- 1 Paleoclimates, Paleovegetation, and Paleofloras of North America
- 2 North of Mexico During the Late Quaternary
- 2.1 Climatic Change
- 2.2 Biotic Changes during the Last Glacial-Interglacial Cycle
Paleoclimates, Paleovegetation, and Paleofloras of North America
North of Mexico During the Late Quaternary
Paul A. Delcourt
Hazel R. Delcourt
The origin of extant vascular plant genera and of many species occurred largely prior to the Quaternary. The Quaternary is characterized more by changes in the distributions of plant taxa and in the organization of plant communities (C.W. Barnosky 1987; L.E. Heusser and J.E. King 1988) than by the evolution of genera and species. One important event of the late Quaternary that affected subsequent biological diversity was the series of megafaunal extinctions that coincided with the spread of PaleoIndians between 12,000 and 10,000 years Before Present (yr B.P.) (P.S. Martin and R.G. Klein 1984). Prehistoric Amerinds influenced the distributions of certain plant species as well as the composition of plant communities, particularly during the last 5000 years, during which aboriginal populations became sedentary and grew native and introduced plants for food (R.J. Hebda and R.W. Mathewes 1984; P.A. Delcourt et al. 1986; H.R. Delcourt 1987; J.H. McAndrews 1988).
In this chapter, we summarize current views regarding the effects of climatic change on the vegetation and flora of North America during the Quaternary. We focus primarily on the past 20,000 years. This time interval represents late Pleistocene full-glacial conditions (20,000--15,000 yr B.P.), late-glacial climatic amelioration from 15,000 to 10,000 yr B.P., and Holocene interglacial conditions of the last 10,000 years. We review the broad patterns and timing of climatic, vegetational, and floristic changes that have occurred in North America and Greenland during this last glacial-interglacial cycle.
The closure of the Isthmus of Panama by tectonic activity in the late Pliocene epoch 3 million years ago (Ma) produced a continuous land bridge between North America and South America, prevented exchange of equatorial water between the Pacific and Atlantic oceans, and resulted in the strengthening of the Gulf Stream (W.A. Berggren and C.D. Hollister 1974). Moisture from storms that tracked northward along the warm Gulf Stream fed the accumulation of glacial ice in Greenland and northeastern Canada. Glacial tills in the midcontinental region of North America, dated from 2.4 to 2.2 Ma, are direct evidence of the expansion of major ice sheets (J.Boellstorff 1978).
Between 3 and 2.5 Ma, mountain glaciers developed at high latitudes of the Northern Hemisphere. During the last 2 million years, they expanded in suitable high elevation sites at middle latitudes concurrently with the uplift of the western Cordillera, particularly along the western Coast Ranges and Sierra Nevada (P.W. Birkeland et al. 1976; J.J. Clague 1989).
Because of these changes in configuration of continents and oceans, the Quaternary climates at middle and high latitudes became increasingly sensitive to seasonal variations in incoming solar radiation, and the onset of glacial-interglacial cycles began. Major papers by G.H. Denton and T.J. Hughes (1983) and by W.S. Broecker and G.H. Denton (1989) summarize the role of long-term changes in solar radiation in both initiating and amplifying climatic interactions among glaciers, oceans, and the atmosphere.
J.Chappell (1978) provided a balanced summary of the many published theories that attempt to account for cyclic climatic changes and repeated episodes of continental glaciation during the Quaternary. Milutin Milankovitch [Milankovic](1941) hypothesized that variations in the earth's orbit around the sun result in cyclic variation in the amount of solar radiation received by the earth, producing a "solar pacemaker" (J.D. Hays et al. 1976), which ultimately led to the climatic changes associated with the glacial-interglacial cycles (A.L. Berger 1978). Over each 100,000-year Milankovitch cycle, the amount of solar radiation received by the earth varies by 3.5%; the variation is attributed to changes in the eccentricity of the earth's elliptical path around the sun.
Two other Milankovitch cycles influence the seasonal contrast in radiation received at different latitudes. A 41,000-year cycle results from changes in the tilt of the earth's rotational axis, which in turn affects the seasonal contrasts between summer and winter. A 21,000-year cycle results from the precession of the equinoxes, in which the earth's axis of rotation moves slowly along a circular path. High latitudes above 65° N and 65° S, where continental ice sheets form and decay, are primarily affected by the 41,000-year cycle. The 21,000-year cycle amplifies climatic variations between latitudes 65° N and 65° S. Together, these two cycles are responsible for the glacial advances and retreats that are superimposed upon the 100,000-year glacial-interglacial cycles (J.Imbrie and K.P. Imbrie 1979).
Isotope curves developed from measurements of the ratio of 18O to 16O in the fossil shells of marine plankton reflect long-term oscillations in the volume of glacial ice and indicate that the 100,000-year Milankovitch cycle has become dominant in driving climatic change during the last 900,000 years (N.G. Pisias and T.C. Moore Jr. 1981). Over the course of each 100,000-year cycle (fig. 4.1A), glacial ice builds up slowly on the continents, with the glacial mode typically lasting for about 90,000 years (W.S. Broecker and J.van Donk 1970; W.A. Watts 1988). Glacial conditions then terminate rapidly as glaciers melt in response to increased solar radiation and maximum seasonal contrast in middle latitudes, conditions associated with the combined peaks of both the 21,000-year and the 100,000-year Milankovitch cycles (A.L. Berger 1978; J.Imbrie and K.P. Imbrie 1979).
The rapid termination of each glacial interval is triggered when the time of highest solar radiation coincides with the maximum seasonal range in solar radiation (fig.4.1B; A.L. Berger 1978; J.E. Kutzbach 1987; W.A. Watts 1988). Increased summer temperatures melt glacial ice on land. Glacial meltwater returns to the oceans, forming a freshwater cap over denser, saline ocean water. Colder winter temperatures freeze upper marine waters, creating a layer of pack ice that in turn cuts off the supply of moisture evaporating from oceans to the atmosphere, further reducing the snow supply to disintegrating continental glaciers (W.F. Ruddiman and A.McIntyre 1981). The subsequent interglacial mode lasts for only 10,000 to 15,000 years. The overall shape of the glacial-interglacial curve of global ice volume (fig. 4.1B), as determined by 18O/16O ratios (d18O), is asymmetrical and sawtoothed, not smoothly sinusoidal as are the fluctuations of solar radiation attributed to individual Milankovitch cycles (W.S. Broecker and J.van Donk 1970).
The Full-Glacial Climate
Proxy evidence of climatic change recorded on land
(S.C. Porter 1983; H.E. Wright Jr. 1983; R. B. Morrison 1991) and in the oceans (CLIMAP Project 1981; W.F. Ruddiman and H.E. Wright Jr. 1987) provides constraints to guide computer simulations of full-glacial climate 18,000 yr B.P. (J.E. Kutzbach 1987; J.E. Kutzbach and H.E. Wright Jr. 1985; J.E. Kutzbach and P.J. Guetter 1986; COHMAP 1988; P.J. Guetter and J.E. Kutzbach 1990). This section is a synthesis of the results from computer simulations and available paleoclimatic proxy data.
During full-glacial conditions 18,000 years ago, three complexes of continental ice sheets dominated the northern half of North America, Greenland, and northern Europe (delineated by stippled areas of fig. 4.2). The North American complex included the western Cordilleran and eastern Laurentide ice sheets. The accumulation of water on land as glacial ice resulted in sea levels of full-glacial oceans being 100--120 m below modern levels. Lowered sea levels exposed extensive areas of coastal plains, particularly in Beringia (the land bridge that connected eastern Siberia and Alaska [D. M. Hopkins et al. 1982]), southern North America, southern Europe, and North Africa (J.E. Kutzbach and H.E. Wright Jr. 1985). The diagonal pattern on figure 4.2 identifies the winter extent of sea pack ice north of Beringia in the Arctic Ocean, and south to about 42° N in the North Atlantic Ocean (CLIMAP Project 1981).
Between latitudes 30° N and 60° N, fossil data and computer simulations indicate that full-glacial winters on land were an average of 6° C colder, and summers about 2° C cooler than today. In winter the polar jet stream split in the North Pacific Ocean (fig. 4.2). A northern branch of this polar jet stream passed along the northern perimeter of the Cordilleran and Laurentide ice sheets in northern Canada. The southern branch crossed the American Southwest at about latitude 30° N and flowed over the Southern High Plains and the middle latitudes of unglaciated southeastern North America. The southern jet stream would have funneled storms along this route and would have caused substantially cooler temperatures and increased winter precipitation across the Southwest and from about latitude 33° N to the glacial margin (about latitude 40° N) in the Southeast (fig.4.2; J.E. Kutzbach and H.E. Wright Jr. 1985; J.E. Kutzbach and P.J. Guetter 1986).
In western North America, this "pluvial" climatic regime led to a 10-meter rise in the groundwater table and the expansion of large, permanent pluvial lakes in which the water level rose as much as 335 m above modern levels (G.I. Smith and F.A. Street-Perrott 1983; L.Benson and R.S. Thompson 1987). Computer simulations of full-glacial climate in North America also specify an anticyclonic circulation of winds radiating outward from the ice sheets that would have brought cold, dry climatic conditions to the unglaciated area of the Pacific Northwest, and cool but very dry conditions to Beringia (fig. 4.2; J.E. Kutzbach and H.E. Wright Jr. 1985).
The Late-Glacial and Holocene Intervals
Computer simulations of climatic conditions (J.E. Kutzbach and P.J. Guetter 1986) show that during the portion of the late-glacial interval from 15,000 to 12,000 yr B.P., the two full-glacial branches of the jet stream (fig. 4.2) shifted in location and merged along one route as climate moderated (J.E. Kutzbach 1987). Between 12,000 and 9000 yr B.P., the route of this polar jet stream extended from latitude 52° N over the western mountain ranges of British Columbia and tracked southeastward across the midcontinent to about latitude 45° N along the southern Laurentide glacial margin. In the last 6000 years, the position of the jet stream has shifted seasonally in location, in winter occupying the midlatitudinal route from British Columbia to the northeastern United States (comparable with the path from 12,000 to 9000 yr B.P.), and in summer occupying a northerly route across Alaska and northern Canada from about 65° N to 70° N.
Changing environmental conditions through the past 18,000 years include a global decrease in extent of glacial ice, a postglacial increase of about 4° C in the global mean values for annual sea-surface temperature (SST), and seasonal changes in solar radiation received by the earth during summer and winter (fig. 4.3). Reconstructions based on Milankovitch cycle calculations indicate that, for the Northern Hemisphere, in the interval between 12,000 and 9000 yr B.P., there was maximum seasonal contrast in temperatures, with warmer summers and cooler winters (A.L. Berger 1978; J.E. Kutzbach 1987; P.J. Bartlein 1988). The Pleistocene-Holocene boundary at 10,000 yr B.P. consequently was a time of fundamental changeover in environmental and climatic conditions from a glacial to an interglacial mode (fig. 4.3). Holocene interglacial climates have been characterized by relatively warm conditions, by decreasing amounts of incoming solar radiation during summer, and by increasing solar radiation levels during the winter season, with an overall decrease in seasonal contrast (fig. 4.3; J.E. Kutzbach 1987).
Biotic Changes during the Last Glacial-Interglacial Cycle
Eastern North America
Various authors have synthesized the late Quaternary literature of plant fossil data for the eastern United States and Canada (J.C. Bernabo and T.WebbIII 1977; P.J.H. Richard 1977; H.E. Wright Jr. 1981; M.B. Davis 1976, 1981, 1983; W.A. Watts 1983; V.M. Bryant Jr. and R.G. Holloway 1985; R.B. Davis and G.L. Jacobson Jr. 1985; P.A. Delcourt and H.R. Delcourt 1987, 1987b; G.L. Jacobson Jr. et al. 1987; J.C. Ritchie 1987; T.WebbIII et al. 1987; T.WebbIII 1988; J.V. Matthews Jr. et al. 1989; H.R. Delcourt and P.A. Delcourt 1991), and for northeastern Canada and Greenland (B.Fredskild 1973, 1985; P.J.H. Richard 1977; J.T. Andrews 1985; H.F. Lamb and M.E. Edwards 1988; J.V. Matthews Jr. et al. 1989). This regional review, based substantially on examples documented in P.A. Delcourt and H.R. Delcourt (1987, 1987b) and in H.R. Delcourt and P.A. Delcourt (1991), illustrates the response of temperate and boreal plant species and communities to climatic and environmental changes over the last glacial-interglacial cycle in eastern North America (from 25° N to 60° N and 50° W to 100° W).
During the peak of the last continental glaciation, between 20,000 and 18,000 yr B.P., the Laurentide Ice Sheet extended south to about latitude 40° N across the Great Lakes region (fig. 4.4). The Laurentide Ice Sheet influenced the full-glacial climate, locking the frigid Arctic air mass north of the great ice dome. The prevailing westerlies of the Pacific air mass thus dominated the region south of the glacial ice to 33° N. The southern branch of the full-glacial jet stream (fig. 4.2) extended across the middle latitudes as the Polar Frontal Zone that marked the climatic boundary between prevalence of the Pacific air mass and the Maritime Tropical air mass to the south (fig. 4.4). In a narrow, periglacial zone along the southern and eastern margin of the ice sheet, tundra communities were present only in localized areas of severe permafrost environments. These occurred in glacial reentrants such as in southeastern Iowa (R.G. Baker et al. 1986), at middle to high elevations within the central Appalachian Mountains, and on the coastal plain of the Northeast.
Boreal forest dominated in the belt south of the Laurentide ice to approximately latitude 34° N, extending westward across the Great Plains. Elements of deciduous forest in the southern region of the full-glacial boreal forest indicate that the boreal climate was less extreme than that of today (P.A. Delcourt et al. 1980). Based on quantitative comparison of fossil and modern pollen assemblages, midwestern plant communities dated from 28,000 until 15,000 yr B.P. were very similar in both forest composition and structure with their modern counterparts in the boreal forest biome (R.G. Baker et al. 1989). Prairie elements may have occurred in the understory of this boreal forest, but extensive prairie tracts were not present (H.E. Wright Jr. 1981). Between latitudes 34° N and 33° N, a narrow transition zone marked the ecotone between more northern boreal and more southern temperate communities.
South of latitude 33° N, across the southern Atlantic and Gulf coastal plains, floristic elements of temperate deciduous forest occurred with plant taxa characteristic today of southeastern evergreen forests (fig. 4.4; P.A. Delcourt and H.R. Delcourt 1987). In the Southeast, average temperatures during the full-glacial interval were probably similar to modern values; late Pleistocene evidence documents that sea-surface temperatures 18,000 yr B.P. in the northern Gulf of Mexico were less than 2° C cooler than those of today (CLIMAP Project 1981). This slight oceanic cooling decreased summer rates of evaporation from marine waters and substantially reduced the precipitation supplied by tropical storms and hurricanes across the interior of the Southeast (W.M. Wendland 1977; P.A. Delcourt and H.R. Delcourt 1984).
Sand dune scrub occupied the broad, exposed coastal plain of central and southern Florida. Tropical plant species were eliminated from southernmost Florida. Dry climate and a decline in water table by up to 20 m, resulting from lowered sea level (W.A. Watts 1983), may have been responsible for these xeric vegetation patterns across peninsular Florida.
In general, major vegetation patterns remained relatively consistent from 20,000 to 15,000 yr B.P. (P.A. Delcourt and H.R. Delcourt 1984). Between 15,000 and 10,000 yr B.P., late-glacial climatic warming and the northward retreat of the ice sheet resulted in widespread changes in climate and vegetation. Vegetation regions and biomes today correlate with distinctive climatic regions, delineated by mean positions of air mass boundaries (R.A. Bryson 1966; R.A. Bryson and F.K. Hare 1974b). This principle can be extended back in time to the analysis of former vegetation patterns (fig. 4.3; R.A. Bryson and W.M. Wendland 1967; P.A. Delcourt and H.R. Delcourt 1983, 1984). By 10,000 yr B.P., boreal forest spread throughout the deglaciated region immediately south of the Laurentide Ice Sheet, through the boreal climate zone where the Arctic air mass dominated in winter and the Pacific air mass was important in the summer. Portions of Nova Scotia, Newfoundland, and Labrador, however, were occupied by tundra (J.V. Matthews Jr. et al. 1989).
Across the Great Lakes and New England regions, the latitudinal belt of mixed conifer--northern hardwoods forest corresponded in distribution with the zone influenced by the average location of the polar jet stream, that is, the broad Polar Frontal Zone representing the prominent climatic boundary between boreal and temperate climatic regions. Deciduous forest spread northward with the areal expansion of the Maritime Tropical air mass. The southeastern evergreen forest remained confined to the Gulf and Atlantic coastal plains and in Florida it replaced the sand dune scrub. Prairie vegetation had developed by 10,000 yr B.P. in the central and southern Great Plains, where the relatively dry Pacific air mass dominated throughout the year.
By 6000 yr B.P. (fig. 4.4), remnants of continental ice sheets had largely retreated north of latitude 60° N. Tundra communities were widespread north of latitude 55° N both to the east and to the west of Hudson Bay, within the climatic region dominated exclusively by the Arctic air mass. Boreal forest occupied the zone between latitudes 50°N and 55° N, where Arctic and Pacific air masses prevailed in winter and summer seasons, respectively. The Polar Frontal Zone and mixed conifer--northern hardwoods forest occupied a wide band across the Great Lakes and New England regions.
A wedge of the warm, dry Pacific air mass and prairie vegetation extended eastward into Illinois and beyond as the prairie peninsula reached its easternmost extent during the middle Holocene (the Hypsithermal Interval, generally dated between 9000 and 4000 yr B.P. [H.E. Wright Jr. 1976]). With further northward influence of the warm, moist Maritime Tropical air mass, deciduous forest communities became widespread across the eastern United States, and southeastern evergreen forests spread northward along the Atlantic Coastal Plain and the Mississippi Alluvial Valley (P.D. Royall et al. 1991). With the rise in sea level and the increase in sea-surface temperatures to modern levels by 5000 yr B.P., the southern portion of the Florida peninsula was invaded by subtropical plant species.
Vegetation patterns for presettlement times, 500 yr B.P. (fig. 4.4), differ from those of the middle Holocene interval. During the last 3000 years, global climatic cooling resulted in a southward displacement of the ecotone between tundra and boreal forest across northeastern Canada (S.Payette and R.Gagnon 1985; S.Payette et al. 1989; J.V. Matthews Jr. et al. 1989). Similarly, the ecotone between boreal and mixed conifer--northern hardwoods forest was displaced southward across the Great Lakes region and New England (J.V. Matthews Jr. et al. 1989). A reduction in seasonal contrast in temperature within the late Holocene (fig. 4.3) diminished the frequency and intensity of summer droughts, reduced the occurrence of wildfire, favored invasion of grasslands by trees, and resulted in westward retraction of the prairie peninsula and its replacement by eastern forests from Minnesota to Illinois (fig. 4.4; E.C. Grimm 1983; G.L. Jacobson Jr. et al. 1987).
Changes in vegetation resulting from late Quaternary climatic changes can be mapped along boundaries between biomes that coincide with the thresholds of physiological tolerance of many dominants of the biomes (B.F. Chabot and H.A. Mooney 1985). It is widely recognized, however, that the responses of plant taxa to climatic change are individualistic (M.B. Davis 1981, 1983; T.WebbIII 1988; P.A. Delcourt and H.R. Delcourt 1987, 1987b; H.R. Delcourt and P.A. Delcourt 1991). Figure 4.5 illustrates distribution patterns for the last 18,000 years for three major eastern North American tree taxa: spruce (Picea), oak (Quercus), and both the northern (N) and southern (S) groups of pine (Pinus). Procedures for producing these paleo-population maps (contoured maps with inferred population abundance are expressed as percent forest composition) are given in P.A. Delcourt and H.R. Delcourt (1987, 1987b). Differences among these taxa are evident in the mean rates of migration, in the migration routes taken northward following glacial retreat, and in the changing locations of major population centers through time.
Spruces---including white spruce (Picea glauca), black spruce (P. mariana), and red spruce (P. rubens)--- were dominant within boreal forests 18,000 yr B.P. (fig. 4.5). A full-glacial population center was located west of the Appalachian Mountains. As the Laurentide Ice Sheet retreated (P.F. Karrow and P.E. Calkin 1985; P.F. Karrow and S.Occhietti 1989), spruce developed its highest populations at the margins of proglacial lakes, i.e., lakes formed by ice-melt near the southern margin of the Laurentide glacier. By 10,000 yr B.P., outlying populations were persisting south of the main range of spruce at middle to high elevations in the central and southern Appalachian Mountains. Between 10,000 and 6000 yr B.P., the primary population centers of spruce shifted north to between 47° and 55° N. It reached its northernmost limit by 4000 yr B.P. as a major component of the boreal forest through central and eastern Canada (fig. 4.4). In the late Holocene, spruce populations generally shifted southward. Overall, the average late Quaternary rate for migration of spruce populations was 14.1 km per century (P.A. Delcourt and H.R. Delcourt 1987).
Eastern oak species were an important component of the eastern deciduous and southeastern evergreen forests south of 33° N latitude during full-glacial times (figs. 4.4, 4.5). Minor populations occurred northward within the boreal forest. By the early Holocene, oak species had invaded throughout the boreal forest, advanced to the tundra and glacial ice limits, and migrated as far north as latitude 51° N in central Canada. As early as 10,000 yr B.P., oak species were dominant or subdominant in southeastern evergreen forest as well as eastern deciduous forest that had expanded areally north to latitude 42° N. Overall, the mean rate of late Quaternary migration of oak populations was 12.6 km per century (P.A. Delcourt and H.R. Delcourt 1987).
Eastern North American pines are differentiated into two groups (W.B. Critchfield 1984; D.I. Axelrod 1986b) that today are either distributed primarily in the Appalachian Mountains and northward, or are generally restricted to the southeastern coastal plain. During the full-glacial interval, northern pines (including jack pine [Pinus banksiana], red pine [P. resinosa], and eastern white pine [P. strobus]) were dominant north of 34° N, within boreal forests located in the rain shadow east of the Appalachian Mountains on the central Atlantic Coastal Plain (fig. 4.5). Southern pine species were restricted to latitudes south of 33° N.
By 14,000 yr B.P., populations of jack and/or red pine that had survived the full-glacial period west of the Appalachians became extinct there. By 10,000 yr B.P., distributions of northern and southern groups of pines were more widely separated geographically than before. Population centers of northern pines had shifted northwest across the Great Lakes region. Southern pines had become important subdominants of the canopy on the southeastern Gulf Coastal Plain.
By 6000 yr B.P., northern pines became structural dominants within the mixed conifer--northern hardwoods forest and the southern half of the boreal forest (J.Terasmae and T.W. Anderson 1970; G.L. Jacobson Jr. 1979). Southern pines spread northward along the Atlantic Coastal Plain to southern New England. In the late Holocene, northern pines became widespread throughout the eastern boreal forest but maintained a primary population center in central Canada. Southern pines became dominants of the southeastern evergreen forest of the southern Atlantic and Gulf coastal plains (figs. 4.4, 4.5). During the late Quaternary, northern pines spread northward at a mean rate of 13.5 km per century; southern pines advanced northward along the Atlantic Coastal Plain at an average rate of 8.1 km per century (P.A. Delcourt and H.R. Delcourt 1987).
The difference between past forest communities and those of today can be measured by using an ordination technique called Detrended Correspondence Analysis (DCA)(fig. 4.6; P.A. Delcourt and H.R. Delcourt 1983, 1987, 1987b). This technique allows the identification of plant communities that may have persisted through long periods of time as well as those that have been ephemeral. We studied the record of the compositional changes in plant communities that have occurred during the past 20,000 years along a south-to-north transect, following longitude 85° W from the Gulf of Mexico to the Laurentide glacier. The glacier retreated northward to Hudson Bay in the middle Holocene (fig. 4.7).
Several kinds of changes in plant communities occurred during the late Quaternary in eastern North America (fig.4.6). The warm temperate vegetation of the southeastern coastal plains has remained relatively constant in its floristic composition through the last glacial-interglacial cycle. The eastern deciduous forest, however, which was restricted to a few small refuge areas during the late Pleistocene, emerged as a major vegetation type only in the Holocene. Today's mixed conifer--northern hardwoods forest has close analogues throughout the Holocene, but in late-glacial times it formed ephemeral communities unlike any in existence today. Since the last glacial maximum, boreal forest communities have exhibited a composition heterogeneous both in space and time. Close modern analogues exist for full-glacial boreal forest communities, but there are only poor modern analogues for the boreal forest communities of the late-glacial/early Holocene transition (P.A. Delcourt and H.R. Delcourt 1987).
From these ordination results, further conclusions can be drawn concerning the degree of vegetational and floristic stability across eastern North America during the past 20,000 years (P.A. Delcourt and H.R. Delcourt 1983, 1987, 1987b). Over this time interval, no major relocations of warm temperate taxa have occurred in the region of the southern coastal plains, and the vegetation there has been in dynamic equilibrium (fig. 4.7). Between 34° N and 43° N (fig. 4.7), the vegetation has shifted from one state of dynamic equilibrium (boreal forest and boreal flora) in full-glacial times to another one (deciduous forest and corresponding temperate flora) in interglacial times, following a late-glacial and early Holocene transition period during which boreal taxa became locally extinct and temperate deciduous taxa immigrated. At higher latitudes (fig. 4.7), successive waves of invasion following deglaciation resulted in continual disequilibrium as tundra and then boreal forest communities became established through the late-glacial and early Holocene intervals. A trend toward equilibrium in species composition occurred only during the last several thousand years of the late Holocene.
Eastern Canadian Arctic and Greenland
During the full-glacial interval, the northernmost glacial margin of the Laurentide Ice Sheet extended from Labrador to the eastern margin of Baffin Island, and westward at about latitude 74° N through the Canadian Arctic Archipelago to Banks Island (J.T. Andrews 1987). Farther north, in the High Arctic zone of the Queen Elizabeth Islands and the corridor between Ellesmere Island and northern Greenland, the Quaternary history is not well known within the area occupied by local ice caps and unglaciated nunataks. During the retreat of the glacial ice from 12,000 until 9000 yr B.P., a rise in sea level and subsequent isostatic uplift of the ocean floor resulted in the exposure of land across formerly glaciated areas of the eastern Canadian Arctic and along the coastal perimeter of the Greenland icecap (references in J.T. Andrews 1985, 1987).
Concomitant with substantial retreat of glacial ice on the Ungava Plateau, the Canadian Archipelago, and Greenland in the early Holocene (10,000--9000 yr B.P.), a pioneer stage of arctic tundra developed on the newly exposed coastal zone (P.J.H. Richard 1977; B.Fredskild 1973, 1985; S.K. Short et al. 1985; H.F. Lamb and M.E. Edwards 1988; J.V. Matthews Jr. et al. 1989). By 10,000 yr B.P. at about latitude 60° N, in coastal zones of northern Quebec and southern Greenland, a High Arctic herb tundra, a sparse cover of arctic or arctic-alpine species of herbs (such as Oxyria digyna, Koenigia islandica, and Saxifraga oppositifolia), ferns, and club-mosses initially colonized the severe periglacial environments of unstable fell-fields and frost-churned soils. This herb tundra community was replaced by a heath tundra between 9100 and 8500 years ago with the invasion of heath dwarf shrubs such as crowberry (Empetrum hermaphroditum) and bilberry (Vaccinium uliginosum). Along both the western and eastern coastlines of Greenland, similar communities developed between 9600 and about 8000 yr B.P. In northwestern Greenland (latitude 75°N) comparable stages of pioneer herb tundra and heath dwarf shrub tundra were established between 8600 and 7700 yr B.P. and 7700 and 6600 yr B.P., respectively (B.Fredskild 1985).
During the middle Holocene interval, shrub populations of shrub willow and dwarf birch (fig. 4.8) invaded northward into Greenland along two dispersal routes, forming communities of Low Arctic shrub tundra. Long-distance dispersal of seeds, probably from northwestern Europe, facilitated colonization by shrub willow (Salix), possibly as early as 9400 yr B.P. in southern Greenland (fig. 4.8A). By 8000 yr B.P. in eastern Greenland, populations of three shrub willows (Salix arctica, S. herbacea, and S. glauca) and one dwarf birch (Betula nana) were successfully established (B.Fredskild 1973, 1985). The maritime corridor of northeastern North America provided a second route for plant migration to southern and western Greenland (J.V. Matthews Jr. et al. 1989). Shrub willows colonized between 8900 and 6700 yr B.P. These were followed by successive invasions of dwarf birches, first Betula nana (8000--6500 yr B.P.), then B. glandulosa (5700--3800 yr B.P.), as well as populations of juniper (Juniperus communis) (7000--6000 yr B.P.) (fig. 4.8B, C; B.Fredskild 1985).
The development of subarctic vegetation in southernmost Greenland coincided with the arrival of tree birch (Betula pubescens) between 3800 and 3600 yr B.P., an example of "sweepstakes dispersal" from European seed sources. Tundra species typically reached their northernmost distributional limits between 9000 and 5000 yr B.P., a time of relatively high seasonal contrast and of peak warm, dry climatic conditions.
Two stepwise changes in climate occurred in the last half of the Holocene interglacial. Both climatic shifts involved increases in precipitation and markedly cooler temperatures, with the first climatic change occurring in the middle Holocene (between 5000 and 3500 yr B.P.) and the second change occurring in the late Holocene (ca. 2000 yr B.P.). Pronounced late Holocene cooling resulted in reduced populations of juniper across coastal Greenland. This juniper decline and reduced amounts of total pollen influx observed at many sites have been interpreted as a progressive replacement of flowering plants by lichens and bog mosses within the arctic tundra vegetation (B.Fredskild 1985).
Western North America
Literature for late Quaternary changes in floras and vegetation has been summarized for western North America (V.M. Bryant Jr. and R.G. Holloway 1985; R.S. Thompson 1988), for the American Southwest (R.G. Baker 1983; W.G. Spaulding et al. 1983; P.V. Wells 1983; B.F. Jacobs et al. 1985; T.R. Van Devender et al. 1987; J.L. Betancourt et al. 1990), for the northwestern United States and western Canada (M.Tsukada 1982; R.G. Baker 1983; C.J. Heusser 1983; J.C. Ritchie 1980, 1987; P. J. Mehringer 1985; C.W. Barnosky et al. 1987; J.V. Matthews Jr. et al. 1989; J. M. Beiswenger 1991), and for northwestern Canada and Alaska (J.V. Matthews Jr. 1974; D.M. Hopkins et al. 1982; T.A. Ager 1983; C.J. Heusser 1983, 1983b; J.C. Ritchie 1984, 1987; C.W. Barnosky et al. 1987; H.F. Lamb and M.E. Edwards 1988; J.V. Matthews et al. 1989).
The American Southwest
During the last full-glacial interval, the Aleutian low pressure system in the North Pacific Ocean intensified, and both the zone of prevailing westerly winds and one branch of the polar jet stream shifted southward (fig. 4.2). This resulted in regional cooling of 7° to 8° C and a persistent Pacific frontal track of winter storms that brought moisture across the American Southwest (G.R. Brakenridge 1978; J.E. Kutzbach and H.E. Wright Jr. 1985; J.E. Kutzbach and P.J. Guetter 1986). Cooling at high elevations produced an elevational lowering of periglacial environments (T.L. Péwé 1983) and regional snowlines. Elevational limits for growth of mountain glaciers were lowered by as much as 1000 m (S.C. Porter et al. 1983). As far south as northern Arizona, the timberline was lowered by 800 to 1000 m (G.R. Brakenridge 1978).
Across the unglaciated portions of southwestern North America, at lower elevations on valley floors and lower montane slopes of closed basins, the cool, moist "pluvial" climate led to the expansion of large freshwater "pluvial" lakes (fig. 4.9; see L. Brouillet and R.D. Whetstone, chap.1). The pluvial lakes produced a lake effect on regional climates, locally enhancing precipitation and reducing seasonal temperature contrasts (L.Benson and R.S. Thompson 1987). Four factors--- increased precipitation, reduced evaporation, high groundwater tables, and increased soil-moisture levels---favored overland water flow. Drainage networks of perennial streams expanded, increasing habitat available for riparian plant communities (G.I. Smith and F.A. Street-Perrott 1983).
Under this full-glacial climatic regime, woody perennials encountered a wide spectrum of habitats through a broad elevational range. Equable full-glacial conditions permitted range extensions of forest plants along their lower elevational and southern distributional limits, where they overlapped the ranges of, but did not necessarily replace, desert scrub taxa (T.R. Van Devender et al. 1987). The comparison of modern ranges of individual species with their Pleistocene fossil occurrences (W.G. Spaulding et al. 1983) indicates a general elevational shift downward in their distributions of 200--1200 m (fig. 4.10). This is especially apparent along the lower elevational limits of many tree species and the upper limits of desert shrub taxa.
This full-glacial reassortment of plant species did not simply produce a "telescoping" of biomes with the same communities of coevolved species. Instead, new combinations of species competed for space, light, and nutrients (K.Cole 1985). Full-glacial pluvial climates favored the expansion of alpine tundra and steppe communities at high elevations, of subalpine forest communities at intermediate elevations, and of woodlands dominated by trees such as juniper (Juniperus spp.) across lower montane slopes (fig. 4.11). Full-glacial biotic communities were bracketed between highland mountain glaciers and lowland pluvial lakes.
Woodland communities rather than desert communities prevailed during most of the Quaternary. Many species that evolved as early as 8 to 5 Ma in the late Miocene (D.I. Axelrod 1979) and that are constituents of the present floras of the Chihuahuan, Sonoran, Mojave, and Great Basin deserts (M.G. Barbour and N.L. Christensen, chap.5) survived as understory plants or in disturbed openings of glacial-age woodlands. In the late-glacial period, with a shift northward of the polar jet stream, the climates changed from a glacial (pluvial) mode to an interglacial (interpluvial) mode. Warmer temperatures, diminished moisture supply, and increased evaporation resulted in pronounced drying and shrinking of pluvial lakes between 14,000 and 10,000 yr B.P. (fig. 4.12).
By 8000 yr B.P., climate-induced reshuffling of plant species in the late Pleistocene and early Holocene intervals produced new desert scrub and desert communities at low elevations that displaced woodlands and invaded exposed playa flats (former pluvial lakes; fig. 4.12). Forest communities persisted at intermediate elevations and expanded to higher mountain summits (K.Cole 1985).
During the early and middle Holocene, the winter peak in precipitation favored one suite of taxa characteristic of the Mojave and Great Basin deserts (T.R. Van Devender et al. 1987). The early Holocene development of summer monsoons with a new summer peak in moisture, the result of the seasonal shift of the Maritime Tropical air mass expanding northwestward from the Gulf of Mexico, favored a second suite of desert plants characteristic of the modern Chihuahuan and Sonoran regions. Migration of desert taxa continued to enrich modern desert scrub and desert grassland communities until stabilization of the floras was attained at about 4000 yr B.P. (W.G. Spaulding et al. 1983; T.R. Van Devender et al. 1987; R.S. Thompson 1988).
In the coastal zone of western California, in southern Arizona, and in western Mexico, the interglacial (interpluvial) climate was characterized by moderate temperatures, cool wet winters and hot dry summers (Mediterranean-type climate), and seasonal droughtiness. Fires increased in frequency, particularly during the Holocene interglacial peak in warmth and aridity. Sclerophyllous scrub vegetation, including chaparral, became more widespread during the middle Holocene Hypsithermal interval from about 8000 to 4000 yr B.P. (D.I. Axelrod 1989).
Chaparral taxa evolved over the past 50 million years in response to the generation of suitable montane habitat, and their populations expanded in response to major tectonic uplift in the late Pliocene and Quaternary. While chaparral plants were not the evolutionary product of Mediterranean-type climates established during the Quaternary interglacials, their predominance on the landscape has been affected by glacial-interglacial cycles of climatic change, which have alternately suppressed and then augmented wildfire disturbance regimes (D.I. Axelrod 1989). These chaparral taxa have presumably assembled into Mediterranean scrub communities during interpluvial periods and disassembled during pluvial periods. Earlier throughout the Tertiary period and during the pluvial glacial ages of the Quaternary, the chaparral taxa persisted as understory scrub within woodland communities.
Northwestern United States and Southwestern Canada
During full-glacial times, the Cordilleran Ice Sheet extended from southern Alaska through British Columbia to its southernmost limit in northern Washington, Idaho, and Montana (D.B. Booth 1987; J.J. Clague 1989). It flowed westward across a 50-km stretch of Pacific Coastal Plain exposed by a lowering of sea level by at least 100 m (C.W. Barnosky et al. 1987). Small pockets of unglaciated land (J.J. Clague 1989), such as in the Queen Charlotte Islands in coastal British Columbia, served as glacial-age, floristically diverse refugia for herbs, willow, Sitka spruce (Picea sitchensis), and lodgepole pine (Pinus contorta) (B.G. Warner et al. 1982; J.V. Matthews Jr. et al. 1989).
Along the eastern base of the Rocky Mountains, glacial ice flowing eastward from the Cordilleran Ice Sheet merged with the western and southwestern limits of the Laurentide Ice Sheet, resulting in continuous late Wisconsinan ice contact of these two ice sheets from west central Alberta, across northeastern British Columbia, to the southeastern corner of the Yukon (N.W. Rutter 1984). Thus during full-glacial time 20,000 to 18,000 years ago, a complex of extensive continental glaciers stretched across North America from the eastern coast of the Pacific Ocean to the northwestern margin of the Atlantic Ocean (J.T. Andrews 1987; D.B. Booth 1987; J.J. Clague 1989).
During peak glaciation 18,000 yr B.P., this continental mass of glacial ice produced a climatic regime of "glacial anticyclone" (fig. 4.2), generating strong, cold winds that radiated in a clockwise direction off the southern Cordilleran and Laurentide ice margins and swept toward the west and southwest (J.E. Kutzbach and H.E. Wright Jr. 1985). These strong easterly winds maintained severely cold and dry conditions through most of the unglaciated northwestern United States. Rigorous permafrost environments with patterned ground, ice wedges, and cryoplanation terraces developed across Montana, Wyoming, Idaho, and the eastern two-thirds of Washington (T.L. Péwé 1983; J.M. Beiswenger 1991).
Westerly winds sweeping across the Pacific Ocean brought cool, moist conditions only to the maritime area west of the Coast Ranges, the first orographic barrier encountered. The eastward transport of oceanic moisture by westerly winds was effectively stopped by the combination of the prevailing, periglacial easterly winds and by the coastal mountain ranges, reinforcing the rain shadow farther east. Because it is directly related to distance from moisture source, alpine glaciation was most prominent in the Pacific maritime zone of the Coast Ranges and Cascade Range. Formation of mountain glaciers was progressively more limited farther to the east in the Rocky Mountains (S.C. Porter et al. 1983).
C.J. Heusser (1983) and C.W. Barnosky et al. (1987) reviewed the patterns of vegetation in the Pacific Northwest relative to the last advance and retreat of the Cordilleran Ice Sheet during the last 20,000 years. South of the glacial limit in western Washington 20,000--18,000 yr B.P. (fig. 4.13), cool, humid conditions favored the growth of alpine glaciers in the Olympic Mountains of the Coast Ranges and inland in the Cascade Range. On the exposed Pacific Coastal Plain between 20,000 and 16,800 yr B.P., the vegetation was a subalpine parkland with a mixture of lowland and montane species, including Sitka spruce (Piceasitchensis, western white pine (Pinus monticola), lodgepole pine (P. contorta), mountain hemlock (Tsuga mertensiana), and western hemlock (T. heterophylla) (Bogachiel Valley, fig. 4.13; C.J. Heusser 1983). In the rain shadow east of the Olympic Mountains (the first Coast Range), the lowland vegetation of the Puget Trough was composed of a mosaic of grass-sedge-wormwood (Artemisia) tundra and parkland with trees of lodgepole pine and probably Engelmann spruce (Picea engelmannii) (Davis Lake and Battle Ground Lake, fig. 4.13; C.W. Barnosky et al. 1987).
In the Columbia Basin, east of the glacier-mantled Cascade Range, a sparse vegetation of periglacial steppe persisted from 20,000 to about 10,000 yr B.P. Under the influence of cold, extremely dry, easterly winds, this lowland steppe vegetation consisted of grasses, sagebrush (Artemisia), and alpine herbs (Carp Lake, fig. 4.13). A forest zone may have been compressed elevationally between the mountain glaciers and the lowland steppe, surviving at intermediate elevations along the eastern slope of the Cascade Range, or, alternatively, alpine tundra at its lower limit may have merged with steppe at its upper limit (C.W. Barnosky et al. 1987). Either situation would have allowed the combination of alpine and steppe floras within this important periglacial steppe vegetation (fig. 4.13) that extended from central Washington east to Wyoming (P.J. Mehringer Jr. 1985; C.W. Barnosky et al. 1987; J.M. Beiswenger 1991).
Between 17,000 and 15,000 yr B.P., increased seasonal contrast in temperatures and the warming of summers were first expressed by retreat of alpine glaciers toward mountain summits (S.C. Porter et al. 1983) and the shift from subalpine parkland toward open forest on the lowlands of western Washington (fig. 4.13A; C.J. Heusser 1983). Late-glacial forests diversified from about 16,000 until 12,500 yr B.P. Major expansions of tree populations, particularly of Sitka spruce and red alder (Alnus rubra), occurred on the Pacific Coastal Plain. Inland in the Puget Trough, lodgepole pine colonized deglaciated terrain (C.W. Barnosky et al. 1987).
Substantial melting and retreat of the Cordilleran Ice Sheet occurred between 12,500 and 10,000 yr B.P. in the Pacific Northwest. This reduced glacial mass diminished both the glacial source for and the strength of anticyclonic easterly winds. The enhanced prevailing westerly winds from the Pacific Ocean resulted in widespread warming of temperate postglacial climates. Closed forests of red alder, spruce, lodgepole pine, and hemlocks developed west of the Cascade Range. Populations of temperate trees, such as grand fir (Abies grandis), western hemlock, and mountain hemlock, shifted elevationally from lowlands into nearby montane habitats, and these trees also migrated northward into suitable maritime areas of temperate, humid conditions (fig. 4.13B; C.W. Barnosky et al. 1987).
At this time of transition between glacial and interglacial climatic regimes, the western slopes of the northern Rocky Mountains served as an increasingly effective orographic barrier, screening moisture from the now-prevailing westerly winds. By about 10,500 yr B.P. in western Montana (Tepee Lake, fig. 4.13B), the periglacial tundra was replaced as open coniferous forests were established, dominated by a variety of pines that probably included western white pine (Pinus monticola), white bark pine (P. albicaulis), and lodgepole pine (P. contorta), as well as by subalpine fir (Abies bifolia) and spruce (Picea engelmannii and possibly white spruce [P. glauca]) (R.N. Mack et al. 1983). Tundralike grasslands or sagebrush steppes persisted within dry rainshadow regions directly east of the Cascade Range and in the Great Plains east of the northern Rocky Mountains (fig. 4.13B; R.G. Baker 1983; C.W. Barnosky et al. 1987).
During the Holocene epoch, warmest and driest summers (fig. 4.3) occurred between 10,000 and 7000 yr B.P., accentuating drought stress during the growing season, especially for plants on well-drained or xeric sites (C.W. Barnosky et al. 1987). In response, early Holocene forests developed more open canopies. Mesophytic conifers, such as spruce and hemlock, became less important in these forests, whereas more drought-tolerant species, such as Douglas-fir (Pseudotsuga menziesii) and red alder, became dominant (R.G. Baker 1983; C.W. Barnosky et al. 1987).
Northwestern Canada and Alaska
During the maximum extent of continental glaciation, unglaciated regions with cold, dry climates and sparse treeless vegetation occurred across Beringia to the Mackenzie River Delta in the Northwest Territories (south to 60° N along the margins of the Cordilleran and Laurentide ice sheets) (D.M. Hopkins et al. 1982; J.C. Ritchie 1984; J.V. Matthews Jr. et al. 1989). In this region, full-glacial fossil-pollen assemblages of wormwood (Artemisia), grasses, and sedges have low total values of pollen influx. This has been interpreted to indicate a landscape mosaic that included many types of herbaceous tundra, interspersed by periglacial areas of polar desert communities and open ground maintained by freeze-thaw processes (C.E. Schweger 1982; J.V. Matthews Jr. 1982; T.A. Ager 1983; J.C. Ritchie 1984, 1987; C.W. Barnosky et al. 1987; H.F. Lamb and M.E. Edwards 1988). This interpretation differs from earlier proposals of either a continuous "arctic steppe" herb tundra (J.V. Matthews Jr. 1976; R.D. Guthrie 1984) or of barren polar desert or fell-field (L.C. Cwynar and J.C. Ritchie 1980; J.C. Ritchie and L.C. Cwynar 1982). Low icecaps occupied the Alaskan Brooks Range and the Ahklun Mountains (T.D. Hamilton and R.M. Thorson 1983).
The Cordilleran Ice Sheet served as a major barrier to the northward dispersal of boreal and temperate plants during full-glacial times. At the maximum extent of Pleistocene glaciation, 20,000 to 18,000 years ago, relatively few summits of the Coast Ranges, MacKenzie Mountains, and Rocky Mountains rose above the Cordilleran glacial ice surface. Generally above 2500 m elevation, these isolated montane summits provided sites for alpine glaciers and periglacial "nunataks," serving as possible refugia for polar desert and periglacial steppe communities (fig. 1.12 in J.J. Clague 1989).
During the late-glacial interval the southern margin of the Cordilleran glacier rapidly retreated northward after 14,000 yr B.P. It retracted to the United States--Canada border by 11,000 yr B.P. (D.B. Booth 1987). Between about 13,500 and 11,000 yr B.P., deglaciated terrain and proglacial lakes formed an ice-free corridor between the retreating margins of the Cordilleran and Laurentide ice sheets (N.W. Rutter 1984), creating a continuous corridor for plant and animal migration across the western interior of Canada (J.C. Ritchie 1980, 1987; J.V. Matthews Jr. et al. 1989). This midcontinental route was used by both eastern and western plant species between 13,500 and 9000 yr B.P.
Boreal trees, such as white spruce (Picea glauca), black spruce (P. mariana), and jack pine (Pinus banksiana), migrated northward along this route from refugia in the western Great Lakes region (J.C. Ritchie and G.M. MacDonald 1986; J.C. Ritchie 1987). Western populations of lodgepole pine expanded from a full-glacial refuge in the unglaciated portions of the Pacific Northwest and initiated their northward migration through this corridor more than 12,200 years ago (G.M. MacDonald and L.C. Cwynar 1985; L.C. Cwynar and G.M. MacDonald 1987).
Cordilleran ice retreated across the Pacific Coastal Plain between 13,500 and 9500 yr B.P. The plain was then repeatedly uplifted by isostatic rebound and submerged by rising sea levels (J.J. Clague 1989). The Aleutian Islands were deglaciated between 12,000 and 10,000 yr B.P., and the Gulf of Alaska was ice-free by 10,000 yr B.P. The present coastline of western British Columbia and southern Alaska became free of glacial ice between 13,500 and 9500 yr B.P., providing a western maritime route for plant invasions (T.D. Hamilton and R.M. Thorson 1983; J.J. Clague 1989).
In unglaciated regions of eastern Beringia, i.e., Alaska, the northern Yukon, and Northwest Territories, full-glacial climatic conditions provided a longitudinal gradient of cool, mesic conditions in the west and much colder and drier conditions nearer the continental ice sheets. This environmental gradient was reflected in the change from mesic tundra meadows of grasses, sedge, Artemisia, and shrub willow (Salix) in western Alaska to more xeric, discontinuous herb tundra and polar desert communities in northwestern Canada (fig. 4.14; T.A. Ager 1983; J.C. Ritchie 1984, 1987).
By 14,000 yr B.P., climate ameliorated across Beringia. Summer temperatures rose concomitant with the increased seasonality of solar radiation. Coastal portions of the Bering land bridge were inundated as sea levels rose (C.W. Barnosky et al. 1987). Increased availability of water from melting glacier margins, and closer proximity to coastal waters, combined to produce more precipitation. This late-glacial shift toward warmer, wetter summers favored the expansion of formerly small, scattered populations of dwarf birch and many heath species (family Ericaceae). The formerly extensive herb tundra was replaced at about 14,000 yr B.P. by birch--heath shrub tundra through virtually all of the Alaskan mainland (T.A. Ager 1983; C.W. Barnosky et al. 1987). Maritime meadows of herb tundra, dominated by grasses and sedges, were established on deglaciated terrain. These have persisted for the last 10,000 years on the chain of the Aleutian Islands, which were geographically isolated by Holocene sea-level rise.
The opening of the ice-free corridor in the interior of western Canada between 13,500 and 11,000 yr B.P. provided the opportunity for a sequence of rapid northward immigrations of plant taxa between the southern Alberta corridor portal and the northern portal of the unglaciated Yukon and Northwest Territories. Other species migrated southward from the Beringian refugium onto terrain recently exposed by retreating glaciers, many tracking along the eastern edge of the Rocky Mountain foothills southward for thousands of kilometers (J.C. Ritchie 1984; J.V. Matthews Jr. et al. 1989). By 10,500 yr B.P., herb tundra communities were invaded by shrubs such as dwarf birch and heaths.
Passage of a fundamental bioclimatic threshold 10,000 years ago shifted the competitive balance from dominance of late Pleistocene herbaceous plants to arboreal plants in the early Holocene. J.C. Ritchie et al. (1983) presented environmental and paleobotanical evidence to suggest the existence of a regional maximum in summer temperatures at high latitudes in Beringia 10,000 years ago, a climatic phenomenon related to maximum seasonal contrast that was a consequence of Milankovitch cycles (fig. 4.3). Between 10,000 and 9000 yr B.P., large shrubs and pioneer trees became established, including willow, juniper (Juniperus communis), balsam poplar (Populus balsamifera), and white and black spruce. The modern landscape mosaic of herb and shrub tundra and woodland was established by 9000 yr B.P. in the far northwest of Canada (J.C. Ritchie 1984).
Populations of trees established in the Yukon in the early Holocene provided the seed source for dispersal into central and western Alaska along the westward-flowing network of the Yukon River and its tributaries (fig. 4.14). The forests of central Alaska, containing white and black spruce, white birch (Betula papyrifera), balsam poplar, and trembling (quaking) aspen (Populus tremuloides), were established between 10,000 and 8000 yr B.P. (fig. 4.14). By about 5000 yr B.P., they were present along their current western limit (C.J. Heusser 1983; T.A. Ager 1983).
The deglaciated coast of western British Columbia and southern Alaska provided the second principal route for postglacial plant migration. By 11,000 yr B.P., lodgepole pine and red alder colonized disturbed sites, forming an open parkland as far north as the panhandle of southeastern Alaska. Coastal coniferous forests were established across northwestern British Columbia and southeastern Alaska. The entry of Sitka spruce occurred between 10,500 and 8500 yr B.P. Successive invasions of western and mountain hemlock occurred in this region between about 8500 and 2700 yr B.P. Coastal forests dominated by Sitka spruce and mountain hemlock reached their modern western limits in south central Alaska within the last 4000 years (T.A. Ager 1983; C.W. Barnosky et al. 1987). This relatively late northwestward extension of Pacific coastal forests has been attributed to a late Holocene increase in the frequency of storms and the supply of moisture they brought (C.J. Heusser 1983b; C.W. Barnosky et al. 1987).
R.J. Hebda and R.W. Mathewes (1984) documented the postglacial advance of western redcedar (Thuja plicata), an endemic species that is today an important component of the coniferous forests of the cool, moist Pacific coastal slope. Increased percentages of pollen grains similar to that of western redcedar in the fossil record indicate that it initially colonized the Puget Trough of western Washington as early as 10,000 years ago. Populations of western redcedar advanced through the deglaciated Fraser Lowlands of southwestern British Columbia between 8000 and 6500 yr B.P., achieving their present northern range in northwestern British Columbia by about 6000 years ago. A maritime climatic shift toward increased moisture favored a major increase in population size of western redcedar from 5000 to 2000 yr B.P. This late Holocene expansion of cedar resulted in the relatively recent codominance in Pacific coastal forests of western redcedar and western hemlock, the previous Holocene dominant.
The modern flora of North America has developed as the constituent plant taxa were evolving at different times and in response to different kinds of selection pressures (B.H. Tiffney 1985, 1985b). The development of the flora during the last 3 million years of the late Tertiary and Quaternary reflects contributions from three groups of floristic elements: (1) "relict floras" that evolved much earlier in the Mesozoic or early Cenozoic and that persist as relatively old, unchanging relicts of past floras; (2) "orthoselection floras" that include a suite of species changing in response to a consistent, long-term environmental trend; and (3) "migration floras" that exhibit substantial shifts in distributional ranges as they migrate across relatively large distances, tracking major climatic changes between alternating glacial and interglacial regimes (V.P. Grichuk 1984).
The relatively ancient or relict floras comprise taxa that have survived the late Cenozoic onset of increasing environmental and climatic oscillations of ever-greater magnitude (E.B. Leopold and H.D. MacGinitie 1972; D.Q. Bowen 1985), and that have persisted to the present day within refugial regions such as the Gulf Coastal Plain (P.S. Martin and B.E. Harrell 1957; E.L. Little Jr. 1971b; A.Graham 1972b, 1973b; A.J. Sharp 1972, 1972b; see A.Graham, chap. 3). In part, these relict floras represent plants with closely related vicariant species, species pairs with disjunct populations in widely separated regions of the world today (H.-L. Li 1972). Elements of these floras have been displaced by the gradual plate-tectonic movements of continents, and their distributional ranges have been fragmented by mountain building, climatic change, and loss of migrational corridors such as the Beringian and North Atlantic land bridges (J.A. Wolfe 1972; B.H. Tiffney 1985, 1985b; D.I. Axelrod 1986, 1986b).
Two Quaternary examples illustrate orthoselection floras (V.P. Grichuk 1984). These are floras that have responded to long-term trends in climate and in regional plift of terrain. At high latitudes during the late Pliocene and early Pleistocene, long-term climatic cooling shaped the patterns of taxonomic differentiation and favored the latitudinal expansion of tundra and taiga communities across the Arctic (D.I. Axelrod 1986b). This late Cenozoic, climate-driven shift in both composition and dynamics of tundra ecosystems provoked the in situ evolution of plant and animal taxa (H.Hara 1972; J.A. Wolfe 1972; "phyletic ecosystem evolution" sensu J.V. Matthews Jr. 1974).
At middle latitudes in the American West, the combination of climatic cooling and continued late Cenozoic mountain uplift, particularly within the Coast Ranges, the Cascade Range, the Sierra Nevada, and the southern Rocky Mountains, intensified the orographic influence of the rainshadow effect within and immediately east of the north-south trending ranges of the Cordillera. This long-term cooling, decreased summer precipitation, and increased continentality of montane sites both impoverished the modern Cordilleran woody flora and enhanced the elevational differentiation of alpine tundra and herbaceous steppe floras, as it did also with assemblages of subalpine arboreal conifers (E.B. Leopold and H.D. MacGinitie 1972; D.I. Axelrod and P.H. Raven 1985; D.I. Axelrod 1986, 1986b, 1988; C.W. Barnosky 1987).
W.B. Critchfield (1984) used data from plant fossil and contemporary genetics to develop a model of response of the North American temperate and boreal conifers to glacial-interglacial cycles. He suggested that, as each episode of continental glaciation was initiated, the combination of areal expansion of glacial ice masses and shifting climatic zones would displace some conifer populations, resulting in both contraction and fragmentation of their ranges. During times of maximum continental glaciation, the fragmented conifer populations may undergo genetic differentiation within isolated refugia. With the onset of interglacial conditions, migration and expansion of distributional ranges would reestablish genetic exchanges between populations within species and among closely related species that were formerly isolated.
The model by W.B. Critchfield (1984) identifies four principal kinds of Pleistocene climatic impacts on the genetic structure of conifer populations and on patterns of floristic change: (1) reduced genetic variation, (2)extinction of taxa, (3) geographic and temporal redistribution of genetic variation, and (4) increased genetic variation in a taxon. The first of these population-level responses to Quaternary environmental changes is the long-term loss of genetic variability. This may be reflected in reduced "ecological amplitude" and progressive restriction in the spectrum of habitats that plants occupy (R.O. Kapp 1977).
In extreme cases, extinction may result, a second type of population response. For example, in the late-glacial interval 11,000 years ago, a large-coned morphotype of white spruce (Picea glauca) became extinct in the southeastern United States (P.A. Delcourt et al. 1980; W.A. Watts 1980b; P.D. Royall et al. 1991). W.B. Critchfield (1984) suggested that this now-extinct fossil spruce may have represented a distinct species rather than an ecotype of white spruce.
Populations may also respond to change by redistribution of genetic variation. During the Quaternary, this may have occurred in two ways within North American conifers (W.B. Critchfield 1984). First, with the onset of glacial climates and more severe environments, genetic impoverishment might have resulted from the early-glacial fragmentation and local elimination of populations, with the genetic pool residing in transitory, migratory races. Second, geographic races that were isolated in glacial-age refugia might have experienced relatively infrequent and brief flushes of genetic exchange during interglacial intervals. Within the current distributional range of some trees, such as lodgepole pine (Pinus contorta), geographic gradients in genetic structure from central to marginal populations may reflect interglacial paths of migration that have resulted from multiple and successive events of long-distance dispersal and subsequent establishment of outlier populations (G.M. MacDonald and L.C. Cwynar 1985; L.C. Cwynar and G.M. MacDonald 1987).
The last type of Pleistocene climatic impact on the coniferous tree flora is that of enhanced genetic variation. Opportunities were generated for hybridization and introgression as interglacial expansion in distributional ranges removed intervening barriers and facilitated genetic exchange among populations both within species and among closely related species groups (W.B. Critchfield 1984).
Through the last 900,000 years of the Quaternary, a period of 90,000 years of each 100,000-year glacial-interglacial cycle has been characterized by gradual cooling (N.G. Pisias and T.C. Moore Jr. 1981), and the onset of each 10,000-year interglacial interval has been marked by rapid climatic warming and increased seasonality of climates (fig. 4.15; J.Imbrie and K.P. Imbrie 1979). Because climates and environments have not been constant during the Quaternary, we might anticipate that any evolution that has occurred within the North American flora during the past several million years has not been gradualistic. Rather, we hypothesize that bursts or pulses of macroevolution (S.M. Stanley 1979) would have occurred every 100,000 years, modulated by the climatic and environmental changes triggered by Milankovitch cycles and coinciding with the transition from glacial to interglacial climatic regimes (fig. 4.1).
The relatively short time span from peak glacial to peak interglacial conditions is the time of greatest environmental instability during a glacial-interglacial cycle (fig. 4.3), and it is characterized both by the greatest magnitude and the greatest rate of change of global temperature (fig. 4.15). We postulate that a pulse of extinctions would result during times of rapid change from glacial to interglacial conditions, because the thresholds of temperature tolerance of plant species would be exceeded during these periods of high seasonal contrast. Extinctions would trigger a subsequent wave of speciation as other taxa diversify and fill the recently vacated niches. Minor readjustments would continue into the interglacial interval, so that eventually the rate of speciation would offset the rate of extinction. During the last stages of each interglacial interval and into the next developing glacial interval, evolution would be gradualistic (fig. 4.15).
During glacial intervals North American plant populations would be fragmented in two ways: (1) mountain and continental glaciers expand to occupy formerly vegetated areas and (2) climatic cooling and shifts in climatic zones result in displacement of migration floras in unglaciated regions (R.A. Bryson and W.M. Wendland 1967; P.A. Delcourt and H.R. Delcourt 1984; J.E. Kutzbach and H.E. Wright Jr. 1985). For example, in the American West, full-glacial conditions provoke elevational adjustments in species ranges and their dispersal along montane corridors (W.G. Spaulding et al. 1983; P.V. Wells 1983; K.Cole 1985; C.W. Barnosky et al. 1987; T.R. Van Devender et al. 1987; R.S. Thompson 1988). Likewise, in middle latitudes of the unglaciated Southeast, a widespread latitudinal displacement in species distributions occurs during glacial times (M.B. Davis 1981, 1983, 1986; W.A. Watts 1983; P.A. Delcourt and H.R. Delcourt 1987; T.WebbIII 1988). Thus glacial-age conditions may isolate populations within separate refugia, facilitating genetic drift and interpopulation heterogeneity.
During each glacial-interglacial transition, floras would be affected in several ways. Heightened seasonal contrast would produce unstable and retreating glaciers, resulting in availability of newly deglaciated territories for plant colonization. Increased seasonality of climate would also release some plant species whose populations were formerly constrained by some limiting bioclimatic threshold factor. Therefore individualistic plant migrations and the intermingling of floras would be facilitated (W.A. Watts 1988). The extinction of some species would occur because of geographic restriction coupled with a loss of genetic variability. In contrast, opportunities for genetic divergence leading to new ecotypes, subspecies, and species would increase for other plants, as selection intensified because of environmental change and competition accentuated among ephemeral assemblages of migrating species.
In the region of the unglaciated southeastern United States, the results of both types of processes can be seen in the development of the Holocene interglacial flora. With the onset of the present interglacial, climatic warming and increased seasonality caused the decline of formerly widespread populations of eastern spruce species along their southern periphery. Meanwhile, the slow rate of melting of the Laurentide Ice Sheet prevented the rapid spread of spruce northward and resulted in an environmental bottleneck that caused the collapse of former population centers of spruce, the extirpation of fragmented populations along the southern periphery of their range, and the extinction of one morphotype or species of white spruce (J.C. Bernabo and T.WebbIII 1977; P.A. Delcourt et al. 1980; W.B. Critchfield 1984; P.A. Delcourt and H.R. Delcourt 1987).
In the central and southern Appalachian Mountains, populations of arctic-alpine and boreal taxa were progressively restricted during the early Holocene to high mountain peaks. Interglacial range restrictions may have resulted in speciation of narrowly endemic plants such as Geum peckii (P.S. White 1984b), as well as of certain boreal and cool temperate trees including red spruce (Picea rubens), Fraser fir (Abies fraseri), and Carolina hemlock (Tsuga caroliniana) that are today restricted in distributional range to the central and southern Appalachian region (E.L. Little Jr. 1971b).