Quaternary History of Eastern Ontario: Impacts on Physical Landscape and Biota

Stephen C. Lougheed and Natalie Morrill
Department of Biology, Queen’s University, Kingston, Ontario Canada K7L 3N6
email: steve.lougheed@queens.ca, natalie.morrill@gmail.com

    Events of the last two and one-half million years profoundly influenced the physical and biotic landscape of northern North America. The immense impact of repeated glaciations on the region we now know as Ontario is widely acknowledged, yet many questions remain, including the role of glaciations in generating new species, the routes of colonization into Canada after the last glacier receded, and their influence on the physiognomy of Ontario. Here we provide a brief overview of the Quaternary period framing the discussion as it relates to landscapes and biodiversity of Canada and particularly Eastern Ontario.

Background

The geological time scale is divided hierarchically into a nested series of temporal classifications: eons, eras, periods, and epochs (Figure 1). Some classify it even more finely with epochs further split into ages. These divisions are based on chronostratigraphic (order of rock deposition and radiometric dating of strata), biostratigraphic evidence (sequence and dating of fossil assemblages) and other evidence (e.g. shifts in sea levels, shifts in Earth’s magnetic fields). Divisions between consecutive geological divisions may be based on major events (e.g. mass extinction between Cretaceous and Paleogene) or more subtle differences among strata.

Figure 1. Geological time scale with different temporal divisions at increasingly finer scales from eons on the left through to epochs on the extreme right. Note that the dates are approximate as there is much contention over these. The Hadean is often not considered a formal eon nor is the division between the Archean and Hadean particularly precise. The Quaternary is divided up into Pleistocene and Holocene, the latter of which began approximately 12,000 years ago and is comprised of the thin rectangle above the Pleistocene in the temporal sequence of Epochs to the right of this figure. PH = Phanerozoic. CE = Cenozoic. Modified from Gradstein et al. (2004) and the British Geological Survey.

The Quaternary is the last period of the Cenozoic, split into two epochs: the Pleistocene and the Holocene. The Pleistocene spans some two and one-half million years during which there were repeated glaciations that hugely impacted the Northern Hemisphere, while the Holocene covers the last approximately 12,000 years following the last glacial retreat. There has been debate as to whether the Quaternary merits true designation as a geological period and vociferous discussions as to the span of time it encompasses (essentially the temporal boundary between Pliocene and Pleistocene epochs). This issue seems to have been resolved with the Quaternary (and thus Pleistocene) beginning at 2.58 mya (Gibbard et al. 2010).

Glacial deposits and other signatures of glaciation were traditionally used to demarcate four major glaciations separated by warm interglacial periods: the Nebraskan, the Kansan, the Illinoian and the Wisconsinan (each named after the US state where the southernmost terminus of each respective ice sheet was purported to have reached; Redfern 2001 – this is the older parlance and we may update this in the future). The challenge here is that each subsequent glacial event erased physical evidence of the preceding one and it has become increasingly clear that there were not four but many glaciations over the Quaternary. For example, recent indirect evidence, including studies of the ratio of oxygen isotopes in marine sediments (δ18O – the ratio of 18O to 16O), implies that global climate cooling events may have corresponded to tens of glacial oscillations over the last 2.5 million years (Sosdian and Rosenthal, 2009). Each major glaciation involved massive ice sheets covering most or all of present-day Canada and a portion of northern continental USA with maximum thickness in excess of 2000 metres (i.e. approximately equivalent to the summed length of 20 soccer fields).

Formation of the Great Lakes Landscape
Of particular interest to those studying the landscape of central Canada is the impact of repeated glaciations on the Laurentian Great Lakes Basin. The Great Lakes watershed, which today covers an area of 765,990 km2 across Ontario, Canada, and eight of the northern U.S. states, is the product of multiple glaciations and redirected water drainage during the retreat of the Laurentide Ice Sheet, the most recent of several massive ice sheets that have covered North America (Larson and Schaetzl 2001). The Laurentian Great Lakes watershed can be divided, geologically, into a southern lowland region (particularly the Erie, Michigan, Huron, and Ontario basins) defined by Paleozoic, relatively non-deformed and horizontally bedded sedimentary bedrock, and a northern upland region (most of Lake Superior and Georgian Bay basins, as well as parts of Lake Ontario), which has highly deformed Precambrian metamorphic rock as its bedrock, usually referred to as the Canadian Shield (Larson and Schaetzl 2001). (Refer back to Figure 1 for a sense of the bedrock’s relative age.) These differences in bedrock type are the main reason there are regional differences in response to the movement of the ice sheets, and help to define the extent and morphology of the historical and current Great Lakes.

As mentioned, geological evidence for previous glaciations is difficult to obtain, as subsequent glaciations tend to obliterate previous sedimentary records and alter landscape features even further. The two most recent glacial periods in North America, and those for which the best geological records have been preserved, are the Illinoian (302-132 ka BP) and Wisconsinan (79-10 ka BP) glaciations. In Canada, almost all geological evidence for the Illinoian glacial period has been eradicated by more recent glacial activity; however, the nature of existing sediment and fossil records of plants and aquatic species, such as diatoms, suggest that the last interglacial period was similar in climate to the current one within the Great Lakes region, or perhaps a few degrees Celsius warmer (reviewed in Larson and Schaetzl 2001).

The Wisconsinan glaciation began 79-75 thousand years “before present” (ka BP) (Gilbert 1994, Larson and Schaetzl 2001). Ice advanced from the northeast of the Great Lakes watershed, and dammed the Lake Ontario basin between 79-65 ka BP (Karrow 1984). From 65-35 ka BP, the margin of the ice sheet advanced and retreated within the Lake Ontario basin, bordered by a proglacial lake (i.e. a lake with a glacial margin), while the southern parts of the Great Lakes watershed (including the Lakes Erie, Huron and Michigan basins) seem to have remained ice-free from 64.5-25 ka BP (reviewed in Larson and Schaetzl 2001). Between 55-27 ka BP, ice may have advanced into the upper Mississippi River valley; if this is the case, then the ice sheet likely extended into the Lake Superior basin and perhaps into the northern end of the Lake Michigan basin by the middle Wisconsinan (Grimley 2000). By 27.5 ka BP, ice extended into the Lake Erie basin but probably did not reach beyond the southern shore of present-day Lake Erie (Fullerton 1986)

Figure 2. Ice coverage of North America during the Wisconsinan Ice Age, at about 18,000 years before present. Adapted from USGS

The period from 24-16 ka BP saw a major ice advance, with glaciers extending beyond southwestern Ontario (Terasmae 1981). In that same period advancing ice eventually covered the entire Great Lakes watershed (Grimley 2000). The maximum extent of the Laurentide Ice Sheet, or Last Glacial Maximum (LGM), occurred approximately 18,000 years ago (Figure 2). During peak glacial conditions, the ice sheet was about 2 km thick over Eastern Ontario and as much as 3 km thick to the north over what is now Hudson Bay. At this time, in the east, ice extended almost as far south as present-day New York City, and throughout northern Pennsylvania (Gilbert 1994). The eastern part of the ice margin was constrained by the Appalachian plateau (extending across several of the present-day eastern States), whereas its southward extent in the west was greater: the ice sheet had by about 20 ka BP reached south-central Illinois (Hansel and Johnson 1992).
Orbital changes in the Earth’s revolution around the sun are considered to be the primary “pacemakers” of ice ages on our planet (Hays et al. 1976), and by about 15 ka BP, the geometry of the Earth’s orbit began to allow for higher insolation (i.e. more radiant energy from the sun on a given area of the Earth’s surface). The timing of this orbital change appears to coincide with the disintegration and subsequent northern retreat of the Laurentide Ice Sheet (Delcourt 1980), and by as early as 16 ka BP the first in a sequence of proglacial lakes began to form in the Great Lakes watershed. What followed was a series of regional ice sheet advances (stadials) and retreats (interstadials). In the first few thousand years of ice retreat, proglacial lakes formed in the southern Lake Michigan basin and the Erie basin, though these lasted for only 1000 years, and were destroyed as the ice re-advanced between 15.5-13.5 ka BP to south of present day Lake Erie. Subsequent ice retreat produced the first pro-glacial lake in the Huron basin. It is interesting to note differences between the drainage of these early pro-glacial lakes and that of the present-day Great Lakes: glacial Lake Maumee, for example, which formed in the Erie Basin during the second period of ice retreat (the Mackinaw Interstadial), drained to the southwest.
The first pro-glacial lake that formed in the Ontario basin, glacial Lake Iroquois, began to form after about 12.5 ka BP, and drained through an outlet near Rome, New York (Terasmae 1981). The period between 12.5 and 11.8 ka BP (the Two Creeks Interstadial) was an eventful one in the natural history of the Lake Ontario Basin. Pollen from lake sediments at the time indicate that spruce was growing in the Toronto area by 12.5 ka BP (Terasmae 1981), and ice had withdrawn completely from the Lake Ontario basin by 12.4 ka BP. The retreating ice front acted as a dam that allowed glacial Lake Iroquois to form, and the lake reached its maximum extent at about this time, before draining in stages marked by beaches still visible around Trenton, Ontario (Gilbert 1994). By 12 ka BP, all of southeastern Ontario was ice-free (Gilbert 1994).

The St. Lawrence Valley became ice-free at about 11.8 ka BP, and Lake Ontario began to drain via this outlet. At the same time, an incursion of the Atlantic Ocean known as the Champlain Sea spread into the isostatically depressed Ottawa and upper St. Lawrence valleys. Isostatic depression refers to the effect that occurs when the massive thickness of an ice sheet loads the earth’s crust to such an extent that it causes the crust to sink. When the ice retreats, the crust rebounds relatively slowly, and the ocean can sometimes flow in to cover the sunken land before it rises again. Ongoing crustal rebound (i.e. the crust rising back up after the ice sheet retreats) causes the sea level to appear to become lower through time. Because of this, the Champlain Sea retreated as the land rebounded, but dominated the landscape of eastern Ontario and southern Quebec from 11 800 to 8800 BP (Gilbert 1994).

The Champlain Sea reached its maximum northwesterly extent at 11 ka BP, between Pembroke and Mattawa, Ontario (Gilbert 1994). The “Younger Dryas” climatic event (11-10 ka BP) also occurred at this time, as a retreating lobe of the Laurentide Ice Sheet allowed glacial Lake Agassiz, a large proglacial lake in Manitoba, to drain into the Gulf of St. Lawrence. This drainage of fresh water into the ocean is thought to have caused a near-shutdown of the North Atlantic Deep Water (NADW) circulation due to the stabilizing effect of fresh water over the denser salt water (Broecker et al. 1989). Since the NADW is responsible for transporting a massive amount of tropical warm water and air from the equatorial to the northern Atlantic, and defines the relatively mild climate of much of western Europe, this massive drainage event had catastrophic effects on global climate. It is recorded throughout the Northern Hemisphere as an atypically cold period during the overall warming that followed peak glacial conditions. By 10.6 ka BP, however, the climate of eastern Ontario had likely changed from “cold” to “cool and dry,” and the more cold-adapted vegetation had migrated northward (Gilbert 1994).

A major climatic and ecological change occurred at 10.5 ka BP, signaling the end of the Wisconsinan glacial stage, which seems to have occurred with remarkable abruptness and regional synchronicity (Ogden 1967). Ice retreat opened a channel from Georgian Bay to the Ottawa River through Lake Nipissing at around 10.1 ka BP, and all drainage from the Huron and Michigan basins flowed through this outlet for several thousand years (Gilbert 1994). Due to ongoing orbital changes, the amount of incident solar energy striking this region around 9000 BP was higher than present in the summer, and lower in the winter (Delcourt and Delcourt 1980), and the climate of northeastern North America was about 2ºC warmer than that of today (Fuller 1997). Isostatic rebound (the rebound of the Earth’s crust following the retreat of heavy ice sheets) in eastern Ontario occurred at a rate of 20 mm/annum at 8000 BP, but by 6000 BP had slowed to 10 mm/anum (Gilbert 1994). Isostatic rebound continues in the region to this day.

At present, the Laurentian Great Lakes basin includes (from upstream to downstream) Lake Superior, Lake Huron, Lake Michigan, Lake Erie and Lake Ontario, as well as many smaller lakes and rivers, all draining via the St. Lawrence River to the Atlantic Ocean. Lake depth can likely be explained by glacial scouring in each basin. Lake Superior, on the Canadian Shield, represents the deepest and most extensive of the Great Lakes, and Erie, the smallest and shallowest.

Effects on the Physical Landscape
Repeated cycles of glaciation have left their mark on the physical landscape of eastern Ontario. Included among the most obvious artifacts of glaciation are distinctive landforms that include drumlin fields, eskers, rattails and other marks in the bedrock, as well deposits of glacial sediment, sometimes tens of metres thick. Drumlins are an example of a large-scale landscape feature caused by glacial erosion: these are streamlined landforms with widths of up to 0.5 km and lengths in excess of 5 km, appearing as a ridge with parallel furrows on either side (Shaw and Sharpe 1987). The direction of the ridge indicates the direction of glacial motion, often with a crescent scour at the upstream end, which tends to fill with water, producing small lakes. Drumlins, like other, smaller scours in the bedrock, are believed to have been caused by either glacial erosion of pre-existing sediment and rock, or alternatively, by deposition of sediment by turbulent meltwater flow in the cavity beneath the ice (Shaw and Sharpe 1987), and perhaps equally by each of these forces in different regions. In Eastern Ontario, one prominent drumlin field can be found near Peterborough, the dominant morphology of which indicates north-northwestward ice or meltwater flow.

Eskers are another prominent and large-scale glacial landform that can be found in Eastern Ontario. These winding ridges with steeply sloping sides represent the remains of subglacial meltwater channels, which left thick sediment deposits that accumulated as the channel thawed upwards into the ice. When the ice sheets melted, the former stream beds became ridges on an otherwise scoured landscape. Eskers can stretch for several kilometres. Two well-known eskers in Eastern Ontario are the Greenborough Esker, in the Ottawa Valley, and the Tweed Esker, north of Prince Edward County.

Smaller glacial landforms can vary in scale from metres down to centimetres (Glasser and Bennett 2004). Some of the more obvious of these include “rattails,” which are essentially centimetre- or metre-scale drumlins carved in bedrock. Like (at least some) drumlins, rattails form when an upstream obstacle causes downstream turbulence to erode a channel in the rock face, which tapers off to a thin “tail.” The upstream end is usually marked by a small round or crescent scour, and the direction of the “tail” indicates the direction of meltwater flow beneath the glacier. Small features such as these are apparent throughout the formerly glaciated regions of North America. Numerous rattails are found on the Queen’s University Biological Station property, including at the easily-accessible QUBS Point, just west of Chaffey’s Lock, Ontario.

Perhaps one of the most important landscape effects that recent ice sheets have had on Eastern Ontario is in depositing glacial sediment. In areas west of Trenton, deposits up to tens of metres thick cover the limestone bedrock, while in most other areas, the sediment deposits are much thinner, perhaps one to two metres thick. Thicker “pockets” of glacial sediment do exist throughout the Canadian Shield landscape, as well as in southeastern Ontario (Gilbert 1994).

In a more general sense, glacial scouring in Eastern Ontario contributed to the topographical variation that defines the landscape, with softer rock being scoured more deeply. These deeply-scoured pockets became today’s lake beds, many of which are quite deep. Furthermore, such scouring seems to have resulted in the “poorly organized” drainage patterns in the region of the Queen’s University Biological Station in the Rideau Lakes system: lakes and wetlands are interconnected, but there are few large rivers, which represents a marked difference from the landscape as nearby as Kingston, Ontario, and the shores of Lake Ontario and the St. Lawrence and Ottawa River.

Impacts on the Evolution of Diversity and Present-day Species Distributions
Pleistocene glaciations have profoundly influenced species distributions and biotic community composition in Canada, which is perhaps obvious given the fact that mere millennia before present the majority of the country was covered by ice. Aside from this, there is much discussion regarding the importance of repeated glaciations and concomitant range fragmentation has been in the generation of new species, the assembly and similarity of biotic communities past and present, and the paths of recolonization into present-day Canada. Below we discuss each of these in turn.

The present-day distributions of some sister species (species that share a common ancestor with each other and no other taxon) originally suggested to biologists that many Northern Hemisphere species trace their origins to fragmentation of geographical distributions by repeated advance and retreat of glacial ice sheets. Much of our understanding of Pleistocene speciation and diversification comes from birds. For example, Robert Mengel in his classic 1964 monograph on wood warblers asserted that the present-day distributions of some wood warbler species (family: Parulidae) reflected separation of the ancestral range(s) of antecedent species with subsequent diversification among geographically separated populations. Thus, some believed that the Pleistocene was an “engine” of diversification and speciation with a large impact on contemporary diversity. As one example, Mengel cited the black-throated green warbler species complex, comprised of five species, which he asserted arose from the repeated glaciations of the late Pleistocene: one ranging broadly in boreal forest of Canada and the USA (also summer breeding resident at QUBS) – the black-throated green warbler (Dendroica virens); three distributed in western North America: hermit warbler (D. occidentalis), Townsend’s warbler (D. townsendi), black-throated grey warbler (D. nigrescens); and one species a rare endemic of central Texas – the golden-cheeked warbler (D. chrysoparia). DNA evidence suggests that the origins of species of the black-throated green warbler complex are somewhat different than envisioned by Mengel (1964), with some western species derived from an eastern ancestor, with others arising in situ in western North America because of isolation in the intermontane regions (Bermingham et al., 1992). Thus while shifting distributions shaped by the waxing and waning of glaciers and vegetations undoubtedly shaped evolutionary trajectories, it is clearly not as simply as repeated east-west splits of taxa.

One of the great debates in ecology extending back to the very origins of the discipline (Gleason, 1924 versus Clements, 1904) has centered on the nature of biotic communities (Stewart 2009). Are they specific assemblages of species that recur over evolutionary time because of necessary associations among similar species, and natural successional processes (Clements 1904)? Or are communities unique assemblages of species that arise contingently in a particular place and time because of species’ individualistic responses to a particular transitory constellation of environmental factors (Gleason 1924)? For example, in Eastern Ontario, including the Queen’s University Biological Station, many forests comprise a mixture of conifers (eastern hemlock, Tsuga canadensis, red pine, Pinus resinosa, white pine, P. strobus) and deciduous species, including sugar maple (Acer saccharum), red maple (A. rubrum), red oak (Quercus rubra), and American beech (Fagus grandifolia) (see Flader 1983) – is this forest community consistently found throughout the Quaternary and particularly the Holocene? Palynological (fossil pollen) and climate simulations appear quite clear on these matters. Unique forest communities arise as a result of myriad plant species individually responding to climatic conditions occurring at one particular juncture in history (Williams and Jackson 2007). For example, Williams et al. (2001) found many North American plant assemblages inferred from pollen records dating between 17 and 9 ka BP lack modern analogs (i.e. they comprise tree and other plant species that are not found together today). Climate simulations suggest that these “no-analog communities” resulted from late-glacial climates with colder-than-present winters, warmer-than-present summers, and lower-than-present precipitation. Moreover, some authors make the provocative claim that megafaunal (species > 100 kgs) population collapses and subsequent release from herbivore pressure some 14.8 to 137 ka BP produced marked increase in hardwood species abundance and incidence of fire also dramatically altering forest plant species composition (Gill et al. 2009).

The northward retreat of glaciers has been used as the basis for the hypothesis that established plant communities dispersed northward from their peak-glacial habitats. Pollen records and fossil plant remains provide a means of testing this idea. As discussed above, the retreat of glaciers was accompanied by the development of a sequence of proglacial lakes. The size and location of these lakes undoubtedly influenced the rate at which plant species recolonized the region (Terasmae 1981). In particular, the low level of Lake Erie throughout the Two Creeks Interstadial (12.5-11.8 ka BP) permitted founder populations of cold-adapted tree species (e.g. black spruce) to colonize southern Ontario, as indicated by pollen records. One thousand years after this, the Ontario climate was from cold to cool and dry, and in Eastern Ontario spruce could be found intermingled with willow, pine, wormwood and ragweed (Gilbert 1994). By 10,000 BP, spruce were gone from southern Ontario, and at 9500 BP, pine appears to have increased in abundance (Kapp 1986). By 7 500 BP, the “Great Lakes Forest” was established, with a mix of coniferous and deciduous trees established in response to warmer conditions. Hemlock dominated, with pine declining, and basswood (Tilia) and hickory (Carya) increasing (Gilbert 1994). Since then, there appear to have been two major influences on forest composition: disease (such as the hypothesized hemlock disease that persisted from 4 700-3 000 BP, and the more recent chestnut blight and Dutch elm disease), and deforestation by colonists.  The summation of this evidence suggests that the post-glacial dispersal of plant communities is complex, and dependent on several factors besides climate. It also serves to reinforce the notion that forest communities with no modern analog have existed.

Literature & Further Reading

  1. Bermingham, E., S. Rohwer, S. Freeman and C. Wood.1992. Vicariance biogeography in the Pleistocene and speciation in North American wood warblers. A test of Mengel’s model. Proc. Natl. Acad. Sci. USA 89: 6624-6628.
  2. Clements, F.E. 1904. Plant Succession: An Analysis of the Development of Vegetation. Carnegie Institute, Washington.
  3. Flader, S.L. (ed.) 1983. The Great Lakes Forest: An Environmental and Social History. Univ. Minnesota Press.
  4. Fuller, J.L. 1997. Holocene forest dynamics in southern Ontario, Canada: fine-resolution pollen data. Can. J. Bot. 75: 1714-1727.
  5. Fullerton, D.S. 1986. Stratigraphy and correlation of glacial deposits from Indiana to New York and New Jersey. Quaternary Science Reviews 5: 23-37.
  6. Gibbard, P.L., M.J. Head, M.J.C. Walker and the Subcommission on Quaternary Stratigraphy. 2010. Formal ratification of the Quaternary System/Period and the Pleistocene Series/Epoch with a base at 2.58 Ma. J. Quaternary Science. 25: 96-102.
  7. Gilbert, R. 1994. A field guide to the glacial and postglacial landscape of southeastern Ontario and part of Quebec. Geological Survey of Canada Bulletin, bulletin 453.
  8. Gildor, H. and E. Tziperman. 2000. Sea ice as the glacial cycles’ climate switch: Role of seasonal and orbital forcing. Paleoceanography 15: 605–615.
  9. Gill, J.L., J.W. Williams, S.T. Jackson, K.B. Lininger, and G.S. Robinson. 2009. Pleistocene megafaunal collapse, novel plant communities, and enhanced fire regimes in North America. Science 326: 1100-1103.
  10. Glasser, N.F. And M.R. Bennett. 2004. Glacial erosional landforms: origins and significance for paleoglaciology. Progress in Physical Geography 28: 43-75.
  11. Gleason, H.A. 1926. The individualistic concept of the plant association. Bulletin of the Torrey Botanical Club 53: 7–26.
  12. Gradstein, F.M., J.G. Ogg, A.G. Smith, et al. 2004. A Geologic Time Scale. Cambridge Univ. Press.
  13. Grimley, D.A. 2000. Glacial and nonglacial sediment contributions to Wisconsin episode loess in the central United States. Geological Society of America Bulletin 112: 1475-1495.
  14. Hays, J.D., J. Imbrie and N.J. Shackleton. 1976. Variations in the Earth’s orbit: Pacemaker of the ice ages. Science 194: 1121–1132.
  15. Kapp, R.O. 1986. Late-glacial pollen and macrofossils associated with the Rappuhn Mastodont (Lapeer County, Michigan). The American Midland Naturalist 116: 368-377.
  16. Karrow, P.F. 1974. Till Stratigraphy of Parts of Southwestern Ontario. Geological Society of America Bulletin 85: 761-768.
  17. Larson, G. and R. Schaetzl. 2001. Origin and Evolution of the Great Lakes. Journal of Great Lakes Research 297: 518-547.
  18. Mengel, R.M. 1964. The probable history of species formation in some northern wood warblers (Parulidae). Living Bird 3: 9–43.
  19. Milankovitch, M. 1941. Canon of insolation and the ice-age problem 
(Kanon der Erdbestrahlung und seine Anwendung auf das Eiszeitenproblem) Belgrade, 1941. Translated from German. Published in 1969 by Israel Program for Scientific Translations.
  20. Redfern, R. 2001. Origins: The Evolution of Continents, Oceans and Life. University of Oklahoma Press.
  21. Shaw, J. and D.R. Sharpe. 1987. Drumlin formation by subglacial meltwater erosion. Can. J. Earth Sci. 24: 2316-2322.
  22. Sosdian, S. and Y. Rosenthal. 2009. Deep-sea temperature and ice volume changes across the Pliocene-Pleistocene climate transitions. Science 325: 306–310.
  23. Stewart, J.R. 2009. The evolutionary consequence of the individualistic response to climate change. J. Evol. Biol. 22: 2363-2375.
  24. Terasmae, J. 1980. Some problems of late Wisconsinan history and geochronology in southeastern Ontario. Can. J. Earth Sci. 17: 361-381.
  25. Terasmae, J. 1981. Late Wisconsinan deglaciation and migration of spruce into southeastern Ontario, Canada. In: Geobotany II, ed. R. Romans (New York: Plenum Press) 75-90.
  26. Weir, J.T. and D. Schluter. 2004. Ice sheets promote speciation in boreal birds. Proc. Roc. Soc. London Ser B Biol. Sci. 271: 1881–1887.
  27. Williams, J.W. 2001. Dissimilarity analyses of late-Quaternary vegetation and climate in eastern North America. Ecology 82: 3346-3362.
  28. Williams J.W. and S.T. Jackson. 2007. Novel climates, no-analog communities, and ecological surprises. Front. Ecol. Environ. 5: 475–482.

Reviewed by: Paul Handford (Univ. Western Ontario) and Scott Lamoureux (Queen’s Univ.)

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