LATE PLEISTOCENE AND HOLOCENE PALEOENVIRONMENTS OF … · 2015-09-04 · (Meltzer, 1993). Biotic interchange between continents is one reason why the paleoclimates and paleogeography

Pergamon Quaternary Science Reviews, Vol. 14, pp. 449-47 1, 1995.

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LATE PLEISTOCENE AND HOLOCENE PALEOENVIRONMENTS OF THE NORTH PACIFIC COAST

DANIEL H. MANN* and THOMAS D. HAMILTON? *Alaska Quaternary Center; University of Alaska Museum, 907 Yukon Drive, Fairbanks, AK 99775, U.S.A.

I-US. Geological Survey, 4200 University Drive, Anchorage, AK 99508, U.S.A.

Abstract - Unlike the North Atlantic, the North Pacific Ocean probably remained free of sea ice during the last glacial maximum (LGM), 22,000 to 17,000 BP. Following a eustatic low in sea level of ca. -120 m at 19,000 BP, a marine transgression had flooded the Bering and Chukchi shelves by 10,000 BP. Post-glacial sea-level history varied widely in other parts of the North

QSR Pacific coastline according to the magnitude and timing of local tectonism and glacio-isostatic rebound. Glaciers covered much of the continental shelf between the Alaska Peninsula and British Columbia during the LGM. Maximum glacier extent during the LGM was out of phase between southern Alaska and southern British Columbia with northern glaciers reaching their outer limits earlier, between 23,000 and 16,000 BP, compared to 15,00&14,000 BP in the south. Glacier retreat was also time-transgressive, with glaciers retreating from the continental shelf of southern Alaska before 16,000 BP but not until 14,000-13,000 BP in southwestern British Columbia. Major climatic transitions occurred in the North Pacific at 24,000-22,000, 15,000-13,000 and 11 ,OOO-9000 BP. Rapid climate changes occurred within these intervals, including a possible Younger Dryas episode. An interval of climate warmer and drier than today occurred in the early Holocene. Cooler and wetter conditions accompanied widespread Neoglaciation, beginning in some mountain ranges as early as the middle Holocene, but reaching full development after 3000 BP.

INTRODUCTION

The North Pacific Ocean borders the Asian and North American continents along an intricate coastline arcing between Japan and California (Fig. 1). Unlike the Atlantic Ocean, the North Pacific is virtually barred from the Arctic Ocean, being connected only through the shallow and narrow Bering Strait. Across this strait and surrounding parts of the Bering Platform, the fauna and flora of Asia and North America have mingled. By this route, humans probably first entered the New World (Meltzer, 1993). Biotic interchange between continents is one reason why the paleoclimates and paleogeography of the North Pacific coastline are of special interest. Another reason is that, like the North Atlantic, the North Pacific greatly affects the climate and weather over the surrounding landmasses, especially downwind in North America (e.g. Charles et al., 1994). Past and future climatic changes are intimately tied to oceanographic and climatic processes in the Pacific Ocean.

Although the largest of the three major oceans, the Pacific is poorly understood in terms of its late Pleistocene history. This lack of information is especially notable in the Russian sectors of the coast. Paleoceanographic research has lagged in the North Pacific because the carbonate-compensation depth, the water depth where calcareous microfossils dissolve, is above the bottom over large reaches of this deep ocean (Archer and Maier-Reimer, 1994). In other oceans, most notably in the North Atlantic, our reconstructions of paleoceanography and paleotemperature are largely based on

calcareous microfossils such as foraminifera contained in deep sea cores.

We review here the late Pleistocene (25,000-10,000 BP) and Holocene (10,000-O BP) paleoclimate and paleogeography of the coastal regions bordering the North Pacific from northern Japan, through Alaska, and down the northwest coast of North America through British Columbia into Washington state. Emphasis is placed on coastal areas of Alaska, the areas we know best. All ages are expressed as uncalibrated radiocarbon years before present.

The reader is cautioned that the wide geographical separation of data sites introduces assumptions about spa- tial scales into the paleogeographic reconstructions we present. As the distance between data sites increases, so does the uncertainty in the reconstructed image of the paleolandscape. The patch size relevant to an organism, whether human or spruce tree, can be far smaller than the resolution of our reconstructions. Scale problems also derive from the available temporal control since most studies rely on radiocarbon chronologies whose resolu- tions range from hundreds to thousands of years. Rates of climatic or geographic changes, which may have been critical in determining species survival or extinction, generally are inadequately known at present.

THE NORTHWEST PACIFIC OCEAN

The northwest Pacific Ocean stretches from Hokkaido to the Commander Islands and includes the Sea of Okhotsk and the Sea of Japan (Fig. 1). Today this sector

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450 Quaternary Science Reviews: Volume 14

FIG. 1. North Pacific region, showing landmasses, continental shelves (light grey), and major ocean currents (arrows). The Tsushima Current flows northward through the Sea of Japan and into the Pacific Ocean through Tsugaru Strait south

of Hokkaido.

of the Pacific, lying beneath the shifting border between cool Siberian air masses and warmer maritime air, is the birthplace of many of the cyclonic storms that travel to the North American coastline along the prevailing westerlies (Wendland and Bryson, 1981; Terada and Hanzawa, 1984). Hence the paleoclimatology of this region is of key interest in understanding the climatic history of other sectors of the North Pacific coastline.

Paleoclimatic events before 15,000 BP are poorly understood in the northwest Pacific. During the last glacial maximum (LGM), the Polar Front was displaced perhaps 5” of latitude further south (CLIMAP, 198 1; Morley and Heusser, 1989). Steepening of temperature gradients along the coast of northeastern Asia probably caused an intensification of wind speeds at the 500 mb level and a more zonal orientation of the subpolar jet stream (Kutzbach et al., 1993). The Japanese archipelago was much cooler and drier than today as maritime air masses were displaced south and eastwards and replaced by the Siberian High for much of the year (Morley and Heusser, 1989; Winkler and Wang, 1993). With lowered sea level, Japan was connected to the Asian mainland. Seasonal temperature contrasts were greater than today, probably because low-salinity surface waters stabilized the water column, preventing deep convection in

summer, and allowing surface waters to warm in summer (L.E. Heusser and Morley, 1985). Boreal-forest vegetation dominated by Picea, Abies, Tsuga and Pinus spread downslope and southward to cover most of the Japanese archipelago (Tsukada, 1983, 1985). Mean annual temperatures in the northern archipelago may have been as much as 8-9OC below present values and mean annual precipitation may have been one-third lower during the LGM (Tsukada, 1986; Kerschner, 1987; Morley and Heusser, 1989). Temperatures and precipitations predict- ed by global circulation modeling for the LGM in this region have similar signs and magnitudes as estimates based on the field data (Winkler and Wang, 1993).

In northeastern Siberia during the LGM, subpolar desert and montane tundra (Grichuk, 1984) covered the lowlands between mountain massifs supporting isolated ice caps and valley glaciers (Glushkova, 1994). Open forests of spruce and birch grew on the southern tip of Kamchatka and in the Amur River region (Fig. 2). At the same time, the upper Kolyma River valley supported Arremisia-Gramineae tundra growing on poorly-developed, frost-disturbed soils (Lozhkin et al., 1993).

The first indications of climatic amelioration in the upper Kolyma valley occurred ca. 12,500-12,000 BP with a transition from herb tundra to Betula shrub-tundra

D.H. Mann and T.D. Hamilton: Paleotinvironments of the North Pacific Coast

i/ 60°N-

Sea

‘Kamchatka

fl\-1’00 m bathymetric

451

FIG. 2. Vegetation zones of northeast Asia during the last glacial maximum (LGM) ca. 22,000-18,000 BP as inferred by Grichuk (1984). Heavy lines depict modem coastline; black areas represent glaciers.

(Lozhkin et al., 1993). Larix duhurica forests became established ca. 11,600 BP indicating summer temperatures of at least 12°C. The last major element of the modem vegetation to arrive, the stone pine (Pinus pumilu), reached the upper Kolyma River about 9000 BP. The autecology of this conifer implies an amelioration in winter conditions that may have included deeper snow packs and warmer temperatures (Lozhkin et al., 1993). Pollen accumulation rates suggest that altitudinal treeline may have risen in the Kolyma area between 9000 and 7000 BP (Lozhkin et al., 1993) accompanying a northward advance in latitudinal treeline on the Kolyma plain (Kind, 1967).

Deep-sea sediment records indicate that full-glacial conditions persisted in the northwest Pacific Ocean until ca. 14,000-13,000 BP, when the first of three abrupt changes in water temperature and/or salinity occurred (Fig. 3). A major influx of ice-rafted debris between 13,000 and 12,000 BP signaled rapid retreat of calving glaciers somewhere in the North Pacific basin (Keigwin et al., 1992). After about 3000 years of stable or slowly- declining temperature, water temperature warmed rapidly around 10,500 BP (Keigwin et al., 1992).

A possible late Glacial cold episode analogous to the Younger Dryas in the North Atlantic is currently a topic of active research in the northwest Pacific. Chinzei et al. (1987) and Kallel et al. (1988) interpret isotope and foraminifera data to indicate a re-advance of subpoiar water along the Japanese coast between 11,000 and 10,000 BP, correlative with the re-advance of the polar front in the North Atlantic during the Younger Dryas. A Younger Dryas signal may be present in the Chinese loess record (An et al., 1993) and possibly in pollen records near Beijing (Liu, 1988). However, Keigwin and Gorbarenko (1992) ascribe possible Younger Dryas analogs near northern Japan to an episode of freshwater discharge from the Sea of Japan.

During the early Holocene, sea-surface temperatures were 1-4”C lower than today (L.E. Heusser and Morley, 1990). The final warming step occurred in the deep waters of the northwest Pacific between 8000 and 6000 BP (Fig. 3). Temperate forest communities reached their maximum abundance around the Sea of Okhotsk during the middle Holocene, whereas the spread of spruce forest after 4000 BP probably accompanied regional cooling (Morley and Heusser, 1991). Northern expansion in the


h

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OTcl-41;.;r:t

15) -- cooler warmer

PIG. 3. Oxygen isotope values of foraminifera from a deep-sea core taken southwest of Kamchatka (Keigwin et al., 1992). Shaded areas mark intervals of rapid warming at 14,000-13,000 BP, 10,500-10,000 BP, and again ca. 7000 BP. The term ‘ka

BP’ signifies millennia before 1950 A.D.

range of Cryptomeria japonica, a tree requiring moisture from heavy winter snowfall, also occurred after 4000 BP, suggesting cooler, wetter winters after that time (Tsukada, 1986).

On Sakhalin Island, shrub birch, larch and Pinus pumila were replaced by spruce after 10,000 BP (Khotinskiy, 1984). Deciduous tree species reached their greatest Holocene abundance between 9500 and 8500 BP suggesting that summers were warmer than today. Historical records of climate in northern China document declining temperatures corresponding to the onset of Neoglaciation after 5000 BP (Zhang, 1991) or, alternate- ly, after 3000-4000 BP (Liu, 1988; Winkler and Wang, 1993). Low temperatures in these Chinese records between 1200 and 1900 A.D. (Zhang, 1991) record the occurrence of Little Ice Age cooling in northeast Asia (Fig. 4).

Marginal seas (the Okhotsk, Japan and Bering Seas) (Fig. 1) comprise much of the coastline of the northwest and northern Pacific. The paleoenvironmental histories of these seas were, in some cases, quite different from events in the open Pacific. The late Pleistocene history of the Japan Sea illustrates the potential complexity of climatic and oceanographic changes in the marginal seas caused by interactions between changing sea level and shallow straits. The Japan Sea today experiences a continuous inflow of warm water from the Tsushima Current, a western branch of the Kuroshio Current. However, during the LGM, falling sea level cut off the Tsushima Current, causing sea surface temperatures to drop and salinity to decline under the influence of increased freshwater input from the diverted Amur River (Oba et al., 1991; Keigwin and Gorbarenko, 1992). Primary productivity fell in the Japan Sea because the cold, low-salinity surface waters prevented deep convection Sometime between 20,000 and 10,000 BP, the cold Oyashio Current entered the northern Japan Sea restoring

3000 2000 1000 0 500 1000 1500 1750 1800 1900 1950 B.CjA.D.

Calendar Years PIG. 4. Historical records of late-Holocene temperature fluctuations in China (redrawn from Zhang, 1991). This curve depicts temperature changes during the Neoglacial Interval (ca. 3000-O B.C.), including the Little Ice Age (ca. 1200 to

1900 A.D.).

D.H. Mann and T.D. Hamilton: Paleoenvironments of the North Pacific Coast

deep ventilation and productivity (Oba et al., 1991). The Oyashio Current was excluded and flickering inflow of the Tsushima Current began after 10,000 BP. Modern oceanographic conditions were not established until after 8000 BP (Oba et al., 1991). Sea level was l-l.5 m higher than present in the Sea of Japan during a mid-Holocene warm interval dated by thermophilous bivalves between 5000 and 7000 BP (Jones et al., 1994).

BERING SEA REGION

During the last ice age, global sea level fell to ca. -120 to -130 m (Fairbanks, 1989; Bard et al., 1990), exposing large portions of the continental shelf beneath the shallow Bering and Chukchi Seas. Exposure of the continental shelf had two major effects in the Bering Sea region: it made the climates on Chukotka and western Alaska more continental and it unified terrestrial habitats on both sides of the Bering Strait. At a hemispheric scale, changes in the amount of relatively fresh Pacific water flowing northwards through Bering Strait into the Arctic Ocean possibly influenced North Atlantic thermohaline circulation and climate (Shaffer and Bendtsen, 1994).

0

20

E 40

rc ; 6o

80

After a low stand of about -120 m, reached ca. 18,000 BP (Fig. 5), eustatic sea level rose rapidly and reached near-modern levels by 4000 BP (Mason and Jordan, 1993). Rising sea level progressively flooded

453

the Bering Strait separating Chukotka and Alaska between 11,000 and 8500 BP (Elias et al., 1992). Only after 8500 BP could the modern circulation patterns between the Bering, Chukchi and Beaufort Seas have become established. In contrast to the relative tectonic stability of the Bering Platform, the Aleutian Arc is seismically active and the relative sea-level history is in part controlled by tectonism (Hamilton and Thorson, 1986; Taber et cd., 1989). Hence it is unwise to apply Fairbank’s (1989) eustatic sea-level curve to this portion of the Bering Sea.

During the LGM, the southeastern Bering Sea and Sea of Okhotsk had ice cover for perhaps 9 months of the year (Sancetta, 1983; Sancetta et al., 1985). During the summer, ice floes were advected across the southern Bering Sea and Sea of Okhotsk into the northwest Pacific Ocean, creating a surface layer of cold, low-salinity water that prevented vertical mixing and lowered primary productivity (Sancetta, 1983). According to Sancetta’s

Sea Level

Shpanberg Straft

ong / Herald I 3nadyr Straits <.

0 2 4 6 8 10 12 14 16 18

AGE (ka B.P.) FIG. 5. The bathymetry of the Bering Strait region (panel a) (redrawn from Porter, 1988) and global (eustatic) sea level approximated by the Barbados sea-level data of Fairbanks (1989) (panel b). In panel b, horizontal lines are the sill altitudes of straits within the Bering and Chukchi Seas: L = Long Strait, H = Herald Strait, A = Anadyr Strait, B = Bering Strait, and S = Shpanberg Strait. The Chukchi Sea lies north of the Bering Strait and the Bering Sea lies to the south. In panel a, vertical arrows point to the probable date these straits were flooded in postglacial times. The dark circle and arrowhead represent the age of terrestrial peats presently at a depth of -50 m in the Chukchi Sea (Elias et al., 1992). Time scale is in radiocarbon years BP See Bard et al. (1990) for the slightly revised, Barbados sea-level chronology in U-Th

years BP


model, cooling of North Pacific waters may have intensified the Polar Front and shifted it southward. This caused increased storminess in the mid-latitude North Pacific, decreasing evaporation from the higher latitude Pacific, and consequently decreasing precipitation in Beringia during the LGM. Ocean circulation in the southern Bering Sea was radically different during the LGM. Lowered sea levels and expanded glaciers over the Aleutians and the Alaska Peninsula prevented the Alaska Coastal Current from entering the southeastern Bering Sea as it does today (Fig. 1) through passes in the eastern Aleutians. While the shelf was subaerial, the Yukon and Kuskokwim Rivers entered the southern Bering Sea rather than having most of their water shunted northward through the Bering Strait as they do today (Sancetta et al., 1985). The resulting fresh water lid would have intensified sea-ice formation. The Pleistocene to Holocene transition was accompanied by decreasing sea ice and re-entry of the Alaska Coastal Current into the southeastern Bering Sea (Sancetta and Robinson, 1983). Plankton productivity increased with the disappearance of the freshwater lid around 10,000 BP (Sancetta et al., 1985).

ranges of beavers, tree birch, and cattails (McCulloch and Hopkins, 1966; Hopkins, 1982) and thickening of the active layer above permafrost (McCulloch and Hopkins, 1966). Extra-limital Populus wood in northern Alaska dates between 11,300 and 8400 BP (Hamilton and Fulton, in press, Table 4b). Unfortunately the existing age estimates bracketing the Populus peak in pollen diagrams are from bulk radiocarbon dates of dubious accuracy from lake sediments often low in organic con- tent. Beetle remains dated to 11,300-l 1,000 BP from terrestrial peats on the Chukchi shelf suggest that summer climate was appreciably warmer than today (Elias et al., 1992).

Full-glacial, terrestrial climate was drier and colder than today, reflecting in part the increased continentality caused by the vastly increased land area of the exposed Bering shelf. Dry, windy climate caused widespread deposition of loess, sand dunes, and sand sheets throughout Beringia (Hopkins, 1982; Lea and Waythomas, 1990; Hamilton and Ashley, 1993). In modem loess-dominated ecosystems on the Arctic Coastal Plain of northern Alaska, continual disturbance by loess and eolian-sand deposition maintains the vegetation in an early succes- sional state with a significant component of grass species (Walker and Everett, 1991). From pollen records. Colinvaux (1986) and Matthews (1982) advocate a topographically- and regionally-complex mosaic of shrub- and herb-tundra vegetation across Beringia during the LGM, whereas Guthrie (1990), reinterpreting much of the same data, advocates a widespread, steppe-tundra biome with no clear modem analogs.

The Bering and Chukchi coastlines did not experience the subsequent episode of increased summer warmth ca. 8500-5500 BP that caused treeline in Keewatin, on the MacKenzie Delta (Ritchie and Hare, 1971), and on the Kolyma plain (Kind, 1967) to advance hundreds of kilo- meters. Ritchie and Hare (197 1) suggest that the nortb- ward shift of the average position of the Polar Front in western Alaska during this interval is consistent with increased onshore flow of cool, cloudy, maritime air from the Bering and Chukchi Seas. Lozhkin ef al. (1993) suggest the flooding of the Bering and Chukchi platforms in the early Holocene caused a transition from continental to maritime climate. Certainly the timing of the interval of maximum post-glacial warmth seems to have been time-transgressive in the Beringia region.

Forests along the eastern coast of the Bering and Chukchi Seas developed their modem distribution and species composition by ca. 6000-5000 BP (Anderson. 1985, 1988). On St. Paul Island, no vegetation change is evident in the pollen record since 9500 BP. However, on St. George Island, slight changes in the pollen record during the middle to late Holocene may record a shift to cooler Neoglacial climate in the southern Bering Sea region (Ager, 1982).

Pollen records from western Alaska indicate that major climatic changes occurred 14,000- 13,000 and 11 ,OOO-9000 BP in the Bering-Chukchi Sea region (Ager, 1982, 1983; Ager and Brubaker, 1985; Anderson and Brubaker, 1993). Pollen diagrams from the Yukon delta, Norton Sound, and Kotzebue Sound areas document a shift from predominately herbaceous tundra vegetation to mesic, shrub-birch tundra ca. 14,00&13,000 BP (Lozhkin et al., 1993).

Existing spit complexes on the coast of northwestern Alaska originated as sea level stabilized after 4400 BP (Mason and Jordan, 1993; Plug and Mann, 1994). Disconformities and truncations in many of these complexes may be the results of erosion during intervals of increased summer storminess between 3300-l 700 BP and 1200-900 BP (Mason and Jordan, 1993 ). These authors correlate increased storminess in the Bering and Chukchi Seas with northward shifts in Pacific storm tracks. Mason and Jordan’s (1993) intervals of increased storminess show weak correspondence to intervals of dune activity and cirque-glacier expansion in northern Alaska (Galloway and Carter, 1993).

Maximum post-glacial warmth occurred in notthwest- We have no information on changes in the extent and em Alaska before 8000 BP and possibly before 10,000 seasonal persistence of sea ice during the Holocene in the BP (Brubaker et al., 1983; Lozhkin et al.. 1993; Edwards Bering and Chukchi Seas. Both these phenomena vary and Barker, 1994) perhaps in response to a solar-radiation according to regional climatic patterns, including strength maximum centered on 10,000 BP (Ritchie et al., 1983; and persistence of the Siberian High and strength and ori- Barnosky et al., 1987; Bartlein et al., 1991). This warm entation of the Aleutian Low (Niebauer et al., 1989). interval saw the establishment of Populus woodland in Polynyas, areas of open water or thin ice found at pre- areas beyond the present range of this species (Brubaker dictable, recurrent locations within the regional cover of et al., 1983; Ager and Brubaker, 1985; Anderson, 1988; sea ice, presently occur in the lee of islands and capes in Anderson et al., 1990), as well as expansions in the the Bering and Chukchi Seas. Polynya geography is

D.H. Mann and T.D. Hamilton: Paleoenvironments of the North Pacific Coast 455

closely related to settlement patterns& the eastern Canadian High Arctic (Schlederman, 1980), and probably affected settlement in the Bering Sea region as well. Polynya locations depend on wind intensity and direction (Stringer and Grove, 1991).

Three alternative reconstructions exist for glacier extent during the LGM in Beringia. In the most widely accepted view, glaciers were limited to high mountains on both sides of the northern Bering Sea and Chukchi Sea during the LGM (Velichko er al., 1984; PCwt, 1975; Hamilton, 1994). On the Seward Peninsula, snowlines fell to 300-600 m above present sea level (Kaufman and Hopkins, 1986) but most glaciers terminated within mountain valleys. Along the southern margin of the Bering Sea, local ice caps inundated the Aleutian Islands, coalescing within island groups and possibly forming a floating ice shelf in the southern Bering Sea (Thorson and Hamilton, 1986).

In an alternate model, Gros’vald and Vozovik (1984) and Gros’vald (1988) advocate a marine-based ice sheet covering the southern Bering Shelf and flowing southward through the Aleutian Islands. However, no field data exist to support this hypothesis, and evidence against the marine ice-sheet model is presented by Hopkins et al. (1992). The third model is a modified marine ice-sheet hypothesis described by Hughes and Hughes (1994). It states that a continuous mountain-glacier complex rimmed the Bering Sea coast northeast of Kamchatka during the LGM, merging with a marine ice sheet in the Beaufort and Chukchi Seas. Both the Hughes and the Gros’vald models predict that exchange of terrestrial

\ il60"E u \

North Pacific Ocean .-I

biota ,was blocked by glacier barriers even when the Bering platform was dry land. However, preliminary mapping of moraines in Chukotka suggest that the extensive mountain-glacier complex hypothesized by Hughes and Hughes (1994) did not exist during the LGM (Glushkova, 1994; Heiser and Roush, 1994).

General circulation models suggest that the Aleutian Low intensified during the LGM and shifted southwards from its present position (Manabe and Hahn, 1977; Kutzbach et al., 1993). Modeling also suggests that the elevated mass of the Cordilleran and Laurentide Ice Sheets caused a split in the polar jet stream during the last ice age, with a northern branch streaming northeastward over the Bering Strait region (Kutzbach, 1987; Bamosky et al., 1987; Bartlein et al., 1991; Kutzbach et al., 1993). However, actual field evidence for this split jet is lacking or contradictory. For instances, the orientation of full-glacial dune fields in Alaska indicates that the pre- dominant, dune-forming winds blew out of the northeast (Hopkins, 1982; Lea and Waythomas, 1990). Also, snowline altitudes in glaciated areas in southeastern Beringia during the LGM were lower but of similar regional pat- tern as today, implying that LGM storm tracks were like those of today (Hamilton and Thorson, 1983; Detterman, 1986; Mann and Peteet, 1994).

SOUTHWEST ALASKA

Southwest Alaska (Fig. 6) extends from the mouth of Cook Inlet to Unimak Pass in the inner Aleutian Islands, and includes the mountainous Kodiak archipelago in the

-

sland .”

@C% Trinity Isla Ads

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. . _*.._* . . . . . . . . . - . . . edge “’

*. . * . . . . * ($I+ 16O’W

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FIG. 6. Southwestern Alaska. Large lakes presently dammed behind moraines left by late Pleistocene glaciation are shown in black.


Gulf of Alaska, the Alaska Peninsula, and the low-lying faces available as refugia would have been steep, alpine shores of Bristol Bay. Southwest Alaska was the extreme ridges with a total surface area of ca. 900 km2 (Mann southeastern comer of Beringia when sea level was low- and Peteet, 1994). To the east, the ice covering the ered to -120 m (Fairbanks, 1989) during the LGM. Alaska Peninsula merged with other parts of the Fluctuating glaciers in southwest Alaska acted as an ice Cordilleran glacier complex. To the west, along the inner gate that opened and shut between the unglaciated low- Aleutian Islands, it probably continued as a narrow belt lands of western and central Beringia and the Gulf of of small interconnected, island ice caps (Thorson and Alaska coastline. Hamilton, 1986).

During the LGM, the Alaska Peninsula was largely covered by an elongate glacier complex fed by snow catchment areas along the crest and southern flank of the Aleutian Range as well as in the mountains of Kodiak Island (Fig. 7). Ice was more extensive on the Gulf of Alaska side of the peninsula than on the Bristol Bay side. The Gulf of Alaska was the major source of precipitation then, as it is today. Snowlines rose from ca. 350 m on Kodiak Island to more than 900 m on the north side of the Alaska Peninsula, representing a lowering of the snowline about 300-600 m below present-day values (Detterman, 1986; Mann and Peteet, 1994).

Reconstruction of ice thickness using trimline and moraine altitudes onshore, bathymetric charts on the continental shelf, and theoretical calculations of basal- shear stress indicate that the glacier complex covering the Alaska Peninsula and Kodiak Island flowed to the outer edge of the continental shelf where it calved icebergs into the open Pacific (Fig. 7). Low ice saddles connected the mountains of Kodiak Island to the mainland and bridged the entrance of lower Cook Inlet. Small ice-free areas persisted through the last glacial maximum on southwest Kodiak Island (Karlstrom, 1969; Mann and Peteet, 1994). However, the ice-free areas on Kodiak Island were largely filled by proglacial lakes and the sur-

The last glacial maximum was reached in southwest Alaska after 23,000 BP, when ice advancing toward the outer continental shelf dammed valleys on western Kodiak Island and impounded proglacial lakes (Mann and Peteet, 1994). Glaciers had retreated from the Low Cape area of southwestern Kodiak Island before 14,700 BP (Mann and Peteet, 1994). The Kodiak Island ice cap expanded to a late-glacial maximum at 13,400 BP, retreated an unknown distance, and then expanded to a slightly less extensive position at 11,900 BP (Mann and Peteet, unpublished d&z). By 10,000 BP the Kodiak ice cap had undergone final collapse and the inner fjords of the island were ice free (Mann and Peteet, unpublished data). In the Katmai area on the Alaska Peninsula, glaciers re-advanced sometime between 12,000 and 10,000 BP and perhaps again between 9800 and 9500 BP (Pinney and Beget, 1991).

A drier than modem climate is suggested by pollen analysis of peat deposited near sea level on western Kodiak Island between 15,000 and 10,000 BP (Peteet and Mann, unpublished data). Vegetation was probably mesic tundra dominated by mosses and ericad shrubs, with much open ground in near-coastal environments where sand-sheet sediments accumulated. Widespread peat accumulation began on the Kodiak archipelago after

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FIG. 7. Maximum extent of glaciers in southwest Alaska during the last glacial maximum, 22,000--18,ooO BP. From Mann and Peteet (1994).


13,000 BP Macrofossil and pollen evidence suggests an 800 year interval of cooler, drier climate between ca. 10,800 and 10,000 BP, perhaps correlative with the Younger Dryas episode (Peteet and Mann, 1994). Pollen and macrofossils record the sudden establishment of modem plant communities on western Kodiak Island at 10,000 BP

On the southeastern slopes of the Ahklun Mountains north of Bristol Bay (Fig. 6), mesic herb tundra dominated the landscape ca. 13,000-9800 BP (Hu et al., 1995). A climate reversal perhaps correlative with the Younger Dryas episode occurred between 10,800 and 9800 BP Alder arrived and became important in the vegetation by ca. 7400 BP. White spruce arrived in the region 5500-4000 BP and reached its modem distribution by ca. 2000 BP (Ager, 1983; Hu et al., 1995).

Peat sections from the Shumagin Islands and the Alaska Peninsula west of Kodiak Island record a major vegetation change at 10,000 BP from grass/ Artemisialwillow to birch/Empetrumlgrass-dominated vegetation (C.J. Heusser, 1983a). Shrub alder (Alms sinuata) reached the Shumagin Islands around 5000 BP but persisted in relatively low numbers until the last several centuries when it increased markedly (C.J. Heusser, 1983a).

THE GULF OF ALASKA

The Yakataga Formation, which is exposed along the coastline of the eastern Gulf of Alaska, documents the onset of tidewater glaciation during latest Miocene time that accompanied the uplift of the coastal mountains (Eyles et al., 1991; Lagoe et al., 1993). Initial tidewater glaciation occurred about 6.7-5.0 Ma, and a major increase in the intensity of glaciation began between 2.48 and 3.5 Ma following a mid-Pliocene warm interval (Lagoe et al., 1993). Bottom sediments in the Gulf of Alaska record numerous episodes of subsequent iceberg dispersal from calving glaciers (von Huene et al., 1976; von Huene, 1989; Carlson, 1989). Large glacial troughs that extend to the outer continental shelf offshore of major fjord systems indicate extensive glaciations whose ages are poorly constrained (Carlson, 1989; Carlson er al., 1982; Molnia, 1986).

The present climate of this large, sea-ice free sector of the Pacific Ocean stretching between Kodiak Island and the Queen Charlotte Islands is dominated by interactions between ocean-current systems (Fig. 1) and the prevailing westerly winds. Oceanographic and atmospheric events here are critical to climates in downwind North America. The northward-flowing Alaska Current/ Alaska Coastal Current (Reed and Schumacher, 1987) transports large amounts of heat into the coastal environments of southern and southeastern Alaska. Current systems in the Gulf of Alaska are sensitive to small changes in sea surface temperature and frontal positions (Reed and Schumacher, 1987; Royer, 1989; Ebbesmeyer and Ingraham, 1992). Heat is also trans- ported into the region by cyclonic storms traveling northeastward in the prevailing Westerlies. Storm tracks

in the, North Pacific are closely linked to sea surface temperatures (Namias, 1970; Emery and Hamilton, 1985). High coastal mountains bordering the Gulf of Alaska pose major barriers to these storms, which dump their precipitation and release large amounts of latent heat. The combined result of the ocean and atmospheric circulation systems is presently a wet, maritime climate which is comparatively mild given the region’s latitude. Temperatures and precipitation tend to increase from west to east across the northern Gulf of Alaska (Wilson and Overland, 1987).

Some aspects of the present climatic regime were probably also present during the LGM. For instance, sea-surface temperature reconstructions indicate that sea ice was never present in the Gulf of Alaska (CLIMAP, 1981). However, near-surface waters were of lower salinity during LGM summers than today, retarding convection and lowering sea-surface temperatures (Zahn e? al., 1991). Sea-surface temperatures at the LGM in the Gulf of Alaska were probably 34°C lower than at present (CLIMAP, 1981). Global circulation modeling hypothesizes that during the LGM both January and July temperatures over the northern Gulf of Alaska were 2-4”C cooler than today (Kutzbach et al., 1993). Although sea ice probably was absent, icebergs were common, especially during periods of glacier retreat (Keigwin et al., 1992). Reconstructions of sea- surface temperatures based on microfossils suggest that the North Pacific Polar Front shifted southward by only 5” of latitude or less between the present and 18,000 BP (Thompson, 198 1 ), however, ocean circulation intensified and thermal gradients steepened along the front (Moore et al., 1980; Sancetta, 1983). Although agreeing in suggesting an intensification of the Aleutian Low during the LGM, global circulation models hypothesize a southerly shift in the Polar Front and its associated storm tracks by 10-20” latitude during the LGM (Manabe and Hahn, 1977; Kutzbach, 1987; Kutzbach et al., 1993). The precipitation now nourishing glaciers and raising temperatures in the eastern Gulf of Alaska would have been diverted further south into the western United States at that time (Bamosky et al., 1987). The GCM models further suggest that an anti-cyclone over the North American ice sheets created a southerly flow, drawing relatively warm, maritime air masses northward into the eastern Gulf of Alaska (Manabe and Broccoli. 1985; Kutzbach, 1987; Kutzbach et al., 1993). Hopefully, further studies of marine microfossils will resolve the discrepancy between computer simulations and field data.

Oceanographic and atmospheric processes in the Gulf of Alaska control climate along its shoreline yet the exact mechanisms and actual chain of prehistoric events remain obscure. For instance, dendrochronologies from coastal sites between western Washington and Kamchatka record a cooling of the North Pacific in the mid- to late 1800s AD (Jacoby et al., 1994) that probably caused glacier advances throughout southern and southeast Alaska (Wiles and Calkin, 1994). Throughout the Holocene, changes in the seaward influence of continental high

458 Quaternary Scienc :e Reviews: Volume 14

pressure over northern British Columbia and the southern Yukon may have modulated the frequency of storms moving onshore in southeastern Alaska (M. M. Miller and Anderson, 1974). The frequency of these blocking highs (Wilson and Overland, 1987) over the eastern Gulf of Alaska in winter could have exerted an important control over the timing of Neoglaciation in southeastern Alaska.

COASTAL, SOUTH-CENTRAL ALASKA

We define coastal, south-central Alaska as the region between Cook Inlet and Yakutat Bay (Fig. 8). This is a high-wave-energy coast closely backed by the Kenai, Chugach, and St. Elias Mountains, which include some of the highest near-coastal peaks in the world. This region contains the most extensive present-day glaciers in Alaska, including two vast Piedmont ice bodies - Bering Glacier and Malaspina Glacier. Protected shorelines are rare compared to southwest Alaska, the Bering Sea, or the Seas of Japan and Okhotsk. The exception is Prince William Sound, a glacially-scoured fjordland of numerous rocky islands protected from the open Pacific by Montague and Hinchinbrook Islands. From Montague Island westward, the coastline is primarily rocky and eastward it is mostly composed of sand and gravel originating from the Copper River and near-coastal, Piedmont glaciers.

The ongoing tectonism associated with subduction of the Pacific plate along the eastern end of the Aleutian Trench (von Huene, 1989; Page et al., 1991) is an important factor in relative sea-level history in south-central Alaska. Along much of the western portion of this coastline, the land is presently subsiding while farther east it is rising (Plafker, 1969, 1990). Coseismic uplift or down- warp ranges up to several meters, and these abrupt dis-

placements are separated by gradual interseismic movements of the crust, usually in the direction opposite to the coseismic displacement (Plafker, 1969). Great earthquakes tend to be spaced ca. 1000 years apart (Plafker et al., 1992; Combellick, 1993).

During the LGM, glaciers virtually covered the Kenai Peninsula and filled Cook Inlet (Karlstrom, 1964; Hamilton, 1994; Reger and Pinney, in press). The only ice-free areas not covered by proglacial lakes were small- upland areas northwest of the Kenai Mountains. The largest ice bodies in Cook Inlet flowed eastward across the inlet from sources in the Alaska and Aleutian Ranges (Schmoll and Yehle, 1986). They extended onto the Kenai Lowland, where they locally merged with ice from the Kenai Mountains and blocked drainages to impound a chain of lakes (Reger and Pinney, in press). Although no dates on this event are available from Cook Inlet, they probably are close to the bracketing ages of ca. 25,000 and 17,000 BP on the last glacial maximum (McKinley Park I Stade) of the central Alaska Range (Ten Brink and Waythomas, 1985).

The glacier complex in Cook Inlet began to break up before 16,500-16,000 BP, when marine waters invaded the lower inlet (Reger and Pinney, in press). Relative sea level was at least 25 m above its present position at that time, owing to glacial-isostatic depression of at least 112 m. During a subsequent stillstand or re-advance prior to 14,900 BP, glaciers from the north and east coalesced in uppermost Cook Inlet while marine waters invaded to present-day Anchorage (Reger and Pinney, in press). During a final glacial re-advance into Cook Inlet, which occurred sometime between 14,000 and 11.700 BP (Schmoll et al., 1972), glaciers extended into upper Knik and Tumagain Arms but did not coalesce (Schmoll and Yehle, 1986; Reger and Pinney, in press). Glacier retreat

FIG. 8. Coastal. south-central Alaska, showing existing glaciers.


from Knik Arm was nearly complete by 11,400 BP and from Tumagain Arm by 10,700 BP (Regei and Pinney, ih press). Spruce forest invaded the region by about 8400 BP (Ager and Sims, 1984).

The extent of glaciers onto the continental shelf off of south-central Alaska during the LGM is a topic of continued speculation (Pew& 1975; Coulter er al., 1965) LGM ice probably flowed to the outer edge of the continental shelf in the Gulf of Alaska given the probable extent of ice in the Kodiak region (Mann and Peteet, 1994), the abundant evidence for high-altitude glacial erosion on Knight and Montague Islands in Prince William Sound (Mann, unpublished observations), and submarine features suggestive of glacial origin on the continental shelf (von Huene, 1966; Carlson et al., 1982; Carlson, 1989; Molnia, 1986; Sirkin and Tuthill, 1987).

This ice mass had broken up and retreated out of the inner fjords of Prince William Sound by about 14,000 BP (Sirkin and Tuthill, 1987; Reger, 1991). Some coastal glaciers may have persisted until nearly 10,000 BP (Ager, 1992). The interval lO,ooo-6ooo BP was probably slightly warmer and drier than today in coastal south-central Alaska with glaciers persisting in retracted positions near or upvalley from their present positions (Calkin, 1988; Wiles and Calkin, 1990).

Three major intervals of Neoglaciation in the Southern Kenai Mountains are documented by Wiles and Calkin (1990, 1993, 1994). The earliest advance occurred about 3700-3600 BP A subsequent interval of expansion was underway by about 500 A.D. and ended around 900 A.D. The latest and best-documented phase of glacier expansion occurred between ca. 1200 A.D. and 1890 A.D. during the Little Ice Age, when equilibrium-line altitudes (roughly equivalent to late summer snowlines) were depressed loo-150 m below present levels. The most recent maximum position of glaciers in the Kenai fjords was reached in the late 1800s A.D. and was followed by general retreat.

Near-synchronous glacial advances during the last two millennia in Icy Bay were reported by Porter (1989), who dated the outermost recognizable moraine complex between 400-850 A.D. Little Ice Age advances probably began in the 13th century in upper parts of Icy Bay, and they culminated in the early 1800s A.D. Glacier retreat began in Icy Bay by the 1880s A.D. (Porter, 1989). Pollen records support the glacial data in showing wetter and cooler climates associated with the Neoglacial Interval (Heusser, 1985).

Interpretation of Holocene dynamics of alpine glaciers on the Gulf of Alaska coastline is complicated by two factors. Calving glaciers are subject to non-climati- cally controlled instabilities causing advances and retreats not affecting landbased glaciers (Mann, 1986a; Meier and Post, 1987). Second, glaciers seem to have differed in their responses to climate change depending on their geographical position. Glaciers on the seaward flank of the Kenai Mountains seem to have advanced in response to the increased winter precipitation occurring during relatively warm intervals over the last 2000 years while glaciers on the western flank advanced in

response to lower summer temperatures (Wiles and Catkin, 1994). These differential responses may be common to glaciers throughout the gulf coast and com- plicate the interpretation of glacier behavior in terms of climate.

The oldest continuous pollen record in coastal south- central Alaska comes from Hidden Lake on the Kenai Peninsula and dates to 14,000 BP (Ager, 1983; Ager and Sims, 1984; Anderson and Brubaker, 1993). Initial herb tundra with a few shrubs of willow and ericads was replaced by shrub tundra dominated by birch in the interval 13,700-10,300 BP Between 11,000 and 8000 BP, the vegetation changed to a mixture of shrub tundra and deciduous forest communities dominated by poplar and willow. Alder invaded the area between 9000 and 8000 BP (Ager, 1983).

Sedge, willow, and ferns were the earliest plants colo- nizing the inner fjords of Prince William Sound ca. 10,000 BP (C.J. Heusser, 1983b, 1985). In Port Wells, alder become widespread after 8300 BP Sitka spruce and mountain hemlock arrived only ca. 3000 BP (C.J. Heusser, 1983b), although they may have entered the outer islands of Prince William Sound much earlier. Only after 2000 BP did Sitka spruce, mountain hemlock, and western hemlock assume an important role in the vegetation of the inner fjords of Prince William Sound (C.J. Heusser, 1985).

At Icy Cape, in the eastern Gulf of Alaska near Mount St. Elias, peat accumulation began around 10,800 BP under a vegetation of coastal tundra dominated by sedges and crowberry (Empetrum nigrum) (C.J. Heusser, 1960). Tundra was replaced by alder and ferns ca. 10,000 BP (Peteet, 1986). A Hypsithermal Interval, warmer and drier than today, occurred at Icy Cape between ca. 9000 and 7500 BP (Peteet, 1986). Sitka spruce arrived ca. 7600 BP, mountain hemlock ca. 5000 BP, and western hemlock ca. 3800 BP (Peteet, 1986). These arrival times probably were affected by immigration lags caused by long-distance dispersal from full- glacial refugia in western Washington (C.J. Heusser, 1985; Peteet. 1991).

SOUTHEAST ALASKA

Southeast Alaska is the ‘panhandle’ region of Alaska that extends from Yakutat Bay southward to Dixon Entrance (Fig. 9). Most of this region consists of the Alexander Archipelago: mountainous, glacially-scoured islands crisscrossed by deep fjords originating in the Coast Mountains. The extreme northwest part of the Alaskan panhandle is a narrow-coastal foreland running along the seaward flank of the Fait-weather Range and including the fjord systems of Lituya Bay and Yakutat Bay.

The Quatemary history of southeast Alaska is poorly understood. During the LGM, coalescing ice caps flowed westward out of the Coast Mountains and the island massifs of the Alexander Archipelago onto the continental shelf of the Gulf of Alaska. Large outlet glaciers occu- pied what are today the major fjord systems in southeast


;Q”**d Charlottei Islands I

FIG. 9. Southeastern Alaska, the region between Yakutat Bay and Dixon Entrance and seaward of the Coast Mountains.

Alaska (Goldthwait, 1987). Reconstruction of glacier extent on the continental shelf around Lituya Bay and Cross Sound, based on ice flow models and onshore mapping of glacial features (Derksen, 1976; Mann, 1986b), suggests that glaciers reached only the inner continental shelf in areas lying between major fjord entrances during the LGM. Though undocumented to date, it is possible that areas lying between Piedmont ice lobes on the continental shelf remained ice-free and subaerial during the LGM.

Deglaciation probably was rapid in southeast Alaska, with iceberg calving causing glacier termini to retreat to near their modem positions by about 13,500 BP, even in some of the inner fjords like Glacier Bay (McKenzie and Goldthwait, 1971; Goldthwait, 1987). Rapid ice retreat was followed by a marine transgression reaching .50- 230 m above present sea level on the isostatically- depressed coast (R.D. Miller, 1972, 1973a,b, 1975). At the south end of the Chilkat Peninsula, three raised ter- races indicate a complex relative sea-level history between 13,000 and 10,ooO BP (Ackerman et ul., 1979). Based on the rates of isostatic rebound in southwest British Columbia (Clague, 1989; Easterbrook, 1992), modem sea level was probably reached by 9000 BP in most areas of southeast Alaska not affected by tectonic uplift or subsidence. However, such generalized predic- tions may be premature. For instance, Mobley (1988) found major differences over a distance of less than 100 km in the sea-level histories of Heceta and Prince of Wales Islands in the southern Alexander Archipelago. The more seaward Heceta Island experienced an early

Holocene marine transgression culminating ca. 8500 BP. A transgression of similar duration and magnitude occurred on the Queen Charlotte Islands, possibly in response to a migrating forebulge created when mantle material that had been displaced westward during maximum ice loading diffused back towards the mainland after deglaciation (Clague, 1975, 1983; Clague et al., 1982). In contrast, sea level around Prince of Wales Island reached near-modem levels before 9000 BP and no subsequent transgression was detected. Mobley’s (1988) findings point out the complexity of sea-level changes over short distances in southeast Alaska, which are probably due to interactions between glacial history, crustal response to ice loading, and tectonism. Tectonic factors are certainly a major factor in sea-level dynamics in this region today (Hudson et al., 1982; von Huene, 1989).

Herb-tundra vegetation colonized the lowlands of southeast Alaska following glacier retreat and the subsequent marine transgression. Lodgepole pine and alder reached the northern archipelago by 12,500 BP (Engstrom et al., 1990; Cwynar, 1990; Anderson and Brubaker, 1993). A climate reversal, perhaps correlative with the Younger Dryas episode in northwestern Europe. occurred between 10,800 and 9800 BP when lodgepole pine parkland was replaced by shrub- and herb-dominated tundra (Engstrom et al., 1990). After 10,000 BP, Sitka spruce, western hemlock, and mountain hemlock arrived in that order. Tree arrival times were time-transgressive along the coast because of the time it took them to spread northward from southern refugia (Peteet, 1986, 1991; Cwynar, 1990).

Western red cedar, an important tree for northwest coast cultures in historic times, expanded its range into the south part of southeast Alaska only during the middle Holocene (Hebda and Mathewes, 1984). The northward migration of tree taxa was only complete in the northern part of southeast Alaska by ca. 4000 BP (Peteet. 1986; C.J. Heusser, 1985; Cwynar, 1990). Despite the presence of small ice-free areas along the outer coast of Glacier Bay and probably on the outer continental shelf (Worley, 1980; Mann, 1986b), the consistent trend of trees arriving later at more northwesterly sites across southeast and south-central Alaska suggests that no tree species sur- vived the LGM in local refugia. Evidently climate was too extreme during the LGM for the survival of arboreal vegetation.

Opinions vary about the timing of an interval of warmer and/or drier climate in southeast Alaska during the early Holocene. Initial reconstructions from pollen diagrams suggested a time of maximum warmth and minimum precipitation between 5000 and 2000 BP (Heusser, 1960). However, temperature and precipitation trends deciphered by transfer functions using pollen data suggest the Hypsithermal Interval occurred between 9000 and 6000 BP with a precipitation minimum occurring at 8000 BP (Heusser et al., 1985). Wetter and cooler climate definitely is registered by pollen records after ca. 3300 BP (C.J. Heusser, 1985; C.J. Heusser et al., 1985).


Neoglaciation began on the western slopes of the Fairweather Mountains ca. 6000 BP and resulted in three distinct glacial maxima 6000-5000 BP, 3500-2500 BP and 1800-1900 A.D. (Mann and Ugolini, 1985; Mann, 1986b). However, glaciers in Glacier Bay were still retracted behind their present-day positions at 6000 BP (McKenzie and Goldthwait, 1971; Goodwin, 1988). Around 5000 BP, glaciers in the West Arm of Glacier Bay began an advance which, though interrupted by still- stands and minor retreats, was to carry the terminus to the mouth of the bay by ca. 1700 A.D. (Goldthwait, 1966). Glaciers in Muir Inlet had joined this advance by 2700 BP (Goodwin, 1988; Goldthwait, 1966). After a brief retreat ca. 2000 BP, the coalescing glaciers continued to advance towards the mouth of the bay until ca. 900 BP when another stillstand or minor retreat occurred (Goodwin, 1988). Sometime after 850 BP a final advance occurred to the mouth of Glacier Bay. After ca. 1750 A.D., the glacier that filled Glacier Bay underwent a rapid retreat from its maximum Little Ice Age position (Goldthwait, 1966; Field, 1975; Clague and Evans, 1993). This retreat perhaps was initiated by climatic amelioration but was certainly exaggerated by the rapid calving of icebergs in the deep waters of this fjord.

In Glacier Bay, biota has rapidly colonized the land exposed by retreating glaciers over the last 250 years. The rates and patterns of this recolonization, especially by vegetation and by salmonid fish, may be partly analogous to prehistoric events in older deglaciated areas. Salmonids do not colonize cold and turbid streams heading in glaciers (Milner and Bailey, 1989). Even after the disappearance of glacial ice, rapid downcutting in recently-deglaciated stream systems discourage large salmon runs (Benda er al., 1992). However, channels stabilize and suspended sediment load decline following deglaciation as streams approach equilibrium profiles and large woody debris accu- mulates. As streams become clear and water temperatures rise, salmon rapidly colonize them. In Glacier Bay, Oncorrhynchus nerka (sockeye salmon) and 0. kisutch (coho salmon) established runs in Nunatak Creek less than 15 years after the watershed was deglaciated (Milner and Bailey, 1989). However, the rapid colonization of Glacier Bay streams following the Little Ice Age may not be analagous to salmon recolonization following the LGM. Glacier Bay after the Little Ice Age was a small area of newly-exposed terrain within a large region of stream systems already containing salmon. Following the rapid retreat of LGM glaciers in coastal areas, large numbers of streams were created in a region devoid of established salmon runs. We can only speculate that salmon populations took longer, perhaps centuries rather than decades, to colonize late Pleistocene streams as their spawning populations increased explosively in formerly ice-covered areas.

COASTAL BRITISH COLUMBIA AND WASHINGTON

Puget Sound in Washington marks the southern limit of a landscape unit that reaches its northern limit in the Alexander Archipelago. This region of glacially-scoured

fjordlands and associated strand flats is bordered to the west by the Pacific Ocean and to the east by the Coast Mountains and Cascade Range (Fig. 10). These near- coastal mountains support numerous glaciers today and during the Pleistocene comprised the heartland of the Cordilleran Ice Sheet and its ancillary fringe of alpine ice caps and glaciers.

The Cordilleran Ice Sheet advanced twice in southwest British Columbia during the interval 25,000- 10,000 BP (Clague, 1991; Ryder et al., 1991; Easterbrook, 1992) (Fig. 11). The first glacier maximum was reached during the Evans Creek stade (Coquitlam stade in British Columbia) between ca. 22,000 and 19,000 BP probably in response to a short-lived (< 1000

FIG 0. Western British Columbia and western Washington.


I existing icefields I \

FIG. 11. The changing extent of the Cordilleran Ice Sheet over southwest British Columbia and Washington after 25.000 BP (after Clague, 1981). Alpine glaciers in the Cascade Range and in the Olympic Mountains are not shown.

years) but severe drop in regional snowline to as low as 1000 m above sea level (Booth, 1986, 1987). This fall in snow line probably accompanied a temperature depression > 6°C (Porter et al., 1983). Ice failed to reach the Pacific coast on western Vancouver Island and left much of Puget Sound ice-free during the Evans Creek stade (Ryder and Clague, 1989). The second, more extensive advance, culminated ca. 14,500-14,000 BP in southern Puget Sound during the Vashon stade (Thorson. 1989; Ryder and Clague, 1989; Easterbrook. 1992). During the Vashon stade, snowline depression was less than during the Evans Creek stade but lasted longer, probably for several thousand years (Booth, 1987, 199 I ). Lowered Vashon stade snowlines probably were associated with increased precipitation (Mathewes, 199 1) accompanying increased storminess along a southward-displaced polar jet stream (Bamosky et al., 1987). Glaciers reached the open Pacific during the Vashon stade along the west coast of Vancouver Island (Anderson, 1968; Alley and Chatwin, 1979) and buried the inner fjords and straits under thick ice (Ryder and Clague, 1989). The Evans Creek and Vashon advances were separated by the Port

Moody interstade (Hicock and Armstrong, 1985) when ice retreated into the inner fjords and mountain valleys of the Coast Mountains (Clague et al., 1988).

Retreat of ice was rapid after 14,000 BP and involved iceberg calving of glaciers into the deep water of the isostatically depressed Puget Sound and Straits of Georgia (Ryder and Clague, 1989; Easterbrook, 1992). The subsequent Sumas re-advance between 11,500 and 11,000 BP was followed by the final collapse of the southwestern portion of the Cordilleran Ice Sheet before 10,000 BP. Fjords bordering the Straits of Georgia were ice-free by 11,000 BP (Ryder and Clague, 1989).

Farther north, outlet glaciers draining the Coast Mountains between Vancouver Island and the Queen Charlotte Islands terminated in calving margins along the outer edge of the continental shelf during the LGM (Blaise et al., 1990; Josenhans, 1992). Cordilleran ice advanced onto the eastern and northern shores of the Queen Charlotte Islands between 23,000 and 21,000 BP (Clague et al., 1982; Blaise et al., 1990). Mainland- derived glaciers retreated from the Queen Charlottes before 15,000 BP following a maximum advance that

D.H. Mann and T.D. Hamilton: Paleoenvironments of the North Pacific Coast

127”W 125’W

0 2000 4000 6000 8000 1

Calibrated Radiocarbon Years B.P.

463

51”N

47-N

FIG. 12. Holocene sea-level histories for three areas of coastal, southwestern British Columbia. Datum is mean tide level. Redrawn from Hutchinson (1992).

culminated before 16,000-15,500 BP (Clague et al., 1982; Warner et al., 1982; Blaise et al., 1990). By 13,000 BP, ice had retreated to the inner fjords of the mainland (Ryder and Clague, 1989).

Postglacial sea-level changes along the British Columbia and Washington coasts varied in a complex fashion according to ice-loading history, eustatic sea-level history, and local tectonism (Hutchinson, 1992). In areas bordering the eastern shores of the Strait of Georgia that were deeply buried by Late Wisconsin ice, sea level fell between the time of deglaciation and the early Holocene as isostatic rebound uplifted land formerly covered by hundreds of meters of ice and outstripped the rise in eustatic sea level. Isostatic rebound following ice unloading was the major determinant of post-glacial sea level in these areas. Post-glacial marine limits were reached before 11,000 BP at altitudes as high as 200 m above present sea level in fjords along the inner coastline of British Columbia (Clague, 1981). Sea level stabilized between 11,000 and 6ooo BP (Clague, 1989; Hutchinson, 1992) as isostatic rebound slowed, matching the rate of eustatic rise.

In areas close to the margin of the Cordilleran Ice Sheet, post-glacial marine limits also lie tens to hundreds of meters above modem sea level (Clague et al., 1982; Thorson, 1989). However, isostatic rebound ended earlier than in the inner fjords allowing a marine transgression of 5-10 m between ca. 8000 and 3000 BP caused by eustatic sea-level rise (Eronen et al., 1987; Hutchinson, 1992; Dragovich et al., 1994). Slowing of eustatic sea- level rise after ca. 5000 BP allowed expression of tectonic effects on relative sea level in western Washington (Bucknam et al., 1992; Hutchinson, 1992). Areas south of the Cordilleran Ice Sheet in Washington and Oregon that were beyond the effects of ice loading experienced monotonic submergence overprinted by tectonic movements (Atwater, 1992) as eustatic sea level rose following a late Pleistocene lowstand (Hutchinson, 1992).

The west coasts of Vancouver Island and northwest Washington experienced a distinctive sea-level history strongly influenced by tectonism (Hutchinson, 1992). Post-glacial marine limits were relatively low, 20-25 m along the outer coast of Vancouver Island (Clague, 1989; Friele and Hutchinson, 1993). In response to isostatic rebound, sea level fell in the early Holocene to a lowstand varying from +5 to -10 m relative to present sea level. After isostatic recovery from ice loading was com- pleted lO,OOO-8000 BP, eustatic sea level rose, causing a marine transgression that culminated in the middle to late Holocene, 2-5 m above present sea level. As eustatic sea level equilibrated in the middle Holocene, tectonic uplift became the dominant control over sea level causing a fall in relative sea level over the last 4000 years (Friele and Hutchinson, 1993).

More is known about Pleistocene sea-level history from the Queen Charlotte region than from any other sector of the coast. Sea level was below that of the present day from at least 15,000 to ca. 10,000 BP (Clague, 1989). Around 10,600 BP, sea level stood near -100 m and exposed large areas of the floor of Queen Charlotte Sound (Lutemauer et al., 1989). These areas may have remained subaerial until ca. 10,000 BP. The Holocene sea-level history of the Queen Charlotte Islands resem- bles that of western Vancouver Island. Modem sea level was probably reached by 9000 BP after a rapid transgression (Josenhans, 1992). Sea level then rose more gradually in a transgression culminating between 8500 and 7500 BP with levels as much as 15 m above the modem, perhaps in response to the eastward flow of a forebulge in the underlying mantle (Claglie, 1983). Falling sea level on the Queen Charlotte Islands since ca. 7500 BP may be due to tectonic uplift (Clague, 1989). Certainly the active tectonic setting of the British Columbia coast makes large vertical-crustal movements likely (von Huene, 1989; Rogers and Homer, 199 1).


Pollen evidence suggests that climate in southern British Columbia was colder and drier than today between 25,000 and 16,000 BP, although several major climatic oscillations occurred within this interval. During the Evans Creek stade (ca. 22,000-19,000 BP), mean annual temperature was 5-8°C lower and annual precipitation 700-1000 mm less than today (Hicock et al., 1982; Whitlock, 1992; Thompson et al., 1993). At this time, subalpine parkland and possibly alpine plant communities covered the lowlands of eastern Vancouver Island (Clague and MacDonald, 1989). Pollen, beetle remains, and plant macrofossils suggest a relatively warm and moist interval between 19,000 and 18,000 BP during the Port Moody interstade (Hicock et al., 1982; Mathewes, 1991). At this time, surprisingly diverse vegetation exis- ted in the coastal areas of southwestern British Columbia that later were overwhelmed by Vashon stade ice. Near the city of Vancouver, subalpine fir-Engelmann spruce forest and parkland grew under a cold-humid continental climate (Hicock et al., 1983; Clague and MacDonald, 1989).

than today (Pellatt and Mathewes, 1994). The warmer and drier climates of this Hypsithetmal interval gradually were replaced by cooler and wetter conditions that accompanied the expansion of western hemlock and western red cedar in the forests of southwestern British Columbia (Mathewes, 1985).

Neoglaciation began in the Canadian Cordillera as early as 5000 BP (Luckman et al., 1993). Widespread glacier expansion was underway by 3500 BP (Ryder, 1989; Luckman et al., 1993) to maxima that were possibly time transgressive, occurring between ca. 3300 and 1900 BP (Denton and Karlen, 1977; Ryder and Thompson, 1986; Osborn and Luckman, 1986, 1988; Clague and Matthews, 1992). Maximum Holocene extent of many glaciers occurred during the Little Ice Age. Little Ice Age advances in the Coast Mountains and Cascade Range began ca. 1350 A.D. and ended ca. 1900 A.D. with the onset of widespread glacier retreat (Burbank, 1981; Heikkinen, 1984; Ryder, 1989).

SYNTHESIS Climate became warmer and wetter in coastal British

Columbia after 16,000 BP (Nelson and Coope, 1982; Mathewes, 1991), although pollen data suggest it was still cooler and drier than today as late as 10,000 BP (Mathewes, 1985). At ca. 15,000 BP, non-arboreal herb- and shrub-tundra covered the lowlands of the Queen Charlotte Islands (Warner et al., 1982). At 14,000 BP, ice-free areas on northwest Vancouver Island supported heaths, grass-sedge-herb meadows, and spatially-restric- ted conifer vegetation (Hebda, 1992). Significant warming occurred between 12,000 and 10,000 BP when conifers rapidly invaded deglaciated areas (Sugita and Tsukada, 1982). Southwestern coastal British Columbia was recolonized by plants ca. 13,500 BP with trees arriving ca. 13,000 BP (Clague and MacDonald, 1989). Lodgepole pine and poplar probably were the first tree colonists on the Queen Charlotte Islands, arriving about 12,000 BP, followed by spruce immigration ca. 11,200 BP (Warner et al., 1982). Summer temperatures in shallow ponds on the Olympic Peninsula may have reached modern values at 12,000 BP (Petersen et al., 1983). Several pollen diagrams suggest a reversion to cold conditions between 11,500 and 10,000 BP at the time of the Younger Dryas episode in northwest Europe (Mathewes et al., 1993).

In rough synchrony with the worldwide LGM, the North Pacific region saw a maximum extent of glaciers between 22,000 and ca. 17,000 radiocarbon years BP. Due to regional patterns of precipitation and topogra- phy, the northeastern (North American) margin of the Pacific Ocean was more heavily glaciated than the northwestern (Asian) margin. Subsequent deglaciation was rapid in coastal areas where glaciers that were grounded below rising sea levels collapsed rapidly through iceberg calving. The continental shelf of south- em Alaska near Kodiak Island (Mann and Peteet, 1994) and lower Cook Inlet (Reger and Pinney, in press) were deglaciated before ca. 14,700 and 16,000 BP respective- ly. The timing of deglaciation was similar on the Queen Charlotte Islands where ice originating on mainland British Columbia had withdrawn before 15,000 BP (Blaise et al., 1990). By inference, the intervening coastline of south-central and southeast Alaska may also have been deglaciated by this date, at least on the continental shelf. Inner fjords in places like Prince William Sound and the easternmost parts of the Alexander Archipelago may have retained glaciers at least slightly more extensive than today until ca. 10,000 BP.

Rapid warming occurred at ca. 10,000 BP in south- The LGM was accompanied by a ca. 120 m drop in western British Colombia. Forests dominated by Douglas global (eustatic) sea level which exposed large portions fir and alder were established during an interval that was of the unglaciated continental shelves in the Bering- drier and as warm or warmer during the summer than Chukchi Seas and northeast Asia. Following degiacia- today and lasted until ca. 6000 BP (L.E. Heusser, 1983; tion, eustatic sea level rose until ca. 4ooo BP (Fairbanks, C.J. Heusser, 1985). Climatic conditions at 9000 BP were 1989). Generalized curves of eustatic sea-level change the driest of the Holocene with pronounced summer are not applicable to many areas around the North Pacific droughtiness (Thompson et al., 1993). Treeline in the due to strong interference from post-glacial tectonism southeastern Coast Mountains was between 60 and and/or glacio-isostatic rebound. As a result, there is no 130 m higher than today during the interval 9100 to 8200 universal sea-level curve for the North Pacific region. BP (Clague and Mathewes, 1989). Western hemlock Within formerly-glaciated areas and in areas of ongoing (Tsuga heterophylla [Raf.] Sarg. trees grew at higher than tectonism, detailed reconstructions of sea-level history modern elevations in the Queen Charlotte Islands are necessary at a local (ca. 1000 km2) scale. An excep- between 9600 and 8700 BP, suggesting warmer climate tion to this caution may be the Bering and Chukchi


basins where tectonism and glacial isostasy have been relatively minor over the last 15,000 years as evidenced by the similar altitudes (8-12 m) of the Isotope Stage 5e, Pelukian shoreline on the Alaskan coast of the Chukchi Sea (McCulloch, 1967; Brigham, 1985; Brigham-Grette and Hopkins, 1995). Comparisons between the Fairbanks (1989) sea-level record and the bathymetry of the Bering and Chukchi Seas suggest that the Bering land bridge probably was not flooded until ca. 10,000 BP, a timing substantiated by radiocarbon dates on submerged terrestrial peats (Elias et al., 1992).

Pollen and plant macrofossils provide an important source of paleoclimatic data in the North Pacific region. During the onset of the last glacial maximum, forests retreated southward in northeast Asia, but were extirpat- ed in Alaska and the Yukon where their retreat was blocked by glaciated mountains and the ocean. During the LGM, northern Japan and the Asian mainland north of the Amur River was a mostly treeless tundra (Grichuk, 1984; Lozhkin et al., 1993). Beringia probably supported a steppe-tundra vegetation lacking any clear modern analogs but capable of supporting a diverse large-mam- mal fauna of unknown population size (Guthrie, 1990). Following the LGM, important changes in vegetation occurred throughout the region in response to the jerky shift of the global ocean-climate system from glacial into interglacial mode and the retreat of glaciers. With the retreat of glaciers from the northeast Pacific coastline, biota began to move northward from refugia in Washington. Similar northward shifts must have occurred along the Russian coast during the Holocene. Forest vegetation similar to today’s was established by 4000 BP in the northern Alexander Archipelago (Peteet, 1986; Cwynar, 1990) and by 5000 BP in northwest Alaska along the Chukchi Sea coast (Anderson, 1985, 1988). Plant colonization of terrain deglaciated more than 10,000 years ago continues at present in south-central and southwest Alaska (Peteet, 1991; C.J. Heusser, 1985).

Despite several decades of discussion, no indisputable evidence exists for the persistence of glacial refugia along the seaward margin of the Cordilleran Ice Sheet in Canada or Alaska (C.J. Heusser, 1989). Ice-free areas did persist near Lituya Bay (Mann, 1986), in the Caribou Hills east of lower Cook Inlet (Reger and Pinney, in press), and on southwest Kodiak Island (Karlstrom and Ball, 1969; Mann and Petett, 1994), but they probably contained depauperate, tundra communities of little importance in the recolonization of the forested, postglacial landscape.

Early Holocene climates were warmer and drier than today in most regions around the North Pacific. As a con- sequence, treeline was higher and glaciers were in retracted positions between ca. 9500 and 7000 BP in the Coast Mountains of British Columbia (Clague et al., 1992). Interior areas seem to have warmed early in postglacial times in response to a peak in high-latitude solar radiation centered on ca. 10,000 BP (Barnosky et al., 1987). In coastal areas in western Alaska and Chukotka, as well as coastlines in the Gulf of Anadyr and the Sea of Okhotsk, flooding of nearby continental shelves during

the post-glacial marine transgression may have prevented drier and warmer conditions during the early Holocene (Lozhkin et al., 1993). Glaciers remained in retracted positions throughout the coastal mountains rimming the Gulf of Alaska until after 6000 BP

Neoglaciation began in the American Cordillera in response to cooler and probably more moist conditions after the middle Holocene (Calkin, 1988; Wiles and Calkin, 1993; Ryder, 1993). Neoglaciation probably accompanied regionally synchronous shifts in temperature and precipitation, however these events were of smaller magnitude than during the LGM to Holocene transition and consequently are recorded with varying sensitivity by pollen and glacier records. The history of the Little Ice Age, which spanned roughly the last millen- nium, is recorded by both the historical record and by glacier and pollen proxy records.

ACKNOWLEDGEMENTS

We thank Dorothy Peteet, Paul Carlson, John Clague, Julie Brigham-Grette, and Dick Reger for constructive comments on earlier drafts of this review. Owen Mason and Patricia Heiser provided useful discussions and references.

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LATE PLEISTOCENE AND HOLOCENE PALEOENVIRONMENTS OF … · 2015-09-04 · (Meltzer, 1993). Biotic interchange between continents is one reason why the paleoclimates and paleogeography