Water bodies, ancestral to the present lakes including Lake Huron, first appeared in the southern Great Lakes basin about 15,500 14C years (18,800 cal years) BP during the oscillatory northward retreat of the last (Laurentide) ice sheet from its maximum position south of the Great Lakes watershed. Glacial lakes, impounded by a retreating ice margin on their northern shores, were continuously present after 13,000 14C (15,340 cal) BP for 3000 14C (3900 cal) years. Drainage routings varied in time through the Erie and Michigan basins to the Mississippi River system, a probable source for colonizing aquatic organisms, then to the Ontario basin, and finally northeastward to the Ottawa River valley via the isostatically-depressed North Bay outlet by 10,000 14C (11,470 cal) BP. Water levels were generally low between 10,000 and 7500 14C (11,470 and 8300 cal) BP and may have risen several tens of metres for short periods due to overflow of meltwater from upstream subglacial reservoirs or from glacial lakes impounded by residual ice in the Hudson Bay watershed.
About 8000 14C (8890 cal) BP glacial runoff bypassed the Great Lakes, and Huron basin waters descended into hydrologic closure under the influence of the early Holocene dry climate. With increasing precipitation and water supply about 7500 14C (8300 cal) BP the Huron water body again overflowed its North Bay outlet. Differential isostatic uplift (fastest to the north-northeast) raised this outlet and lake level relative to the rest of the basin. The lake finally overflowed southern outlets at Chicago and Port Huron-Sarnia by 5000 14C (5760 cal) BP (during the Nipissing highstand). Enhanced erosion of the latter outlet and continued differential uplift of the basin led to the present configuration of Lake Huron and Georgian Bay.
Lake Huron and Georgian Bay are located centrally among the Laurentian Great Lakes in North America (Fig. 1). This group of large lakes constitutes the headwaters of the St. Lawrence River that drains eastward to Atlantic Ocean. The watershed of Lake Huron and Georgian Bay comprises an area of about 194,000 km2 with 59,600 km2 being lake surface. At its northwestern extremity Lake Huron joins Lake Michigan over a 42 m deep sill via the Straits of Mackinac, and the two lakes, including Green Bay and Georgian Bay, form a confluent water surface at 176 m above sea level. This paper is a short review of the history of lake development in the Huron basin.
The maximum water depths in Lake Huron and Georgian Bay are 229 m and 168 m, respectively, and their combined water volume is 3540 km3 (The Great Lakes Environmental Atlas, 1995). The configurations of the lake shoreline and of the bathymetric basins (Fig. 1) are mainly controlled by the topography of the bedrock surface, which is the result of differential erosion of the underlying bedrock strata. During the previous 50–60 million years of river erosion followed by 2–3 million years of erosion by multiple glaciations, the softer shale and salt formations have been excavated into basins separated by upstanding remnants of the harder layers of limestone and dolomite (Hough, 1958; Larson, and Schaetzl, 2001). These sedimentary rocks rest on harder metamorphic rocks of the Precambrian Shield and dip gently westward and southward into the Michigan structural basin between lakes Huron and Michigan. As a result, the erosion-resistant dolomite formations in Huron basin form east- to north-facing escarpments. The most prominent of these is the Niagara escarpment that forms the peninsula between Lake Huron and Georgian Bay. The underwater sediment distribution and substrates of lakebed habitat, as shown in Thomas et al. (1973), Dobson et al. (1995) and McCarthy et al. (2007) for example, are a mosaic of bedrock outcrop in places, and deposits of the subsequent glacial and lacustrine paleo-environments as described below.
Although lakes undoubtedly existed during previous interglacial periods, the sequence of water bodies leading to the present Lake Huron and Georgian Bay began about 15,500 BP (radiocarbon years before present) or 18,800 cal BP1 following retreat of ice from the last glacial maximum [21,000-18,000 (25,270-21,290 cal) BP] (Fig. 2). This history began in the Lake Erie basin, the southernmost, first deglaciated, and an important entry route for flora and fauna into the glacial Great Lakes. As the Erie ice lobe shrank, glacial lakes expanded northward into the southern Huron basin against the retreating ice dam of the Huron lobe (Eschman and Karrow, 1985; Lewis et al., 1994). Biota had ready access to the Erie basin from the non-glaciated area in the Ohio Valley to the south. A glimpse of that life has recently been documented at Leamington, Essex County, Ontario, dated 13,200 (15,620 cal) BP (Karrow et al., 2007). The southern Lake Michigan basin provided a similar entrance route from the Mississippi and Illinois valleys for biota before 13,000 (15,340 cal) BP. Another route for migration of aquatic organisms may have been from glacial Lake Wisconsin (Fig. 2), which had existed throughout the glacial maximum (Clayton and Attig, 1989), into Lake Oshkosh via overflow from Wisconsin River, and then into Green Bay of Lake Michigan during subsequent ice retreat.
Most of the early glacial lakes, centred in the Erie basin with their northern shores constrained by ice in the southern Huron basin, had initial overflow southwestward down the Wabash Valley to the Ohio and Mississippi valleys (Calkin and Feenstra, 1985). As northward retreat of the Huron lobe opened lower outlets, drainage was westward through the Grand Valley of Michigan to the Lake Michigan basin and then southward along the Illinois and Mississippi valleys. In Saginaw Bay, a smaller Lake Saginaw developed along this drainage route with fluctuating levels as ice retreated and re-advanced about 13,000 (15,340 cal) BP (Fig. 2). The separate Lake Saginaw ended with ice retreat after 13,000 (15,340 cal) BP as its lowering waters joined those in the main Lake Huron basin.
Concurrently in the Lake Michigan basin (Hansel et al., 1985), Lake Glenwood followed the retreating Michigan lobe northward while draining southward by the Illinois River valley to the Mississippi River. Lake Glenwood ended as ice retreat allowed drainage northward by Mackinac Straits and adjacent lowlands into the Huron basin (and eastward) about 13,400 (15,900 cal) BP. The Port Huron ice advance about 13,000 (15,340 cal) BP then restored Lake Glenwood, followed by Lake Calumet, returning outflow temporarily into the Illinois and Mississippi valleys (Fig. 2). Michigan lobe ice then retreated in the Two Creeks Phase, again sufficient to allow Mackinac Straits eastward outflow at 11,900 (13,760 cal) BP (Kaiser, 1994). The subsequent rapid Two Rivers advance of the Michigan lobe again forced southward drainage from restored Lake Calumet. Declining waters with northward ice retreat then allowed the final merging of Michigan and Huron basin water bodies. After two decades of controversy over modelling of crustal tilting, new dates confirm the presence of Algonquin waters in the southern Lake Michigan basin by about 10,500 (12,500 cal) BP (Capps et al., 2007).
Crustal tilting reflects uplift and recovery of Earth's crust from the weight of the previous ice load, and is an ongoing effect of the last glaciation. The uplift is differential in the direction of ice retreat and past thicker ice, with northern areas uplifting faster than southern areas; the northern Huron basin has been uplifted over 200 m more than its southern extremity since 10,600 (12,600 cal) BP (Lewis et al., 2007a). The upward velocity decreases exponentially with time, so the effect was most rapid during early stages of deglaciation, but still it has dominated the evolution of postglacial lakes in the Huron basin.
Erie basin glacial lakes came to an abrupt end when the Erie ice lobe thinned and retreated enough to allow catastrophic drainage eastward across New York to the Hudson Valley and to the Atlantic Ocean (e.g. Hand and Muller, 1972). Huron basin glacial lakes then became isolated from the Erie basin. It is believed there was an early Algonquin lake that formed in the Huron basin that drained to Erie basin, but its shorelines and deposits have not been separately identified. A separate local Lake Schomberg formed in south-central Ontario between the north side of the Oak Ridges moraine and the retreating ice about this time (Fig. 2), which coalesced with Algonquin waters.
Subsequent ice retreat switched Huron basin drainage eastward past Kirkfield (Fig. 1) down the Trent River valley of Ontario into glacial Lake Iroquois of the Lake Ontario basin and ended Early Lake Algonquin to form the Kirkfield Algonquin low water phase, with extensive exposure of former lake bottom around southern Georgian Bay, Lake Huron, and Lake Michigan basins. While the ice margin continued its northward retreat across the area east of Georgian Bay, isostatic crustal tilting raised the Kirkfield outlet sill to Trent River valley (Fig. 3), causing a southward transgression of southern Huron basin lake waters that is well documented with radiocarbon dates and fossil biota at numerous sites in southern Ontario (Karrow et al., 1975, 1995). This transgression spanned at least the period 11,300 to 10,500 (13,190 to 12,500 cal) BP as the rising waters reached twin southern outlets at Chicago and Port Huron to create Main Lake Algonquin (Fig. 3), the largest of the glacial Great Lakes and a focus of much interest by archeologists tracking evidence of coastal occupation by the earliest humans (Paleoindian) (Storck, 2004). Lake Algonquin was the first to form one confluent water body through the Michigan, Huron, and Georgian Bay basins, and for a short time as the retreating ice allowed, to extend into the southern Superior basin as well.
Paleoclimatic theory (COHMAP, 1988) suggests the Lake Huron region was dominated during the last glaciation and deglaciation by a glacial atmospheric anticyclone circulation. Cold, dry air was circulated clockwise over the ice sheet, and extended up to about 100 km south of the ice margin, affecting glacial lakes and their surrounding terrain. Strong winds of this circulation are indicated, for example, by sand spits that prograded from east to west in Lake Algonquin in the upper Michigan peninsula west of the present northern Lake Huron (Krist and Schaetzl, 2001).
Nearshore Lake Algonquin sediments in southern Ontario, as represented by estuarine silts underlying valley terraces (Karrow, 1986) and by lagoon sediments behind or under baymouth bars, have yielded an abundant fossil flora and fauna, clearly showing that diverse life flourished in at least the later glacial lakes – i.e. they were not sterile, barren environments. Detailed studies of pollen, plant macrofossils, ostracodes, and molluscs in Lake Algonquin sediments aided correlation and age estimation for the enclosing sediments, and provided data for interpretation of past environments (Karrow et al., 1975, 1995; Miller et al., 1985; Fitzgerald, 1985; Silis, 1999). Changes in the distributions of species through time relate to changes in drainage and varying linkages between basins as the Great Lakes evolved.
Eventually ice retreat unblocked a series of outlets for Lake Algonquin south of North Bay, shortly before 10,000 (11,470 cal) BP, with resulting catastrophic floods eastward into the Ottawa Valley (Harrison, 1972; Karrow, 2006). This major event diverted all drainage of the upper Great Lakes through the Lake Nipissing lowland past North Bay, Ontario, for most of the subsequent 5000 (5760 cal) years, and resulted in the lowering of water levels to create shrunken low–water phases early Lake Chippewa (Michigan basin), early Lake Stanley (Huron), and early Lake Hough (Georgian Bay). About half the present areas of these lakes dried up, became forested, and were dissected by lengthened streams (Karrow et al., 2007).
Late- and post-glacial paleoclimate
Changes in climate during the deglacial to present period have been derived from the δ18O isotopic composition of cellulose in fossil wood, and in organic matter of lake sediments, and from changes in forest composition evidenced by changes in pollen assemblages recorded in bog and lake sediments. Calibration of the first proxy (Edwards et al., 1996), and statistical transfer function analyses of the second proxy (Bartlein and Webb III, 1985; McAndrews and Campbell, 1993) indicate a sequence of climate during the past 12,000 (13,900 cal) years from cold and dry, through warm and dry, to warm and moist before arriving at the present climate (Fig. 4). This climatic sequence is traditionally interpreted in terms of changes in the relative dominance of different air masses − cold and dry Arctic air, warm and dry Pacific air, and warm and moist Gulf of Mexico air (Bryson and Hare, 1974). Shifts in the mean extent of these air masses could result in significant shifts in mean temperature and water supply to the lake basins (Yu et al., 1997). During the early to middle Holocene major controls of climatic change were retreat and downwasting of the Laurentide Ice Sheet, higher-than-present summer insolation, and possibly northeastward movements of the Atlantic Bermuda high pressure zone which enhanced mid-continent drought. These influences curtailed the glacial anticyclone circulation progressively from west to east and led to increased aridity by about 8100 (9050 cal) BP in the mid-continent region (Shuman et al., 2002), probably decreasing water supply to the Lake Huron basin. Wetter conditions prevailed in the middle and late Holocene with demise of the ice sheet and decline of the insolation anomaly (Shuman et al., 2002). In addition to changing airmass circulation and boundary conditions, other atmospheric processes such as moisture flux, soil moisture recycling, and vertical movements (uplift and subsidence), may have influenced Holocene climatic trends in mid-continent North America (Shinker et al., 2006).
The proxy paleoclimatic record suggests the Lake Huron region was dominated during the deglacial period by cold, dry, anticyclonic winds, followed by frequent southward incursions of cold, dry Arctic air as the ice sheet diminished and melted down in the Hudson Bay area (Yu and Wright, 2001). The cold, dry regime was replaced by a warm-dry regime, well-documented in numerous mid-west pollen records (Shuman et al., 2002). Then about 7500 (8300 cal) years BP, a warmer, wetter regime with moisture advected from the Gulf of Mexico (Shuman et al., 2002) began to affect the Lake Huron region to the present, making significant contributions to the lake's water supply. A warmer-than-present (hypsithermal) climate occurred after about 7500 (8300 cal) years BP (Edwards et al., 1996), possibly interrupted by a short period (century?) of cooling by the North Atlantic 8.2 cold event (Alley and Ágústsdóttir, 2005). This warmer period is expressed in mid-Holocene silty clay sediments of the Huron basin by the more frequent presence of black bands containing iron sulphide. The black bands were deposited during periods of likely anoxia near the sediment-water interface (Odegaard et al., 2003). Small shifts in atmospheric processes and circulation probably caused small variations in water supply and lake level into the present while much larger changes were imposed by abrupt changes in throughflow and the steady influence of uplift.
For 2500 (3170 cal) years after 10,000 (11,470 cal) BP approximately, during the transition from late glacial to postglacial conditions, lake levels fluctuated dramatically through a possible range up to 60 m (Lewis and Anderson, 1989; Lewis et al., 1994, 2007a, 2007b) (Fig. 5). Water supplies were augmented by inflows from upstream Lake Agassiz, which were sometimes large and catastrophic (Teller et al., 2002), or from subglacial outburst floods (Breckenridge et al., 2004; Breckenridge, 2007). Water levels rose 40 to 60 m above the North Bay outlet about three times for short periods to the Mattawa highstands (Fig. 5), which are thought to have been held up by constriction of the large outflows in the Mattawa and Ottawa valleys. Alternations between glacial and nonglacial sources of water during this period are known from oxygen isotopic composition of benthic ostracode valves (Rea et al., 1994a, 1994b; Dettman et al., 1995; Macdonald and Longstaffe, 2008). Between highstands, water levels fell to the North Bay outlet and probably below, indicated by tree stumps in growth position now submerged (Blasco, 2001), and by wave erosion surfaces now in deep water recorded in the sedimentary record (Hough, 1962; Moore et al., 1994; Rea et al., 1994a, 1994b; Lewis et al., 2007b). These lakes likely fell 25 to 30 m below the level of the overflow outlet at North Bay as a result of excessive evaporation in the dry early Holocene climate (Fig. 6). The last lowstand ended about 7500 (8300 cal) BP with the increasing precipitation and water supply brought by the more frequent presence of a warm-moist regime over the Huron basin as mentioned earlier. From this time forward, the Huron basin lakes were hydrologically open, overflowing water bodies, as at present.
With the level of the hydrologically open lake controlled by overflow at North Bay, differential uplift continuously raised this outlet and lake level relative to other parts of the basin from about 7500 to 5000 (8300 to 5760 cal) BP (Fig. 5). River valleys were progressively flooded and infilled with sediment (Karrow et al., 2007). By 5000 (5760 cal) BP the Huron basin water level, known as the Nipissing Great Lakes, had flooded all basins of the upper Great Lakes to levels higher than present, and had begun overflowing to the south at Chicago and Port Huron-Sarnia.
Studies of fossils in sediments of the postglacial Nipissing Great Lakes provide a marked contrast with those of glacial Lake Algonquin, with Nipissing floras and faunas being most like the present. Nipissing time was within the hypsithermal interval, when climate was warmer than the present, with deciduous forest, in contrast to Algonquin time which was during cooler climate associated with boreal conifer forest. In addition to known differences in plants, it is possible to distinguish the molluscs in the sediments of the two lakes by the presence of different species (Miller et al., 1985). Studies of fossil assemblages in the lakes spanning the time from deglaciation to the present show a progressive enrichment, with increasing diversity as the climate warmed and soils matured.
Glacial rebound continued to raise the Huron basin differentially, but with the lake outlets in the southern sector, the trend in water level changed from ascent to descent relative to land. The Chicago outlet was founded on bedrock but the Port Huron outlet was based on glacial sediments that became incised enough so that all overflow was discharged into the St. Clair-Detroit river system to Lake Erie by 4200 (4740 cal) BP (Larsen, 1985). A relatively rapid lowering of the water surface occurred around 3200 (3420 cal) BP when the Algoma shore was abandoned in the Huron and Michigan basins (Larsen, 1985), and a lower St. Clair River delta surface began forming (Thomas et al., 2006). Rapid outlet erosion and reduction in lake level may have resulted from a steepening of the outlet river gradient due to a 3-m drop in the downstream Erie basin water level at about the same time (Pengelly et al., 1997). The combined lowering from outlet erosion and crustal tilting lowered water levels to form lakes Michigan, Huron, and Georgian Bay of today; it also allowed preservation of numerous recessional beach complexes between the Nipissing shoreline and that of the present. Analysis of these complexes has provided detailed hydrographs far back beyond the historical record, revealing several superposed cycles of water level changes related to variations in climate and water supply, which continue today (Fraser et al., 1990; Thompson, 1992; Baedke and Thompson, 2000).
The direct ancestral lakes of Lake Huron and Georgian Bay date only since the last glaciation, although the bathymetric basins are the result of differential erosion of bedrock by previous rivers and glaciations. Retreat of the last ice cover after about 15,500 (18,800 cal) BP led to a sequence of ice-dammed glacial lakes that overflowed southward first from the Erie and then Michigan basins into tributaries of the Mississippi River. Fish and other aquatic organisms could have used these drainage routes to migrate from Mississippi River and possibly other refugia to colonize the Erie, Michigan, and subsequently the Huron basin lakes.
The retreat was punctuated with ice re-advances, and only after 13,000 (15,340 cal) BP did the southern Huron basin become permanently free of a glacial cover. By 11,000-10,500 (12,920-12,500 cal) BP the Huron, Michigan and part of Superior basins were inundated by the largest of the glacial Great Lakes, Main Lake Algonquin. With ice retreat from the isostatically-depressed lowland northeast of Georgian Bay, the upper Great Lakes basins drained over an outlet sill at North Bay to the Ottawa Valley for most of 5000 (5710 cal) years. From 10,000 to 7500 (11,470 to 8300 cal) BP low lake levels were present except for about three short periods when overflow from upstream glacial lakes or subglacial reservoirs flooded into the upper Great Lakes basins, raising levels by 40 to 60 m above the North Bay outlet (Mattawa highstands). The last of the lowstands (between 8000 and 7500 (8890 and 8300 cal) BP) fell below this outlet by 25 to 30 m into hydrologic closure because of excessive evaporation in the early Holocene dry climate. The extreme variation of lake level under climate forcing at this time provides an opportunity to better understand the sensitivity of the Huron lake system to climate change, useful for modelling the future effects of global warming on Lake Huron. The shifts in shoreline positions during this transition from glacial to postglacial lakes are a focus for archeological research in locating evidence of early humans.
By 7500 (8300 cal) BP, moist air masses became more prevalent, and the increased water supply raised lakes in the Huron basin to a hydrologically open, overflowing condition that continued to the present. For the next 2500 years lake levels rose relative to land under control of the isostatically rebounding North Bay outlet, culminating in the Nipissing Great Lakes phase about 5000 (5760 cal) BP, as outflow was transferred from North Bay to southern outlets at Chicago and Port Huron-Sarnia. The Port Huron outlet soon captured the Huron basin drainage entirely, where it remains at present. The previous 2500-year period of rising lakes was followed by a 4000- to 5000-year period of falling relative lake level as a result of outlet downcutting and continued rebound of the basin. Today, many of the deeper bathymetric basins are accumulating fine-grained sediment while parts of the shoreline accumulate sand in beaches. Between these zones, sediment units or bedrock that outcrop on the lakebed reflect various stages in the paleo-environmental history of the Huron basin.
We appreciate discussion with L. Clayton (Wisconsin Geological and Natural History Survey), and reviews of the manuscript by G.R. Brooks and R. Parrott for the Geological Survey of Canada, as well as reviews by A. Breckenridge and an anonymous reviewer for the journal. Figures were prepared by P. Melbourne of the Electronic Publishing Group at Geological Survey of Canada Atlantic. Preparation of this paper was supported by the Earth Sciences Climate Change program of Natural Resources Canada. This is Earth Sciences Sector (Geological Survey of Canada) Contribution 20060698.
cal BP = calibrated 14C age in calendar or solar years before present (BP) where present is AD 1950. Hereafter, dates are reported, for example as 7500 (8300 cal) BP, meaning 7500 14C years BP or 8300 calibrated years BP.