorological observations are sparse, shorter-term
climate averages were used (if available) to estimate the
modern climate or, in some cases, digitized climate data
from the World Meteorological Organization atlas for
North America were utilized.
We assigned climate values to each lake by evaluating
the four 25-km climate grid points nearest to each lake
site. If the lake bordered the ocean, the algorithm
evaluated less than four grid points, rather than
searching for data more than 25 km from the lake. The
‘‘modern climate’’ was assigned to each lake using
the data from the grid point with the closest elevation to
the lake without any additional interpolations. We
considered and dismissed two other algorithms: one
that calculated a lapse rate using climate and elevation
data from the two grid points with the lowest and
highest elevations among the four, and a second that
calculated a lapse rate using the two grid points with
elevations closest to the lake. There was no discernable
improvement in the modern pollen calibrations using a
calculated lapse rate to adjust the climate of the closest
neighboring grid point to the elevation of the lake.
4. Methods
4.1. General assumptions
We applied the response surface (Bartlein et al., 1986;
Webb et al., 1987, 1993b; Overpeck et al., 1991; Prentice
et al., 1991; Webb et al., 1993a) and modern analog
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(Overpeck et al., 1985, 1992; Anderson et al., 1989)
techniques for our July temperature reconstructions. We
did not use the multiple regression approach (e.g., Webb
and Bryson, 1972; Webb and Clark, 1977; Andrews
et al., 1980a; Howe and Webb, 1983; Bartlein et al.,
1984; Bartlein and Webb, 1985; Huntley and Prentice,
1988; Diaz et al., 1989) because no-analog assemblages
are difficult to identify (Weisberg, 1985; Bartlein and
Whitlock, 1993). In addition previous multiple regression studies in high latitude regions have demonstrated
problems where the modern pollen–climate relationship
is non-linear and subject to local environmental and
depositional peculiarities (Guiot, 1990; Huntley, 1990).
Regardless of the reconstruction approach utilized, a
common set of assumptions is required (Webb and
Clark, 1977). We assume that the vegetation of eastern
Canada is currently, and was during the time of our
reconstructions (Holocene), in equilibrium with climate
change. We also assume that variations in the fossil
pollen record are predominantly attributable to climatic
change rather than other non-climatic influences. In
addition, we must assume that the range of modern
pollen, vegetation, and climate in our study area is
sufficient to describe these relationships in the past.
With regard to the first assumption, vegetation is
never completely in equilibrium with the mean state
under changing climate conditions because the vegetation response must lag any change in forcing (Kutzbach,
1976). The fundamental question of equilibrium depends on the time scale of interest (Webb, 1986).
Palynological studies in general deal with fossil samples
that have accumulated over several decades or more
allowing time to smooth out the initial lag between
climate and vegetation (MacDonald and Edwards,
1991). Many paleoclimate reconstructions, including
ours, focus on century to millennial scale changes. On
these time scales, analyses of numerous pollen records
indicate that vegetation does respond to climatic events
of 500–1000 years in duration (e.g., Webb, 1986;
Bartlein and Whitlock, 1993; Grimm et al., 1993;
Whitlock and Bartlein, 1997; Anderson et al., 2002). If
tundra vegetation responded to climate changes during
the Holocene faster than forested vegetation (Overpeck
et al., 1991), one can expect that Holocene arctic
vegetation and climate have been in equilibrium on
century to millennial time scales. In terms of nonclimatic influences on plant ecology, Davis et al. (1986)
demonstrated that seed dispersal variations unrelated to
climate can influence plant communities on a regional
scale (hundreds of kms), but the effect is limited to
minor aspects of thinning and distribution. The existence of numerous fossil pollen assemblages that are
taxonomically similar to modern pollen assemblages
does suggest similar vegetation composition.
Previous analyses of fossil pollen records (Overpeck
et al., 1992) suggest that our modern pollen database
will not include all of the pollen assemblages and climate
regimes that existed in eastern Canada, especially in the
early Holocene. By determining quantitatively when
fossil assemblages are unlike any in the modern
database and identifying these assemblages as noanalog, we reduce the possibility of reconstructing
climates based on vegetation and climates that were
unlike the modern dataset.
Both the response surface and modern analog reconstruction techniques require a two-step process to generate
quantitative paleoclimate estimates from fossil pollen
data. First, a quantitative relationship between modern
pollen and climate must be derived. Then the fossil
assemblages are assigned climate values based on the most
similar modern analog. The two reconstruction techniques
we used differ mainly in how the modern pollen–climate
database is constructed. For the response surface
approach, this database consists of interpolated values
of the individual pollen types and their associated climate
values. The modern analog approach relies on the raw
pollen and climate data without any interpolations.
4.2. Quantitative reconstruction techniques
We utilized software written by Bartlein and described
in Prentice et al. (1991) and Webb (1990) to construct
modern pollen and climate response surfaces. These
response surfaces are useful for visualizing how the
abundance of key pollen taxa fluctuate in climate space,
and for reconstructing past climates based on the
relationships portrayed in the surfaces (Kerwin, 2000;
Markgraf et al., 2002). Our experimentation with the
various fitting and smoothing algorithms ultimately lead
us to construct two different sets of response surfaces,
one for N Quebec and Labrador where the climate
gradient is relatively steep and one for Baffin Island
where the climate gradient is shallow. For the boreal
and Subarctic regions of mainland Canada (where the
pollen sum included Alnus), we used a window width
that was 1/5 the size of each axis with one smoothing
pass. For the Baffin Island reconstructions (where the
pollen sum excluded Alnus) we used a window width
that was 1/10 the size of each axis with no smoothing
pass (Kerwin, 2000). These gridding algorithms resulted
in a set of response surfaces, one for each pollen taxon
that illustrates the spatial distribution of each pollen
taxon in climate space. Stacking the individual response
surfaces in climate space enabled synthetic modern
pollen spectra to be defined for each grid point in the
multi-dimensional climate space (Webb et al., 1993a).
These synthetic pollen spectra were then used as the
modern pollen–climate database from which to compare
with the fossil pollen data.
The method of modern analog (Overpeck et al., 1985,
1992; Guiot, 1987; Anderson et al., 1989; Bartlein and
Whitlock, 1993; Sawada et al., 1999) compares fossil
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pollen assemblages from an unknown climate, to a
database of modern pollen assemblages that originated
from a known climate. The method is based upon the
expectation that a particular climatic regime, both today
and in the past, will foster a unique set of floral and
pollen assemblages. If a fossil pollen assemblage can be
numerically linked to a modern assemblage, we assume
that the modern environmental and climatic conditions
that produced the modern assemblage are equivalent to
the past conditions that gave rise to the fossil pollen
assemblage. In practice, the modern analog approach is
dependent upon the use of a dissimilarity coefficient that
measures the numerical similarity of multivariate
samples, and is essential for identifying no-analog
assemblages (defined as samples that are unlikely to
have a modern counterpart), which can, when not
properly identified, lead to misleading paleoclimate
results.
1907
similar to at least one of the response surface spectra
below a SCD of 0.20 (Supplementary Table 5).
Uncertainty in the reconstructed climatic values for
each reconstruction approach was estimated by the
standard deviation of the weighted mean (using the
inverse of the SCD) of the climate values associated with
the ten best analogs. Bartlein and Whitlock (1993)
determined that this standard practice underestimates
the true uncertainty because SCD weighted standard
deviations can be expected to down weight large
deviations from the mean. Nonetheless it is still the
most rigorous way to estimate uncertainty in the
reconstructed climatic values (Bartlein and Whitlock,
1993).
5. Results
5.1. Pollen response surfaces
4.3. Determining the similarity between fossil and modern
pollen assemblages
Based on the work of Overpeck et al. (1985), we used
the squared chord distance (SCD) dissimilarity coefficient to reconstruct July temperatures from fossil pollen
assemblages. We choose two different SCD critical
values, one for the analog approach and one for the
response surface approach.
Previous modern analog reconstructions in eastern
North America demonstrated that a SCD value of 0.15
could differentiate between analog and no-analog pollen
assemblages (Overpeck et al., 1992). To evaluate how a
SCD critical value of 0.15 would work in our study area,
we conducted a statistical comparison of modern pollen
samples both from the same vegetation zone and from
different vegetation zones (Supplementary Table 4).
Sixty three percent of all samples on average within each
of the five vegetation zones fell below 0.15, but more
importantly, only 11% of the samples were considered
numerically similar (o0.15) even though they were from
different vegetation zones (Supplementary Table 4). We
considered using 0.18 SCD (i.e., Anderson et al., 1989)
as the critical value for the modern analog approach
because 70% of all samples from the same vegetation
zone fell below 0.18, but choose to use 0.15 as a
conservative estimate for separating analog assemblages
from no-analog assemblages.
For the response surface approach we selected a
critical value of 0.20 SCD following the guidelines
established for eastern North America by Webb et al.
(1987, 1993b, 1998). Webb and his group used 0.20 SCD
as a cutoff because more than 97% of the observed
spectra matched the response surface spectra below 0.20
SCD (Webb et al., 1993a). We evaluated our dataset in
the same manner and found that 99% of the modern
and 90% of fossil pollen assemblages were numerically
Two-dimensional cross sections of the three-dimensional response surfaces illustrate the relationship
between pollen abundance and climate for the 11 pollen
taxa (Alnus was excluded) included in the sum for Baffin
Island and the 12 pollen types included in the sum for
Quebec and Labrador (Fig. 3). Collectively the response
surfaces reveal how individual species are controlled by
climatic variations, and help identify which pollen types
are useful as indicator species for specific climate zones.
The response surface for Picea pollen shows maximum
percentages in the relatively warm and wet portions of
the study area coinciding with the boreal forest where
Picea mariana is the dominant tree species. High Picea
percentages also cluster in parts of the subarctic climate
space where Picea glauca is common and prevailing
southwesterly winds bring Picea mariana pollen from
the nearby forest. Maximum Abies pollen abundances
are associated with the warmest and wettest climate
space of the broadleaf boreal forest where there is a high
concentration of Abies balsamea. Oxyria, Saxifraga, and
the Arctic herbs group are most abundant in the coldest
and driest climate space in the study area associated
with the mid- to high-arctic.
The Betula and Alnus response surfaces are somewhat
complicated by the tree versus shrub control on pollen
concentrations throughout eastern Canada, and the
tendency for pollen grains to be transported large
distances by the wind. Nonetheless, maximum Betula
and Alnus pollen percentages are clustered toward the
middle of the climate space coinciding with the subarctic
vegetation zone where Alnus crispa and Betula glandulosa are co-dominant shrubs (Fig. 3). Percentages
decline in both directions toward the warm, wet
broadleaf boreal forest (although both Alnus and Betula
percentages increase at the most warm and wet corner of
the climate space, where Alnus rugosa and Betula
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Cyperaceae
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Poaceae
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Salix
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0
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(a)
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Annual Precipitation (mm)
Arctic Herbs
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Saxifraga
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>6.0
1500 0
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Annual Precipitation (mm)
3
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>4
1500 0
500
1000
1500
Annual Precipitation (mm)
Fig. 3. Response surfaces (versus July temperature and annual precipitation) based on 11 pollen taxa (A) and 12 pollen taxa (B) showing the impact
of including Alnus in the pollen sum.
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30
20
12
15
60
10
July Temperature (ºC)
15
20
15
60
18
Betula
0
5
10
15
40
20
20
>40
>20
>20
0
18
5
20
35
8
6
16
14
12
5
10
3
1
2
8
0
5
10
15
20
>20
Poaceae
3
0
3
6
8
4
25
0
7
Cyperaceae
7
6
5 4
40
Ericaceae
2
5
5
30
2
12
6 14
16
3
4
6
10
8
6
40
4
9
1
10
15
25
35
15
10
2
12
10
July Temperature (ºC)
15
15
20
>20
9
12
>12
0
15
3
0
0
2
0
1
12
1
2
3
9
3
6
3
6
3
2
2 1
Artemisia
0
5
10
0
1
2
0
1
2
15
20
3
4
3
4
>20
>4
>4
3
3
Abies
1
4
12
9
6
0
Salix
2
3
9
12
July Temperature (ºC)
0
1
2
3
0
18
0
18
0
0
3
4.5
6.0
>6.0
0
500
1000
Annual Precipitation (mm)
0
1
1.5
3.0
3
1.5
3.0
Arctic Herbs
0
1
0
2
3
0
Saxifraga
0
Oxyria
1
4
2
0
3
6
1
3
2
3
(b)
2
0
1
1
9
0
1
2
2
12
2
July Temperature (ºC)
15
4.5
6.0
>6.0
1500 0
500
1000
Annual Precipitation (mm)
Fig. 3 (continued).
3
4
>4
1500 0
500
1000
Annual Precipitation (mm)
1500
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
papyrifera species are prevalent), and toward the cold,
dry high-arctic (where these taxa are entirely exotic).
The response surfaces for the ubiquitous taxa including
Ericaceae, Cyperaceae, Poaceae, and Salix suggest that
these pollen taxa are dominant throughout the tundra
regions of the study area. Variations in Cyperaceae and
Poaceae pollen should be useful for differentiating the
climate space of tundra from that of the forest.
Variations in Ericaceae, and Salix pollen can be used
to differentiate climate zones within the tundra. Salix
peaks in the climate space of the mid- to high arcitc,
whereas Ericaceae is abundant in all tundra climate
zones except the subarctic climate space of N Quebec
and Labrador. The response surface for Artemisia is
non-monotonic, showing maximum percentages in the
climate space coinciding with the low-arctic vegetation
zone on Baffin Island.
5.2. Comparison of reconstruction approaches
Throughout much of the Holocene at most of the 25
sites, the July temperature reconstructions are broadly
similar between the RS and AN approaches (Fig. 4 and
Supplementary Table 6), suggesting that both methods
are capable of picking up the large scale and subtle
fluctuations in fossil pollen records from eastern
Canada. Discrepancies between the two techniques are
most apparent in the records from six sites in the boreal
forest and forest–tundra regions of Quebec and Labrador (Bereziuk, CHISM 2, Delorme, GB2, Gravel Ridge,
and Hebron). At these sites the RS reconstructions show
high sample-to-sample variability (1–2 C) that is inconsistent with the AN reconstructions and the fossil
pollen assemblage data (Kerwin, 2000). This variability
can be attributed to the selection of averaging
and smoothing algorithms in constructing response
surfaces (Kerwin, 2000). The RS fitting procedure
smooths the modern pollen data, removing much of
the spatial variability inherent to the raw data. In
previous studies from eastern North America, this
spatial variability was considered unrelated to climate
and hence noise (Bartlein et al., 1986; Bartlein and
Whitlock, 1993; Webb et al., 1993a). In our much
smaller study area, this smoothing appears to produce
inconsistent results, within a 1–2 C July temperature
window, because many modern pollen assemblages from
the forest–tundra climate space are averaged together
(see Fig. 4D, Gravel Ridge for an example). Despite
these areas of mismatch between the two methods, the
overall pattern of Holocene July temperature fluctuations in eastern Canada is generally consistent between
the methods. Thus, rather than limiting our reconstructions to one method, we utilize the results from both the
RS and AN approaches to maximize the statistical
robustness of the reconstructed values (Webb and
Woodhouse, 2003).
5.3. Consolidation of individual lake records
We consolidated the 25 individual records into 10
‘‘regional records’’ by averaging the paleoclimate
reconstructions from nearby lakes (Fig. 5). Lakes in
close proximity (and elevation) to each other were
assumed to have experienced similar Holocene environmental and climatic fluctuations. The ten regional
syntheses reveal coherent, regional climate fluctuations
recorded at more than one site, and de-emphasized
individual variations from one lake that were not
recorded elsewhere. Two regions in Baffin Island were
represented by single fossil pollen records: NW Cumberland Sound and Clyde River.
The W Quebec regional record (near James Bay) is
based on three lakes (Bereziuk, CHISM 2, Kanaaupscow) that are currently in the boreal forest and were in
close proximity to several of the large proglacial lakes
and ice sheets during deglaciation (Andrews, 1987; Dyke
and Prest, 1987; Richard, 1995). These lakes probably
experienced large climate fluctuations associated with
episodic melting of ice sheets during the early Holocene
(Barber et al., 1999). Three regional records (10 lakes)
are clustered in central Quebec and NW Labrador along
the forest–tundra transition zone (Short and Nichols,
1977; Stravers, 1981; Richard et al., 1982; Lamb, 1985;
Gajewski and Garralla, 1992; Gajewski et al., 1993).
From west to east these regional records include 3 lakes
(BI2, GB2, LB1) along the eastern margin of Hudson
Bay, 4 lakes (Boundary, Delorme, Lac Hamard,
Tunturi) near the eastern Quebec/W Labrador border,
and 3 lakes (Caribou Hill, Gravel Ridge, Kogaluk) in
NW Labrador. Cumulatively, these 10 lakes provided
the best opportunity to quantify summer temperatures
along the southern margin of the Quebec–Labrador ice
sheet that persisted until approximately 6000 14C yr BP
(COHMAP, 1988; Richard, 1995). Further to the
northeast, one group of 3 lakes (Nain, Hebron, Ublik)
is centered in NE Labrador in close proximity to the
Labrador Sea (Short and Nichols, 1977; Lamb, 1984).
During deglaciation, the Torngat Mountains separated
these Labrador lakes from the last remnants of the
Laurentide Ice Sheet making them more susceptible to
oceanic conditions in the North Atlantic. The final
regional record on mainland Canada consists of two
lakes (Diana 375, VHC1) in N Quebec in close
proximity to Hudson Strait (Richard, 1977, 1981). On
Baffin Island, one regional record of two lakes (Hikwa,
Jake) is located just north of Frobisher Bay in the
middle of the low-arctic vegetation zone (Mode and
Jacobs, 1987). A second regional record of three lakes
(Donard, Dyer, Fog) in located on NE Cumberland
Peninsula in the mid-arctic vegetation zone (Wolfe et al.,
2000; Miller et al., submitted for publication). The other
two regional records are individual lake records from
NW Cumberland Sound (Iglutalik) on the boundary
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
between the low- and mid-arctic vegetation zones and
Clyde River (Patricia Bay) in the high-arctic vegetation
zone.
5.4. Holocene July temperature reconstructions
Sparse fossil pollen data and frequent no-analog
assemblages restricted quantitative reconstructions before 8000 14C yr BP. After 8000 14C yr BP, most of the
fossil pollen samples have modern analogs (Fig. 4). In
general tundra regions in Baffin Island and Quebec
reached or exceeded today’s July temperatures earlier
than the forest and forest–tundra regions in Quebec and
Labrador (Fig. 5). Eleven of the 12 records from modern
day tundra regions suggest that July temperatures ca.
6000 14C yr BP were equal to or greater than those of
today (Fig. 6). In contrast, all 13 records from the boreal
forest and forest–tundra transition zone of non-coastal
Quebec and Labrador suggest that July temperatures
were 1–2 C colder than present ca 6000 14C yr BP. All of
the lakes that show colder than present conditions ca
6000 14C yr BP were in close proximity to the last
remnants of the Laurentide Ice Sheet that persisted in
the Quebec–Labrador Peninsula until approximately
6000 14C yr BP (COHMAP, 1988; Richard, 1995). The
lakes that show no change or warmer conditions ca 6000
14
C yr BP were distant from this ice sheet and therefore
influenced by the 5% higher than present summertime
isolation forcing at 60 N latitude (Berger, 1978) and the
warmer than present sea surface temperatures and
reduced sea-ice in Hudson Strait, the Labrador Sea,
and Davis Strait (Andrews, 1972; Dyke et al., 1996;
Kerwin et al., 1999; Sawada et al., 1999; Gajewski et al.,
2000).
The single record from Patricia Bay Lake (NE Baffin
Island) suggests the maximum regional Holocene
warming (+1 C, relative to today) occurred shortly
after deglaciation (6500 14C yr BP) in the high arctic.
Gradual cooling ensued over the next 3000 years (the
remainder of the record), in agreement with a previous
quantitative July temperature reconstruction based on
this fossil pollen record (Diaz et al., 1989). Further
south along the east coast of Baffin Island, three records
from NE Cumberland Peninsula (Dyer, Donard, and
Fog) suggest that July temperatures peaked (+1 C) at
5000 14C yr BP. This maximum regional Holocene
warming is not recorded in the Iglutalik Lake record,
which shows little temperature variability and no
significant warming during the mid- to late Holocene
(Fig. 4). The reconstruction from Hikwa Lake in
southern Baffin Island suggests that maximum regional
Holocene warming was delayed (peak: 4200–3200
14
C yr BP) and slightly larger in magnitude (+2 C) than
elsewhere on Baffin Island.
In the low-arctic tundra of N Quebec along Hudson
Strait and in the subarctic tundra of NE Labrador along
1911
the Labrador Sea, July temperatures peaked at about
3500 14C yr BP (+1.5 C to 2 C) and were beginning to
decline by 3200 14C yr BP. July temperatures in W
Labrador were about 1 C warmer than present by 4500
14
C yr BP and peaked at about 2000 14C yr BP (+1.5 C).
This particular anomaly at 2000 14C yr BP may be
somewhat misleading since it is driven by an increase in
Picea pollen from 60% to 70% during a time when the
overall pollen record suggests the forest was beginning
to thin (Lamb, 1985). Three forest tundra sites along the
east coast of James Bay in W Quebec suggest minimal
regional Holocene warming (+0.5 C) that lingered
from about 5500 to 2000 14C yr BP. Along the N boreal
forest and forest–tundra transition zone of Quebec,
regional Holocene warming was even less pronounced, if
it existed at all. Summer temperatures warmed to near
present values within 500–1000 years after 6000
14
C yr BP, and remained relatively unchanged and close
to modern values over the remainder of the Holocene
(Fig. 5).
July temperatures declined throughout the tundra
regions of eastern Canada after 4000 14C yr BP, with the
most pronounced cooling (>1 C over 1000 years) in the
coastal regions of southern Baffin Island and NE
Labrador at 3200 14C yr BP (Fig. 5). Although still
apparent the cooling trend after 4000 14C yr BP was
more gradual in E and NE Baffin Island (o0.5 C
over 1500 years). July temperatures declined further
(0.7–1.5 C) beginning about 2500 14C yr BP in E and S
Baffin Island based on the individual records from
Hikwa and Dyer Lakes (Fig. 4). Neoglacial cooling is
less apparent in the forest and forest–tundra sites.
6. Discussion: comparison with other Paleoenvironmental
and Paleoclimatic records
6.1. Quebec and Labrador
6.1.1. 6000 14C yr BP
Prior to 6000 14C yr BP, an ice sheet persisted in the
Quebec–Labrador Peninsula that strongly influenced
the regional climate (COHMAP, 1988; Richard, 1995).
The spatial extent of this ice sheet’s influence is evident
when looking at our July temperature anomalies at 6000
14
C yr BP (Fig. 6). All 13 sites that were in close
proximity to the ice sheet remained colder than present
at 6000 14C yr BP while July temperatures in the rest of
the region had exceeded or at least reached modern
values. Our results for 6000 14C yr BP are consistent with
the reconstructions of Bartlein and Webb (1985), Diaz
et al. (1989) and Webb et al. (1993a), which both suggest
a 1–2 C summer temperature cooling at 6000 14C yr BP
relative to today. Our results disagree with the no
change or ‘‘warmer than present’’ summer temperature
anomalies for 6000 14C yr BP reconstructed by Andrews
1912
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
Fig. 4. July temperature reconstructions based on the response surface (blue diamonds) and modern analog approaches (red circles) from 25 fossil pollen records in NE Canada. Uncertainty in the
reconstructed climatic values is shown as the stippled line bracketing the reconstructed values for each technique, and was estimated by the standard deviation of the mean of the (weighted) climate
values associated with the ten best analogs. The vertical black line in each plot represents the mean modern (1951–1980) July temperature at each site. The time series to the right of each
reconstruction shows the squared-chord distance values for the single closest (stippled line) and average of the top ten closest modern analogs for the response surface (blue lines) and modern analog
(red lines) approaches. The two vertical lines in the squared-chord distance time series show the critical values (0.15 for modern analog approach and 0.20 for response surface approach) used to
identify fossil pollen assemblages that have no analog in our modern pollen database. No-analog assemblages were excluded from the reconstructions and are not included in these reconstructions.
ARTICLE IN PRESS
1913
Fig. 4 (continued).
M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
Fig. 4 (continued).
1914
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
Fig. 4 (continued).
1915
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1916
M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
and Diaz (1981), Gajewski et al. (2000) and Sawada et al.
(1999). Climate model simulations suggest that an ice
sheet in the Quebec–Labrador Peninsula could have
spawned summertime anticyclones that would have
enhanced northerly air flow, creating localized cooling
south of the ice sheet (COHMAP, 1988).
Away from the influence of the Quebec–Labrador ice
sheet, our three records from NE Labrador and two
from N Quebec suggest that July temperatures at 6000
14
C yr BP were generally comparable to those of today
(Fig. 6). The five sites were influenced primarily by
oceanic conditions in the Labrador Sea and Hudson
Strait. Summer sea surface temperature (SST) reconstructions are limited in these regions, but marine
mollusk and dinoflagellate cyst data from the Labrador
Sea and Hudson Bay suggest that SSTs at 6000
14
C yr BP were slightly warmer than present (Andrews,
1972, 1973; Levac and de Vernal, 1997; Andrews et al.,
1999; Kerwin et al., 1999; Sawada et al., 1999).
6.1.2. Late Holocene (5000-1 14C yr BP)
Between 5000 and 4000 14C yr BP, SSTs in the
Labrador Sea and Hudson Strait began to cool (Dyke
et al., 1996) which, combined with a decrease in summer
insolation (Berger, 1978), likely contributed to the late
Holocene summer cooling trend that is apparent in all
five of our records from NE Labrador and N Quebec
between 4600 and 3000 14C yr BP (Fig. 4). This late
Holocene cooling trend was apparently restricted to the
N and NE coastal regions because the interior records
show no cooling until after 2000 14C yr BP interval
(Fig. 6). One possible explanation for the lack of a late
Holocene cooling trend in central Quebec and Labrador
is that the mean position of the polar front may have
been shifted north of its present position during the late
Holocene (5000–2000 14C yr BP). The position of the
polar front in mid-summer tends to follow the northern
boreal forest, and has a large impact on mean summer
temperatures at the forest–tundra boundary (Bryson,
1966). The polar front today inhibits cold northern air
from settling in the regions south of the front. Because
the polar front tends to accurately reflect tree line during
the midsummer, it is likely that its mean position
migrated north during the middle Holocene as tree line
expanded (Payette and Lavoie, 1994). In turn the
position of the polar front likely migrated to the south
1917
as the boreal forest began to thin at about 3000
C yr BP in N Quebec (Richard, 1995) and Labrador
(Lamb, 1985). The lack of a cooling trend in the boreal
and forest–tundra records until 2000 14C yr BP ( 1 C
over 1000 years) in Caribou Hill and Gravel Ridge
Lakes (Labrador) and at 1000 14C yr BP ( 0.5 C in 500
years) in Lakes B12 and LB1 (Quebec) suggests that the
polar front remained north of these sites until 2000 and
1000 14C yr BP, respectively, when the front moved to
the south subjecting these sites to more frequent
summertime Arctic air masses.
14
6.2. Eastern Canadian Arctic
6.2.1. Early postglacial (11,000–7000 14C yr BP)
Deglaciation was underway by 11,000 14C yr BP at
several E Baffin Island locations (Miller et al., 1999;
Wolfe et al., 2000; Moore et al., 2001). The pollen record
from Donard Lake shows an influx of herbaceous pollen
(Rosaceae, Fabaceae, Saxifraga, etc.), Poaceae
(grasses), pteridophytes (ferns, club-moss, etc.), and
Oxyria from 11,000 to 8000 14C yr BP (Miller et al.,
submitted). Similar pollen assemblages are found in
early Holocene sediments from Dyer, Fog, Amakutalik,
Jake, Iglutalik, and Robinson Lakes on southern and
central Baffin Island (Short et al., 1985; Mode and
Jacobs, 1987; Miller et al., 1999; Wolfe et al., 2000).
These early Holocene pollen records commonly contain
percentages of Poaceae and Oxyria exceeding 50%
(max.=99%) and 30% (max.=60%), respectively.
Oxyria digyna and Poaceae plants are dominant species
in newly emergent tundra vegetation (Porsild, 1964).
Almost all of the fossil pollen samples on Baffin Island
prior to 8000 14C yr BP have no modern analogs (SCD
values >0.15?) when compared with the 275 modern
pollen samples in our database, even those from high
Arctic, ice proximal sites on N Baffin Island. As a result,
we did not reconstruct July temperatures on Baffin
Island prior to 8000 14C yr BP.
The early Holocene vegetation on Baffin Island that
produced the no-analog pollen assemblages was likely
similar to the modern vegetation at some of our higharctic lakes, but the pollen deposition 8000–11,000 years
ago was different then compared to now. Exotic pollen
is a large component of high-arctic lakes today (Short
et al., 1994; Kerwin et al., in preparation), with grains
Fig. 5. Composite July temperature anomalies (reconstructed minus modern) based upon 25 fossil pollen records using the modern analog (red
circles) and response surface (blue diamond) reconstruction approaches. Individual lake records were merged into single, composite records for each
region by first calculating July temperature anomalies by subtracting the observed mean July temperatures (1951–1980) from the reconstructed July
temperature values shown in Fig. 4. We then averaged the RS and AN July temperature anomalies for each fossil pollen sample (Bartlein and
Whitlock, 1993). We used a simple smoothing algorithm (second order distance weighted least-squares technique) on the averaged climate anomalies
to reduce the variance in the time series (Velleman, 1980). We plotted all of these data, including the raw temperature anomalies, averaged values
(stippled black line), and the smoothed composite reconstruction (dark black line), for each lake record to show the variability between the two
reconstruction techniques. Note that the Patricia Bay Lake and Iglutalik Lake reconstructions (dark black lines) simply reflect the averages between
the modern analog and response surface reconstructions.
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
reaching Baffin Island from all over central and N
North America (Ritchie and Lichti-Federovich, 1967).
From 11,000 to 8000 14C yr BP, much of Canada was
still unvegetated or covered by sparse tundra species
rather than the dense forest or shrub vegetation that
persists today (Ritchie, 1987). The source vegetation for
much of the exotic pollen that today reaches the high
Canadian Arctic probably did not exist in the early
Holocene. The jet stream was also shifted south of its
present position because of the influence of the
Laurentide Ice Sheet (COHMAP, 1988). These two
factors combined to diminish the flux of exotic pollen to
Baffin Island lakes, thereby increasing the concentration
of local taxa like Poaceae and Oxyria. Development of
pollen climate calibrations based on pollen concentration data (Elliot-Fisk et al., 1982; Short et al., 1985) may
eliminate the problems related to no-analog assemblages, but is difficult due to variations in sedimentation
and accumulation rates from changes in pollen influx.
Excluding exotic pollen taxa from the sum (Picea,
Betula, Salix) may also work.
By 8000 14C yr BP, the local vegetation had begun to
change from an emergent tundra or rock barrens to a
shrub tundra dominated by Salix and Cyperaceae with
some Ericaceae plants (Mode, 1996). This pollen and
vegetation shift coincided with a time when summertime
insolation at 60 N latitude was approximately 7.5%
higher than it is today (Berger, 1978) resulting in warmer
than present summer temperatures across much of the
Northern Hemisphere (Kerwin et al., 1999). Our July
temperature reconstructions for 8000 14C yr BP are
limited to Donard and Fog Lakes where a slight
(+0.5 C) increase relative to today is apparent
from 8000 to 7000 14C yr BP (Fig. 4), probably
associated with final deglaciation in S Baffin Island
(Stravers et al., 1992).
6.2.2. Middle Holocene warming (7000–4000 14C yr BP)
Our July temperature reconstructions from Baffin
Island suggest that the interval from 7000 to 4000
14
C yr BP ka was initially a time of rising temperatures
followed by a consistently warmer than present interval
(by 1–2 C) that peaked at various times (earlier in N
Baffin Island) (Fig. 5). By 6000 14C yr BP, positive
anomalies of 0.5 C near Frobisher Bay, 0.75 C on NE
Cumberland Peninsula, and 1 C near Clyde River (Fig.
6) were reconstructed. Summer warming of 0.5–1 C at
1919
6000 14C yr BP is consistent with marine mollusk data
from the east coast of Baffin Island that suggest SSTs in
W Baffin Bay and the Davis Strait warmer than modern
at 6000 14C yr BP (Dyke et al., 1996; Levac et al., 2001).
Likewise, d18O records from cellulose, extracted from
aquatic mosses preserved in Baffin Island lake sediments
and ice core data from the high-Canadian Arctic,
suggest that summer temperatures across the Eastern
Canadian Arctic were warmer than present at 6000
14
C yr BP (Sauer, 1997; CAPE, 2001).
In southern Baffin Island, an abrupt warming of
> 2 C over about 500 years is recorded in the Hikwa
Lake pollen record beginning about 4500 14C yr BP. This
warming trend coincides with an sudden increase in
Alnus and Betula pollen in four southern Baffin Island
lakes (Miller et al., 1999). This sharp rise in Alnus and
Betula pollen has been interpreted to reflect an increase
in warm, southerly winds that brought pollen across
Hudson Strait from the recently colonized shrub tundra
of Ungava-Nouveau Quebec (Andrews et al., 1981), and
an expansion of local Betula dominated low-arctic
vegetation (Short et al., 1985). An outstanding question
is did the southerly winds increase local temperatures by
3 C during the peak of middle Holocene warming or
simply blow pollen assemblages associated with more
southerly (warmer) climates to Hikwa Lake?
6.2.3. Neoglaciation (5000–2000 14C yr BP)
Evolution of the Neoglaciation was most likely in
response to cooler summer temperatures on Baffin
Island (e.g., Denton and Porter, 1970; Miller, 1973).
Summer cooling began as early as 6000 14C yr BP in NE
Baffin Island (Patricia Bay Lake: 1 C over 3000 years)
and as late 3200 14C yr BP in S Baffin Island (Hikwa
Lake: 1.5 C in 1000 years). Elsewhere July temperatures began to decline ( 0.5–1 C over 2000–3000 years)
between 5000 and 4000 14C yr BP (Fog, Dyer, Donard,
Jake, and Iglutalik Lakes). Pollen records reveal a shift
during this time from being dominated by Alnus and
Betula to containing high percentages of Cyperaceae,
Salix and Ericaceae. Pollen concentration declined at
this time as well (Mode, 1986). A later cooling trend is
also apparent the Dyer Lake record, which shows a
cooling of 0.5–1 C over 1000 years beginning about
2500 14C yr BP.
Our results are compatible with other paleoclimate
and paleoenvironmental records from Baffin Island.
Fig. 6. July temperature anomalies (reconstructed minus modern values) at each fossil pollen site from 6000 to 1000 14C yr BP. Anomalies were
calculated for each sample by first subtracting the mean modern July temperature (1951–1980) at each lake from the reconstructed temperature
derived from the two different reconstruction techniques shown in Fig. 4. We then averaged the two independent July temperature anomalies
(response surface and modern analog) that were generated for each fossil pollen sample. We used a simple smoothing algorithm (second order
distance weighted least-squares technique) on the averaged climate anomalies to reduce the variance in the time series (Velleman, 1980). This
smoothed line was then used as the composite July temperature reconstruction for each site. We generated the above paleoclimate anomalies by using
the single anomaly closest to each 1000 year interval as the reconstructed temperature for that time slice.
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M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
Lake sediment d18O measurements on aquatic mosses
suggest an abrupt cooling in the Cumberland Sound
area at about 5000 14C yr BP and again at 2400 14C yr BP
(Sauer, 1997). Glaciers expanded on Baffin Island
between 5000 and 2000 14C yr BP (Miller, 1973; Davis,
1985), with the greatest increase in glacier activity
recorded at Donard Lake at 2300 14C yr BP (Moore
et al., 2001). SSTs in Baffin Bay and the Labrador Sea
cooled from 3000 to 2000 14C yr BP as recorded by the
shift from subarctic molluscan fauna to cold-water
tolerant species (Dyke et al., 1996). The diatom record
from Fog Lake also shows a 4–5 C cooling in lake water
temperatures beginning at approximately 2500 14C yr BP
(Wolfe et al., 2000). The Devon Ice Cap record from just
north of Baffin Island shows a 2 per mil decrease in d18O
between 5000 and 4000 14C yr BP that has been
correlated to decreasing Arctic temperatures (Paterson
et al., 1977). Together these records clearly show two
major cooling events in the eastern Canadian Arctic; one
beginning between 5000 and 4000 14C yr BP and the
second centered around 2500 14C yr BP.
7. Conclusions
1. Both the method of modern analog and the response
surface approaches can be used to reconstruct
millennial scale changes in July temperatures from
the boreal, subarctic, and Arctic regions of eastern
Canada. Because of the shallow climate gradient in
the Arctic, we used two different response surfaces
for paleoclimate reconstructions and reduced the
search window/eliminated the post-fitting smoothing
algorithms. These alterations resulted in more realistic response surfaces while bringing the response
surface approach closer to the modern analog
approach.
2. Our July temperature reconstructions in the Arctic
are consistent with existing paleoclimate and paleoenvironmental records suggesting that Arctic
pollen dispersal and deposition in lake sediments
are sensitive to July temperature variations. Arctic
lakes are strongly impacted by the irregular accumulation of wind-blown pollen (Barry et al., 1981) and
sediments (Moore et al., 2001), in addition to local
variations in elevation, topography, soil development, and moisture that impact the development of
tundra vegetation. Our reconstructed July temperatures thus likely assimilate several different environmental factors including changes in wind direction,
storm tracks, sedimentation rates, and winter snow
accumulation that are difficult to quantify. The
pollen record (Mode and Jacobs, 1987) and paleoclimate reconstruction from Hikwa Lake in S Baffin
Island illustrate some of the challenges of extracting
quantitative paleoclimate information from Arctic
pollen records. July temperatures increase abruptly at
about 5000 14C yr BP (2 C over 500 years) coincident
with an increase in Betula pollen (Alnus was removed
from the analysis) percentages (Mode and Jacobs,
1987). Betula shrubs are common at Hikwa Lake
today and sensitive to climate changes in the lowarctic vegetation zone of Baffin Island (Andrews et al.,
1980b; Jacobs et al., 1985). Some of the Betula pollen
in Hikwa Lake is exotic having blown in with Alnus
pollen from the subarctic vegetation zone to the south
and west of Hikwa Lake (Short et al., 1985; Mode
and Jacobs, 1987). Unlike Alnus pollen, Betula
cannot be removed from the pollen sum, because
local Betula shrubs grow at the Hikwa Lake site. At
the same time, enhanced southerly winds capable of
blowing Betula pollen from Quebec and Labrador
would have brought warmer summer temperatures to
the Hikwa Lake region (Andrews et al., 1981).
Despite some of these problems, pollen derived
climate records from Baffin Island do appear useful
for studying millennial scale summer temperature
fluctuations.
3. July temperatures at 6000 14C yr BP in Quebec and W
Labrador were 1–2 C colder than present because of
residual Laurentide Ice in the Quebec–Labrador
Peninsula (Richard, 1995), in agreement with previous pollen-based summer temperature reconstructions by Bartlein and Webb (1985) and Webb et al.
(1993a). In contrast, summer temperatures at 6000
14
C yr BP were 0.5–1 C warmer than present in Baffin
Island and NE Labrador because of increased orbital
forcing and warmer than present Atlantic sea surface
temperatures.
4. Neoglaciation was more pronounced in the Arctic
compared to eastern mainland Canada where the
mean position of the polar front sheltered the sites to
the south of the front from cold, Arctic air in the
summertime. Gradual summer temperature cooling
beginning as early as 6000 14C yr BP in NE Baffin
Island and intensified at about 4000 14C yr BP in NE
Labrador and south-central Baffin Island. Decreased
summertime insolation (Berger, 1978) and SST cooling in Baffin Bay and the Labrador Sea from 3000 to
2000 14C yr BP resulted in a 0.5–1 C July temperature
cooling at about 2500 14C yr BP. The fact that the
vegetation responded to cooling at 4000 14C yr BP,
well before the decrease in SST, suggests that
decreased solar insolation first caused the Neoglaciation at about 4000 14C yr BP, which was subsequently
intensified at about 2500 14C yr BP by decreased SSTs
in Baffin Bay and the Labrador Sea.
5. Our Holocene reconstructions reveal a mosaic of
transient warming and cooling trends that are
dependent on local manifestation of changes in the
seasonal distribution of insolation and high latitude
SSTs. The global change message here is that uniform
ARTICLE IN PRESS
M.W. Kerwin et al. / Quaternary Science Reviews 23 (2004) 1901–1924
and synchronous changes in summer temperature in
the high Arctic across distances of less than 300 km
are not ubiquitous in the Holocene. Two alternative
approaches of how to interpret this are: (1) the recent
relatively synchronous and uniform widespread
changes in the Arctic are unprecedented, (2) any
future changes in the Arctic should not be expected
to be either relatively synchronous, uniform, or
widespread.
Acknowledgements
This research was supported by a grant from the US
National Science Foundation (NSF-ATM-9402657). We
thank Gifford Miller and Vera Markgraf for comments
and discussions during the preparation of this manuscript, and Joel Guiot and Brian Huntley for careful
reviews that greatly improved the paper.
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