A late Quaternary paleotemperature record from Hanging Lake, northern Yukon Territory, eastern Beringia

Quaternary Research 72 (2009) 246–257

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Quaternary Research j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / y q r e s

A late Quaternary paleotemperature record from Hanging Lake, northern Yukon Territory, eastern Beringia Joshua Kurek a,⁎, Les C. Cwynar a, Jesse C. Vermaire b a b

Department of Biology, University of New Brunswick, PO Box 4400, Fredericton, NB, E3B 5A3 Canada Department of Biology, McGill University, Canada

a r t i c l e

i n f o

Article history: Received 16 May 2008 Available online 12 June 2009 Keywords: Paleoclimate Chironomids Last glaciation Younger Dryas stade Holocene thermal maximum Pollen Lake sediments Beringia

a b s t r a c t The late Quaternary paleoclimate of eastern Beringia has primarily been studied by drawing qualitative inferences from vegetation shifts. To quantitatively reconstruct summer temperatures, we analyzed lake sediments for fossil chironomids, and additionally we analyzed the sediments for fossil pollen and organic carbon content. A comparison with the δ18O record from Greenland indicates that the general climatic development of the region throughout the last glaciation–Holocene transition differed from that of the North Atlantic region. Between ∼ 17 and 15 ka, mean July air temperature was on average 5°C colder than modern, albeit a period of near-modern temperature at ∼ 16.5 ka. Total pollen accumulation rates ranged between ∼ 180 and 1200 grains cm− 2 yr− 1. At ∼ 15 ka, approximately coeval with the Bølling interstadial, temperatures again reached modern values. At ∼ 14 ka, nearly 1000 yr after warming began, Betula pollen percentages increased substantially and mark the transition to shrub-dominated pollen contributors. Chironomid-based inferences suggest no evidence of the Younger Dryas stade and only subtle evidence of an early Holocene thermal maximum, as temperatures from ∼ 15 ka to the late Holocene were relatively stable. The most recognizable climatic oscillation of the Holocene occurred from ∼ 4.5 to 2 ka. © 2009 University of Washington. Published by Elsevier Inc. All rights reserved.

Introduction During the last glaciation–Holocene transition, between 17 and 10 ka (calibrated thousand years before present), significant changes in vegetation composition, insolation, ice-sheet extent, and atmospheric circulation patterns occurred in eastern Beringia (Bartlein et al., 1991; Edwards et al., 2001; Bigelow et al., 2003; Kaplan et al., 2003). This transitional period is crucial for understanding: 1) the environmental context of human colonization into North America (Dixon, 2001; Marshall, 2001), 2) the extirpation and expansion of dominant large mammals in eastern Beringia (Guthrie, 2006), and 3) broad-scale vegetation responses to changing environmental conditions. However, there are few continuous, quantitative reconstructions of climate from the region (see Bunbury and Gajewski, 2009; Kurek et al., 2009). The paucity of suitable climatic records also makes it a challenge to evaluate potential teleconnections with ocean–atmosphere processes in the North Atlantic region known to have hemispheric or perhaps even global significance. Abrupt, highmagnitude climatic oscillations well-recorded in the North Atlantic region, such as the Younger Dryas stade (∼13 to 11.6 ka), are inadequately understood from inland eastern Beringia and sites often show differing climatic responses (Kokorowski et al., 2008). ⁎ Corresponding author. E-mail addresses: [email protected] (J. Kurek), [email protected] (L.C. Cwynar), [email protected] (J.C. Vermaire).

Documenting the extent, magnitude, and timing of climatic oscillations is a necessary step towards recognizing mechanisms that can explain the global climate system. Climatic conditions varied during the last glaciation–Holocene transition across eastern Beringia, as easternmost Beringia (interior Alaska and Yukon Territory) was colder and drier than regions adjacent to the Bering Land Bridge (Barnosky et al., 1987; Guthrie, 2001). Between ∼14 and 13 ka, some sites in northwestern Alaska record an earlier thermal maximum than those in northwestern Canada (Kaufman et al., 2004). During the early Holocene, Picea populations in northwestern Canada exceeded their present northern and altitudinal range limits (Cwynar and Spear, 1991, 1993; Szeicz and MacDonald, 2001), whereas Picea populations in northwestern Alaska expanded, but never extended past their present day northern range limit (Anderson and Brubaker, 1994). Independent, high-resolution climatic reconstructions may help to improve the understanding of the region's paleoclimatic development and its effect on sub-Arctic vegetation. Here we present a paleoclimate reconstruction based on fossil chironomids from Hanging Lake, northern Yukon Territory. We assess patterns in summer temperatures during the last glaciation–Holocene transition, comparing our inferences with pollen-derived vegetation change at Hanging Lake and the GISP2 δ18O record. This study addresses the following questions: 1) What are the magnitudes and duration of the Younger Dryas stade and Holocene thermal maximum at Hanging Lake? 2) To what extent do regional changes in vegetation coincide with

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J. Kurek et al. / Quaternary Research 72 (2009) 246–257

inferred changes in temperature between ∼17 and 12 ka? and 3) Are inferred temperatures from the last glaciation–Holocene transition synchronous with those of the North Atlantic region? The vegetation history from Hanging Lake (Cwynar and Ritchie, 1980; Cwynar, 1982) ignited a debate about the nature of vegetation during the last glaciation in Beringia. The prevailing view had been that Beringian vegetation was unique and generally indicative of productive grassland (Guthrie, 1968; Matthews, 1982), but low pollen influx during the last glaciation at Hanging Lake suggested a sparsely vegetated landscape more like Arctic herb communities than temperate grassland communities. This led to the “productivity paradox” (Hopkins et al., 1982), namely, how could the presumptive large herds of megafauna grazers (bison, horse, and woolly mammoth) be supported in an unproductive Arctic herb tundra? That debate endures to some extent, although there is now a general acceptance that Beringia was a heterogeneous landscape and not a monolith of any particular vegetation type during the last glaciation. From this historical perspective, Hanging Lake continues to draw attention. The original chronology for Hanging Lake was based on conventional dating of bulk sediments. The new cores and subsequent chronology from Hanging Lake reported here is based on AMS 14C dates of plant macrofossils or chironomid chitin. This new chronology has important implications for regional vegetation histories and climate reconstructions.

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Study site Hanging Lake (62°21′N, 138°21′W) is located in northern Yukon Territory, Canada, within a region of continuous permafrost in the foothills of the Barn Mountains (Fig. 1). Situated in a raised basin composed of shale, siltstone, and sandstone, the site is at an elevation of ∼ 500 m and lies ∼80 km from the Beaufort Sea. The modern vegetation around the lake consists of tussock tundra, wet sedge meadows, heath tundra, and patchy fellfield (Cwynar, 1982). Catchment area is negligible in size compared to the ∼ 60-ha surface area of the lake. There are no inflowing streams, however a small outlet stream is present at the northeast end. In May 2002, a maximum water depth of 9.5 m was measured at the southern end. A pH of 6.2 was also recorded. The modern northern Yukon Territory experiences a sub-Arctic continental climate with low annual precipitation and large ranges in temperature (Wahl et al., 1987). Located near the transition between the Porcupine-Peel Basin and Northern Mountains climatic divisions, Hanging Lake experiences cold and prolonged winters (October to late May) and brief summers with mild, variable weather due to the proximity to the Beaufort Sea (Wahl et al., 1987). The mean annual temperature at Hanging Lake is −11.6°C and the mean air temperature of July is estimated at 10°C (New et al., 2002).

Figure 1. Map of Hanging Lake including elevation contours (m) and hydrologic features. Inset shows regional map with the area surrounded by the black square depicting the extent of Laurentide ice in study region at ∼18 (dotted line), 15, and 13 14C ka BP (Dyke et al., 2003). Stippled areas on the map to the right show the extent of glacial lakes.

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Methods Sediment sampling and laboratory analysis In May 2002, a 3.69-m core was recovered from the ice surface at the southern end of Hanging Lake, close to where Cwynar (1982) cored. A Livingstone piston sampler and a crank drill driver were used to core the silt-rich sediments. Although the final drive penetrated to 4.02-m sediment depth, the basal sediments were not recovered and the resulting core ended at 3.69 m. Cores were measured, wrapped in plastic and aluminum foil in the field, and transported to the laboratory where they were stored at 4°C. Cores were split longitudinally; one half was sub-sampled and the other was archived. Working cores were mostly sub-sampled at 0.5-cm intervals. Loss-on-ignition (LOI) was used to estimate the organic content of sediment (Dean, 1974; Heiri et al., 2001). Sediment samples of 0.5 mL were dried at 105°C overnight and combusted at 550°C for exactly 4 h, with the weight loss measured after each stage once the samples cooled to room temperature in a dessicator. Chronology and pollen analysis of the basal sediments A total of eight AMS 14C dates (Table 1) from unidentified woody macrofossil fragments, moss fragments, or chironomid chitin were used to develop an age–depth model (Fig. 2A). For each chitin sample, 100 mL of sediment was warmed in distilled H2O and washed through a series of nested sieves (500-, 250-, and 106-μm meshes). Only the largest head capsules that did not contain sediment were hand-picked from a Petri plate. Approximately 800 whole head capsule equivalents were placed onto a glass depression slide, dried, weighed, and scraped into a Wheaton-33® 0.3 mL V-Vial. All ages were calibrated to calendar years using CALIB v. 5.0.2 (Stuiver and Reimer, 1993; Reimer et al., 2004). The midpoint of the cumulative probability of the calibration curve was used and calibrated ages were rounded to the nearest decade (Table 1). The age–depth model was developed using a second-order polynomial (Fig. 2A). To compare the Hanging Lake basal core sequence collected in 2002 with the original core of Cwynar (1982), pollen analysis was undertaken at ∼ 10-cm intervals from 364–200 cm. Sediment preparation followed standard methodology (Faegri et al., 1989), with fine-grained sediments removed by sieving through a 7-μm mesh screen (Cwynar et al., 1979). Initial washes in sodium pyrophosphate were omitted. Pollen identification was aided by the use of keys (McAndrews et al., 1978; Faegri et al., 1989) and a modern reference collection at the University of New Brunswick. At least 300 terrestrial pollen grains (min = 305, max = 375, mean = 331) were counted. Pollen percentages were based on the sum of all tree, shrub, and herb pollen and total pollen accumulation rates (grains cm− 2 yr− 1) were based on the sum of the concentration of all identifiable pollen divided by sediment deposition times (Fig. 3).

Figure 2. (A) Age–depth model with ages expressed as AMS 14C dates (closed circles) and their respective calibrated ages (open circles). (B) Comparison of uncalibrated age– depth relationships between eight AMS 14C dates from the record presented here (closed circles) and twenty-one conventional 14C dates (open squares) from Cwynar (1982).

Chironomid sampling and identification Fossil midges were processed following methods given in Walker (2001). Sediment samples (ranging from 0.5–13 mL) were heated in 5% KOH for ∼20 min, then poured onto a 95-μm mesh sieve and rinsed with distilled water, and the residue was then transferred with distilled water to a beaker. Aliquots were poured into a Bogorov counting tray. Using a dissecting microscope, remains were handpicked with fine forceps for each entire sample. Remains were then placed onto a cover glass, separated, and permanently mounted on a slide using Entellan®. Chironomidae head capsules and Chaoboridae mandibles were identified with a compound microscope under bright-field illumination at 400× magnification. Chironomidae were identified with reference to Wiederholm (1983), Oliver and Roussel (1983), and Walker (2007). For consistency between fossil taxa and the modern

Table 1 AMS 14C dates from the 2002 Hanging Lake record. Depth (cm)

Material dated

Lab number

14

14–15 91–92 115–120 134–139 251–255 272–277 272–277 295–300

Unidentified twig Twig fragments Moss fragments Moss fragments Moss fragments Chironomid chitin Moss fragments Chironomid chitin

UCIAMS UCIAMS UCIAMS UCIAMS UCIAMS UCIAMS UCIAMS UCIAMS

1195 ± 15 5960 ± 15 7295 ± 25 8035 ± 30 12,020 ± 60 12,745 ± 45 13,250 ± 70 13,030 ± 60

19030 19031 21611 21610 20985 20987 20984 19032

C yr BP

Calibrated yr BP (2σ range) 1120 (1067–1172) 6780 (6738–6848) 8105 (8029–8171) 8915 (8777–9017) 13,880 (13,754–14,015) 15,060 (14,813–15,281) 15,700 (15,335–16,107) 15,380 (15,104–15,738)

J. Kurek et al. / Quaternary Research 72 (2009) 246–257 Figure 3. Hanging Lake pollen percentage diagrams (selected taxa) of basal sediments from the new record (white silhouette) and Cwynar (1982) record (dark silhouette). Pollen percentages are plotted on a common depth scale from 364–200 cm (∼17 to 12 ka). Dark silhouetted taxa from Cwynar (1982) correspond to the same fossil taxa as those immediately to the left. Chironomid-inferred temperatures are also plotted with sample-specific prediction errors. Zones are based on the chironomid assemblages and the pollen zones of Cwynar (1982) are also shown.

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250 J. Kurek et al. / Quaternary Research 72 (2009) 246–257 Figure 4. Chironomid stratigraphy from Hanging Lake. Taxa are ordered based on their weighted average of location (time) in the record. Temperature optima are in parentheses and taxa without optima are absent from the training set or lumped at higher taxonomic groupings.

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training set, Appendix 1 in Barley (2004) was used for taxonomic groupings distinct to the transfer function applied. A minimum count of 50 head capsules was targeted for reliable inferences (Heiri and Lotter, 2001; Larocque, 2001; Quinlan and Smol, 2001). A total of 114 out of 122 samples met this criterion (head capsule count of all samples: min = 9.5, max = 436, mean = 102). Eight samples with b50 head capsules are still included in the reconstruction. Rarefaction analysis was performed using PRIMER v. 5 (Clarke and Gorley, 2001) to approximate fossil taxon richness based on a constant sample size. Chironomid zones Zones were identified (Fig. 4) using the optimal-splitting-byinformation-content technique in PSIMPOLL v. 4.25 (Bennett, 2005). Proportional data were square-root transformed, taxa never reaching 3% abundance in any one sample were excluded, and proportions were then recalculated prior to zoning. Significant zones were identified using variance reduction as a percent of the total variance with comparison to a broken-stick model (Jackson, 1993; Bennett, 1996). Inference model and analog matching All identifiable midges were included as a percentage of the total identifiable count. Mean July air temperatures were estimated using a model developed by Barley et al. (2006). The weighted averaging-partial least squares (WA-PLS) 1-component model was selected as the most parsimonious inference model (r2boot = 0.78, RMSEPboot = 1.58°C) from the 136 lake training set. Inferences are likely more reliable if the fossil assemblages are similar to those in the modern training set (Birks et al., 1990; Jackson and Williams, 2004). To identify analogs within the modern training set for the Hanging Lake fossil samples, we used squared chord distance (SCD) as an assemblage dissimilarity measure (Overpeck et al., 1985; Laird et al., 1998; Laing et al., 1999). SCD is a multivariate dissimilarity measure that reduces the influence of dominant taxa while enhancing the importance of minor taxa (Overpeck et al., 1985; Wahl, 2004). Routinely used for comparisons of modern and fossil

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pollen assemblages, SCD is accepted as an appropriate dissimilarity measure for fossil data (Overpeck et al., 1985; Gavin et al., 2003). Rare taxa in the training set (present in fewer than 6 out of 136 lakes) and fossil taxa never reaching 3% abundance in any one sample were removed prior to analog matching. Data were transformed into proportions for analysis using modern analog technique (MAT) in C2 (Juggins, 2003). Analogs were established by determining the mean minimum dissimilarity coefficient (min DC) within the modern training set and selecting an upper confidence interval (C.I.) of 95% as the cut-off. Next, we compared the fossil samples to the modern training set and recorded the min DC for each fossil sample. Assemblages greater than the upper 95% C.I. of the mean min DC of the training set are considered poor analogs, whereas assemblages less than the upper 95% C.I. are considered analogs (Fig. 5A). Another method to evaluate the similarity of fossil samples to the training set is to examine the squared residual distance to Axis 1 of a canonical correspondence analysis (CCA) constrained to one explanatory variable (Birks et al., 1990). We used a CCA of the training set constrained only to temperature, with the fossil samples plotted passively, and calculated the squared residual distance to Axis 1 of the fossil samples (Fig. 5B). CCA was performed using CANOCO v. 4.5 (ter Braak and Šmilauer, 2002). Percentage data were square-root transformed and rare species were downweighted. Fossil taxa not present in the training set and those defined as rare in the training set were omitted. The squared residual distance to Axis 1 was log (x + 1) transformed due to right skewness. Fossil samples with a transformed squared residual distance to CCA Axis 1 equal to or greater than the upper 95% C.I. of the mean transformed squared residual distance of the training set are considered poor assemblages from which to infer the variable of interest (Birks et al., 1990; Laing et al., 1999). Results Chronology: new versus the original Hanging Lake record Our age–depth model indicates that the new Hanging Lake record extends to the waning stages of the last glacial period (Fig. 2A). Due to

Figure 5. Analog comparisons of fossil samples to modern assemblages of the Barley et al. (2006) training set. Zones are also plotted. (A) Analog comparison using squared chord distance as the dissimilarity measure. Fossil samples with minimum dissimilarity coefficients (min DC) greater than the analog cut-off are considered poor analogs. (B) Constrained ordination analysis of fossil samples (plotted passively) to temperature. Samples with log (x + 1) squared residual distance greater than the analog cut-off are considered poor inference analogs.

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the absence of dates from the basal sediments, our age–depth model extrapolates ages for the basal ∼64 cm of sediment. Cwynar (1982) derived age estimates for the original 403-cm Hanging Lake record from twenty-one conventional 14C dates on bulk sediment and proposed that the record at 370 cm is “at least 25 000 and possibly 33 000 yr old” (Fig. 2B). The new record we present here is 34 cm shorter than the original record and our age–depth model, based on calibrated AMS 14C ages, estimates an age of ∼17 ka for the basal-most sediments. The discrepancy between the basal ages of Hanging Lake sediments presented here and basal ages from Cwynar (1982) are likely the result of bulk sediment ages much too old and the subsequent age–depth model (Fig. 2B) chosen by Cwynar (1982). Chronologies developed from radiocarbon dating of bulk sediments in Arctic lakes are frequently inaccurate because of low rates of productivity and decomposition of organic matter (see Abbott and Stafford, 1996). Also, although limestone is absent from the surrounding bedrock, the presence of pre-Quaternary spores in the basal sediments indicates that ancient carbon was eroded from the basin. Our record of vegetation change is roughly similar to Cwynar (1982) once a ∼ 20-cm offset is taken into account (Fig. 3). The offset is most obvious for the Betula rise, Salix peak, and Cyperaceae peak, which in the original record are down-core of these events presented here. While we have no clear indication why this offset between pollen percentages exists, LOI (%) from both records increases from

constant values of 5% at approximately 300 cm (∼15.5 ka [Fig. 6]). Therefore, we assume that the basal sediments generally capture similar time periods. Pollen percentage and total accumulation rates ∼17 to 12 cal ka BP Prior to ∼15 ka, pollen assemblages in the new record are dominated by Poaceae (∼30%), Artemisia (∼10–20%), and Cyperaceae (∼5–15%), as well as comparatively high percentages of Brassicaceae (∼5–10%) between ∼ 17 to 16 ka (Fig. 3). The Salix peak occurs abruptly at ∼15 ka and the Betula rise at ∼14 ka. Total pollen accumulation rates (PAR) range between ∼180 and 1200 grains cm− 2 yr− 1 and show a general trend of increasing values after 16 ka. Chironomid zones Seven chironomid zones are identified (Fig. 4). Generally, zones from the basal sediments are brief in duration compared to those after 15 ka. Cricotopus/Orthocladius, Heterotrissocladius, Chironomus, Procladius, and various genera of Tanytarsini are common and abundant throughout the record. Zone HL-1 (364–324 cm core depth; ∼17–16.3 ka) is dominated by Hydrobaenus/Oliveridia. Other taxa at abundances near 20% include Micropsectra-type, Mesocricotopus, Constempellina, and Tanytarsini (und.). Parakiefferiella nigra-type is only present within HL-1.

Figure 6. Comparisons of the δ18O record from GISP2 (Grootes and Stuiver, 1997) and LOI (%), Influx of head capsules (INFLUX (hc)), and mean July air temperatures (mean July temp. (°C)), including sample-specific prediction errors and the modern mean July air temperature at Hanging Lake (vertical line). A LOESS smooth (span = 0.1) captures general trends in the LOI (%) and INFLUX (hc). Zones are based on the chironomid assemblages.

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Endochironomus, Glyptotendipes and Polypedilum (not shown). Chaoboridae are frequent and reach 10% within HL-4. Cricotopus/Orthocladius and Tanytarsini (und.), with abundances N40%, are the most abundant taxa in zone HL-5 (190–92.5 cm core depth; ∼11.7–6.5 ka). Early in the zone, Corynocera oliveri-type is frequent. Common taxa with moderate abundances (∼20%) within zone HL-5 also include: Micropsectra-type, Sergentia, Tanytarsus-type, Chironomus, and Heterotrissocladius. Only Heterotrissocladius exceeds abundances of 40% in zone HL-6 (92.5–54 cm core depth; ∼6.5–4 ka). There are six members of Tanytarsini that are common with moderate abundances (10–20%) in HL-6, including Paratanytarsus-type and Stempellinella/Zavrelia. Other common taxa with relative abundances near 10% include Sergentia, Parakiefferiella triquetra-type, and both Zalutschia-type A and B. Zone HL-7 (54–0 cm core depth; ∼4 ka to the present) is dominated by Heterotrissocladius, Micropsectra-type, Tanytarsus-type, Chironomus, and Stictochironomus. Both Zalutschia-type A and B increase in abundance compared to the preceding zone. Corynocera ambigua-type peaks at ∼5% during the most recent 1000 yr. Temperature inferences

Figure 7. Inferred temperatures and sample-specific prediction errors (1.56–1.89°C) from Hanging Lake at (A) 14 ka to the present, with a LOESS smooth (span = 0.1) and (B) 17 to 14 ka. Note the different scalings of the ordinate axes. Eight samples marked with an asterisk have b 50 whole mentum equivalents. Chironomid zone boundaries are denoted by dashed vertical lines.

Stictochironomus and Chironomus are the only Chironomini present in this zone, albeit at low abundances. One Chaoborus mandible is present in the basal-most sample. Hydrobaenus/Oliveridia dominates zone HL-2 (324–295 cm core depth; ∼ 16.3–15.5 ka). Other taxa present at ∼20% abundance include Cricotopus/Orthocladius, Tanytarsini (und.), Micropsectra-type, and Heterotrissocladius. Minor contributors to the samples include Pseudodiamesa, Mesocricotopus, and Protanypus. Chironomus is the only Chironomini present in HL-2. Average rarified taxon richness within zone HL-3 (295–275 cm core depth; ∼15.5–15 ka) is the lowest of the entire record. The only taxa present are Pseudodiamesa, Micropsectra-type, and Chironomus. Pseudodiamesa is the only taxon recovered from 3 of the 4 samples within HL-3, whereas Micropsectra-type and Chironomus are each found at 10% abundance from one sample. Head capsule sums are also low at 11, 34, 38.5, and 75. Cricotopus/Orthocladius and Micropsectra-type reach their highest abundances (N40%) in the entire record within HL-4 (275–190 cm core depth; ∼ 15–11.7 ka). Other taxa that exceed 20% include Mesocricotopus and Tanytarsini (und.). Eleven taxa (not all shown) also register their first occurrence within this zone, including Corynocera oliveritype, Paracladius, and Monodiamesa. Chironomini diversity also increases with the first occurrence of Microtendipes and Sergentia. Other Chironomini found at low abundance include Cryptochironomus,

Four general features are evident from the temperature reconstruction at Hanging Lake (Figs. 6 and 7): 1) decreasing temperatures between ∼17 and15 ka, 2) extremely low temperatures in HL-3 between 15.5 and 15 ka, 3) rapid temperature increase at the HL-3 to HL-4 zone boundary at 15 ka, and 4) relative temperature stability throughout the last ∼ 15 ka, with the exception of a late-Holocene cooling event between ∼4.5 and 2 ka. Only zones HL-2 and HL-3 have temperatures that never exceed modern. Summary averages of inferred temperatures for each zone are presented in Table 2. LOI (%) and head capsule influx The organic content of sediment (LOI %) is extremely low, ranging from 0.2 to 16.6% (Fig. 6). LOI is ∼5% between 17 and 15.5 ka (HL-1 through HL-3) and demonstrates substantial variability between 15.5 and 14 ka (early in HL-4). The greatest average LOI (%) occurs in zones HL-4 and HL-5, between ∼14 and 8.5 ka. After 8.5 ka, LOI (%) gradually decreases to b10%. Head capsule concentration and influx demonstrate similar trends throughout the entire record, so only influx is presented (Fig. 6). Generally, the influx of head capsules is lowest early in the record within zones HL-1 to HL-3, between ∼17 and 15 ka. In zone HL-4 it reaches ∼60% of the maximum record value that occurs at ∼2.5 ka. After ∼12 ka, the influx of head capsules increases gradually to mid-Holocene values. Influx of head capsules is then generally highest and most variable from ∼6.5 ka to the present. Analog comparison and evaluation of inferences The analog comparison indicates poor analogs throughout most of the record (Fig. 5A). Overall, 86% of fossil samples min DC values exceed the analog cut-off value. The greatest poor analogs occur from

Table 2 Zone characteristics of the 2002 Hanging Lake record. Zones and ages (ka)

# of fossil samples

Zone averages of mean July air temperature (°C)

Deviation from the modern mean July air temperature (°C)

Taxonomic richness (rarefied)

HL-7 HL-6 HL-5 HL-4 HL-3 HL-2 HL-1

38 18 22 16 4 13 11

10.84 11.37 11.45 10.84 0.92 5.52 7.69

0.84 1.37 1.45 0.84 −9.08 −4.48 −2.31

13.0 14.2 11.6 10.5 1.5 6.7 7.7

(4 to present) (6.5 to 4) (11.7 to 6.5) (15 to 11.7) (15.5 to 15) (16.3 to 15.5) (17 to 16.3)

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∼ 17 to 14.4 ka. Of the 17 fossil samples (14% of samples) that are considered analogs, nine occur during two distinct periods: ∼ 13 to 12 ka (i.e., during the Younger Dryas stade) and the most recent ∼1 ka. Eight other analogs occur sporadically throughout the record. The log (x + 1) squared residuals in a CCA constrained to temperature indicate that 67% of fossil samples are suitable for inferring mean July air temperature (Fig. 5B). However, nearly all basal-most samples (∼17 to 14.3 ka) except two exceed the inference analog cut-off value. Clearly, fossil samples in HL-1 to early in HL-4 differ from those in the modern training set (Fig. 5). Also, between 4 and 3.5 ka, a sequence of six samples exceeds the inference analog cut-off value. Discussion Temperature inferences in relation to poor analogs Do the poor analogs limit the climate reconstruction at Hanging Lake? Lotter et al. (1999) found that, despite poor analogs, chironomid-based inference models are able to estimate reliably modern temperatures and may even be applied to late glacial noanalog situations. Our data suggest that early samples (∼17 to 14.4 ka) at Hanging Lake are largely different than modern and are considered poor analogs, especially during zone HL-3 where taxonomic richness is lowest (Fig. 4). Hill's N2 values (i.e., diversity measure that weights taxon abundance by occurrence, see Hill [1973]) differ between taxa common to both the fossil samples and the modern training set of Barley et al. (2006). In fossil samples, nine taxa have Hill's N2 values that are on average 23 units higher than their values in the training set. Six fossil taxa have Hill's N2 values that are on average 19 units lower than their values in the training set. Fifty percent of fossil taxa have Hill's N2 values that are comparable (±10 units) with their Hill's N2 value in the training set. Thus, most samples from Hanging Lake are poor analogs because they are composed of taxa with abundances outside of their range in the training set, not because the fossil taxa differ in type from those in the training set. However, considering that fewer poor analogs were identified with the inference analog test (Fig. 5B), most temperature inferences, except those in zones HL-1 to HL-3, are generally reasonable despite the dissimilarities between some fossil samples and the modern assemblages within the training set (Fig. 5A). For all fossil taxa, except Constempellina and some Tanytarsini, weighted averaging of the taxa within the training set provides estimates of each taxon's temperature optimum. It is then reasonable to assume that fossil samples composed of taxa with adequately estimated optima can be used to infer environmental variability irrespective of subtleties in fossil taxon abundances compared to their modern abundances. We emphasize that the patterns and direction of change, not the absolute values, are the most important aspect of any paleoenvironmental reconstruction. Paleotemperature patterns at Hanging Lake Last glaciation–Holocene transition (∼17 to 12 cal ka BP) Climatic conditions during the early history (HL-1 through HL-3) of Hanging Lake may have been influenced by the proximity (Fig. 1) of the Laurentide Ice Sheet (LIS) to the north and east (Rampton, 1982 and 1988; Dyke and Prest,1987; Dyke et al., 2003). Given our chronology and a minimum age of ∼13 14C ka BP for Hanging Lake sediments (Fig. 2A), it is likely that the onset of sedimentation began around the culmination of the drainage of glacial Lake Old Crow, between ∼15 and 13 4C ka BP (Zazula et al., 2004). At that time, Hanging Lake was approximately 60 to 100 km from the LIS (Fig. 1). At times, cold-adapted taxa (Bennike et al., 2004) dominate, including Hydrobaenus/Oliveridia and Pseudodiamesa (Fig. 4). On average, temperatures are ∼5°C below modern and abrupt environmental changes are reflected by brief chironomid assemblage zones (Table 2). Head capsule influx also indicates low chironomid productivity from ∼17 to 15 ka (Fig. 6). Influx values are highly

dependent on the sedimentation rate and should be viewed with caution given our lack of dates constraining the age–depth model for sediments prior to 15 ka. There is also evidence of a brief period of relative warmth from ∼16.7–16.4 ka (Fig. 7), with temperatures near modern at about 10°C. The regional significance of this warm period remains unclear, but may be related to regional LIS dynamics which characterized its northwestern-most margins during the late Wisconsinan (Dyke et al., 2003; Murton et al., 2007). The coldest temperatures at Hanging Lake occur between 15.5 and 15 ka (zone HL-3) and approximately correlate with the Sitidgi stade, when Laurentide ice re-advanced into the Mackenzie Delta at about ∼13 14C ka BP (Rampton, 1988; Murton et al., 2007), and the Tutsieta Lake Phase, when Laurentide ice re-advanced from the southern Mackenzie Mountains to the northern Richardson Mountains (DukRodkin et al., 1996; Zazula et al., 2004). Temperatures are 2°C below zero; however, these values result from 100% Pseudodiamesa in three of four samples, which are the poorest analogs in the entire record (Fig. 5B). Likely, temperatures were closer to the optimum of Pseudodiamesa estimated, to be ∼5°C (Barley et al., 2006). Diamesinae are cold-stenotherms and may prefer flowing water (Oliver and Roussel, 1983). Therefore, we presume that Hanging Lake received seasonal meltwater from ice or snow pack near the site. Lake levels may also have been low, as there were numerous moss fragments (mainly Warnstorfia fluitans (Hedwig) Loeske) within zone HL-3 sediments. Between 15 and 14 ka, Hanging Lake was ∼ 110 km west of the LIS (Fig. 1). At ∼ 15 ka, concurrent with abrupt Bølling warming in the GISP2 ice core record, temperatures are at least as warm as modern. Chironomid richness (rarefied) also increases by an order of magnitude from ∼ 15 to 12 ka, compared to the previous three zones, and average zone temperature increases by ∼5.0°C (Table 2). After 15 ka, temperatures are relatively stable, showing reduced variability compared to the previous few millennia. There is no evidence of the Younger Dryas stade (YD: ∼ 13 to 11.6 ka in GISP2) at Hanging Lake. The YD is well-recorded in lake sediments from mainly southern Alaska (Engstrom et al., 1990; Bigelow and Edwards, 2001; Mann et al., 2002; Kaufman et al., 2003; Hu et al., 2002, 2006) and evidence indicates its prominence at sites closer to the North Pacific Ocean (Mathewes, 1993; Mikolajewicz et al., 1997; Bigelow and Edwards, 2001; Kokorowski et al., 2008). Modeling exercises suggest teleconnections between the North Atlantic and North Pacific oceans (Mikolajewicz et al., 1997; Resson, 1997) in addition to observational data supporting such coupling (Kienast and McKay, 2001), but evidence for antiphase trends between those regions during deglaciation is also supported (Schmittner et al., 2003; Sarnthein et al., 2006). However, with increasing summer insolation, changing effective moisture (Abbott et al., 2000; Edwards et al., 2001), and large exposed areas of the Bering Land Bridge, cooling in the North Pacific may only have subtly affected regions distant to its coast. In a robust spatial analysis of the affects of the YD across Beringia, Kokorowski et al. (2008) note the event's heterogeneity and in general find little evidence of the YD at sites north of the Alaska Range. They propose that the presence of a strong high pressure system north of Alaska and greater summer insolation may have caused summer climates in much of central and northern Alaska to be warmer than at present. Late Pleistocene–early Holocene thermal maximum Quantitative records of the late Pleistocene–early Holocene thermal maximum (HTM) in eastern Beringia are extremely sparse from ∼135° to 150°W longitudes and regions show inconsistent evidence (Kaufman et al., 2004). Only a subtle HTM is evident, as there are a few warmer than average temperature inferences at ∼12 ka (Fig. 7A), likely resulting from minor increases in Microtendipes, Chironomus, and Chaoborus (Fig. 4). Zoning methods also identify a shift in chironomid assemblages at ∼11.7 ka (Table 2

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and Fig. 4); however, the temperature changes during this period are minor and well within the prediction error of the transfer function applied. Our findings are consistent with regional evidence supporting an early HTM including: 1) pollen percentage increases of Populus balsamifera, Typha latifolia, and Myrica gale (Cwynar, 1982), 2) increased formation of thermokarst lakes in northwestern Canada (Rampton, 1988), 3) regional treeline advance (Ritchie et al., 1983; Cwynar and Spear, 1991; Spear, 1993), and 4) high concentrations of beta-carotene in lake sediments suggesting that total algal abundance was greatest between ∼12 and 9 ka (Pienitz et al., 2000). In general, the late Pleistocene–early Holocene climate at Hanging Lake can best be characterized as near modern (∼12–10°C), with a greater range in temperatures than other periods during the last ∼ 15 ka. Holocene (11.6 cal ka BP to the present) The magnitude of Holocene temperature change is much less than during the last glaciation–Holocene transition at Hanging Lake (zones HL-1 to HL-4). All Holocene temperature changes are within the overall model prediction error (±1.58°C) and suggest that other environmental factors, such as lake levels, likely assume a larger role in driving chironomid assemblage change during the Holocene. For example, effective moisture increases throughout much of eastern Beringia during the mid-Holocene (Hu et al., 1998; Abbott et al., 2000; Pienitz et al., 2000; Anderson et al., 2005a). Alnus, an indicator of greater effective moisture, also becomes a dominant pollen contributor in most regional vegetation histories between ∼ 7 and 5 ka (Anderson and Brubaker, 1994; Cwynar and Spear, 1995; Szeicz and MacDonald, 2001), as well as from the original Hanging Lake pollen record (Cwynar, 1982). Increased abundances of profundal taxa, namely, Heterotrissocladius, Parakiefferiella triquetra-type, Zalutschiatype B, and after ∼ 2 ka, Micropsectra-type occur during the mid-tolate Holocene at Hanging Lake (Fig. 4). These taxa also tend to be associated with colder or moderate air temperatures in the Barley et al. (2006) model, further complicating our reconstruction. Nonetheless, the most pronounced Holocene climate oscillation at Hanging Lake is a ∼1.5°C cooling event from ∼4.5 to 2 ka (Fig. 7), likely reflecting neoglacial climatic conditions (Denton and Karlén, 1973; Davis, 1988). This event is the most significant of our Holocene record, although there is no evidence of it in the original Hanging Lake pollen record (Cwynar, 1982). This late Holocene cooling event contrasts with evidence of treeline advance and/or increased density of Picea populations at 3 ka from high-elevation sites in the Mackenzie Mountains, south of our study area (Szeicz and MacDonald, 2001). If effective moisture increased substantially during the mid- to late Holocene at Hanging Lake, inferences of a cooling event from 4.5 to 2 ka may be confounded by rising lake levels as warm-adapted littoral taxa become less abundant due to increased abundances of profundal types such as Heterotrissocladius and Zalutschia-type B (Fig. 4). Regionally, increased effective moisture throughout the late Holocene is evidenced by core transects suggesting lake level overflow at Birch Lake, interior Alaska (Abbott et al., 2000) and from ∼ 4 to 2 ka at Marcella Lake, southwestern Yukon (Anderson et al., 2005a). A proposed mechanism may be shifts in regional atmospheric circulation via the position and intensity of the Aleutian Low (Spooner et al., 2003; Anderson et al., 2005b); which is identified as the climatic feature responsible for moisture delivery from the Gulf of Alaska to interior Alaska and Yukon. Vegetation responses in relation to temperatures (∼17 to 12 cal ka BP) and re-interpretation of the original Hanging Lake record The willow (Salix) peak and birch (Betula) rise between ∼ 16 and 12 ka are characteristic features of vegetation histories from eastern

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Beringia and mark a change from herb-dominated to shrubdominated communities. This shift to shrub-dominated communities is generally interpreted as a response to increases in temperature and effective moisture (Ager and Brubaker, 1985; Anderson and Brubaker, 1994; Bigelow and Edwards, 2001; Brubaker et al., 2005) or snow depth (Fredskild, 1991). At Hanging Lake, the increase in Salix coincides with a rapid rise in temperature to nearmodern values at ∼ 15 ka (Fig. 3). The birch rise, however, is not associated with a change in temperature, even though modern response surface curves for Betula pollen percentages indicate that northern birch species are strongly controlled by summer temperature (Anderson et al., 1991). The lag of ∼ 1000 yr in the response of shrub birch to the temperature increase at ∼ 15 ka suggests that some other variable limited its expansion; possibilities include effective moisture, poorer dispersal than willow or competition. Hu et al. (2002) also noted that at Nimgun Lake, southwestern Alaska, the establishment of Betula shrub tundra lagged warming by 1000 yr. Given the age differences (Fig. 2B) between the basal sediments of the record presented here and Cwynar (1982), a re-examination of the general timing of major vegetation shifts from the original record is warranted. A comparison of the chronologies (Fig. 2B) indicates that the radiocarbon ages from the original record are substantially too old. For example, the beginning of the Salix rise is at ∼18.5 14C ka BP in the original record (Cwynar 1982), but at ∼ 12.9 14C ka BP in the new record. In addition, there is a ∼ 20-cm offset for the beginning of the Salix rise between the records. Taking into account the younger ages of the new AMS-dated record, the offset, and the similarity of the overall pollen stratigraphy of the two records, the original record may postdate the last glaciation, given that the extrapolated age–depth model places the base of the new record at ∼ 17 ka. However, the absence of dates from the bottom-most 69 cm of the new record (because macrofossils were not found) precludes a definitive assessment of the age of the record. Thus, this new chronology throws into question whether or not the Hanging Lake record extends to the last glaciation as previously reported (Cwynar and Ritchie, 1980; Cwynar, 1982). Comparisons with the GISP2 record during the last glaciation–Holocene transition A comparison of temperature trends during the last glaciation– Holocene transition at Hanging Lake with climatic oscillations recorded in the GISP2 ice core record from Greenland (Grootes and Stuiver, 1997) shows many differences (Fig. 6). For example, between ∼ 17 ka and the start of Bølling interstadial at 14.7 ka, conditions in the North Atlantic region are extremely cold, whereas temperatures at Hanging Lake are already near modern at ∼ 16.5 ka and then show a progressive cooling trend of at least ∼ 5°C (Fig. 7B). In addition, throughout the Bølling–Allerød interstadial and Younger Dryas stade, between ∼ 14.7 and 11.6 ka, temperatures at Hanging Lake are near or slightly warmer than modern and stable. This contrasts markedly with the overall climatic cooling recorded in the GISP2 δ18O record. The only similarity between the records is the onset of abrupt warming at 15 ka from Hanging Lake, nearly coeval with the start of the Bølling interstadial at 14.7 ka. Substantial differences between climatic inferences from Hanging Lake and the GISP2 δ18O record cannot be resolved by single-site comparisons alone and highlight the importance of regional influences (e.g. proximity to the LIS) to our paleoenvironmental record. If indeed our climatic inferences are correct, and near-modern temperatures were attained by 15 ka in eastern Beringia, then shifts in other key environmental variables, such as effective moisture, must be further examined within a multi-proxy approach in order to elucidate causal mechanisms of paleoenvironmental change during the late Quaternary in eastern Beringia.

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Conclusions A new, AMS-based chronology from Hanging Lake demonstrates that the lake sediment record extends from ∼ 17 ka to the present and suggests that the original record and subsequent vegetation history are likely too old. Inferred temperatures during the last glaciation– Holocene transition do not exhibit trends similar to those of the North Atlantic region as recorded by the GISP2 δ18O record. Prior to 15 ka, Hanging Lake experienced three relatively brief periods of considerably colder than modern climatic conditions, with the most severe between ∼ 15.5 and 15 ka. A brief warm interval centered at ∼ 16.5 ka indicates that mean July air temperatures in the region were at least as warm as modern. Significant warming occurs at ∼ 15 ka, coeval with an increase in Salix pollen, whereas warm and stable temperatures near-modern values precede the birch rise by ∼ 1000 yr. Between ∼17 and 12 ka, total pollen accumulation rates never exceed ∼ 1200 grains cm− 2 yr− 1. There is no clear indication of the Younger Dryas stade and the early Holocene thermal maximum is best recorded qualitatively (i.e., assemblage shifts and subtle increases in warm-adapted midges), rather than by significantly warmer temperatures than modern. Holocene temperatures range from ∼12 to 10°C and are relatively stable compared to the often abrupt, highmagnitude changes recorded between ∼17 and 15 ka. The most significant cooling event of the Holocene, a decrease and subsequent increase in temperature of 1.5°C, occurs from ∼4.5 to 2 ka, reflecting neo-glacial climatic conditions in the region. Acknowledgments This research was supported by the Natural Sciences and Engineering Research Council of Canada (NSERC), Collaborative Research Opportunity on “Late Pleistocene paleoclimates of eastern Beringia”. Additional scholarly support was provided by NSERC with a doctoral scholarship to J Kurek. MB Abbott, E Barley, J Racca, and IR Walker assisted in the field. Ray Spear provided constructive comments on an early draft of the manuscript. Thanks to Bruce Bagnell for identifications of the moss fragments. We thank two anonymous reviewers for their insightful comments and suggestions which greatly improved this manuscript. References Abbott, M.B., Stafford, T.W., 1996. Radiocarbon geochemistry of modern and ancient Arctic lake systems, Baffin Island, Canada. Quaternary Research 45, 300–311. Abbott, M.B., Finney, B.P., Edwards, M., Kelts, K.R., 2000. Paleohydrology of Birch Lake, central Alaska: lake-level reconstructions using seismic reflection profiles and core transect approaches. Quaternary Research 23, 154–166. Ager, T.A., Brubaker, L.B., 1985. Quaternary palynology and vegetation history of Alaska. In: Bryant, V.M., Holloway, R.G. (Eds.), Pollen Records of Late Quaternary North American Sediments. InAmerican Association of Stratigraphic Palynologists, Dallas, Texas, pp. 353–384. Anderson, P.M., Brubaker, L.B., 1994. Vegetation history of northcentral Alaska — a mapped summary of late-Quaternary pollen data. Quaternary Science Review 13, 71–92. Anderson, P.M., Bartlein, P.J., Brubaker, L.B., Gajewski, K., Ritchie, J.C., 1991. Vegetation– pollen–climate relationships for the Arcto–Boreal region of North America and Greenland. Journal of Biogeography 18, 565–582. Anderson, L., Abbott, M.B., Finney, B.P., 2005a. Large and rapid Holocene moisture balance shifts in the Yukon Territory, Canada, based on lake-level reconstructions. The Holocene 15, 1172–1183. Anderson, L., Abbott, M.B., Finney, B.P., Burns, S., 2005b. Regional atmospheric circulation change in the North Pacific during the Holocene inferred from lacustrine carbonate oxygen isotopes, Yukon Territory, Canada. Quaternary Research 64, 21–35. Barley, E.M., 2004. Palaeoclimate analysis of southwestern Yukon Territory using subfossil chironomid remains from Antifreeze Pond. MSc. thesis, Simon Fraser University. Barley, E.M., Walker, I.R., Kurek, J., Cwynar, L.C., Mathewes, R.W., Gajewski, K., Finney, B., 2006. A northwest North America training set: distribution of freshwater midges in relation to air temperature and lake depth. Journal of Paleolimnology 36, 295–314. Barnosky, C.W., Anderson, P.M., Bartlein, P.J., 1987. The northwestern US during deglaciation: vegetational history and paleoclimatic implications. In: Ruddiman, W.F., WrightJr. Jr., H.E. (Eds.), North America and Adjacent Oceans during The Last

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