Glacial Isostatic Adjustment of the Laurentian Great Lakes Basin: Using the Empirical Record of Strandline Deformation for Reconstruction of Early Holocene Paleo-Lakes and Discovery of a Hydrologically Closed Phase

Michael Lewis, C.F.; Blasco, Steve; Gareau, Pierre

doi:https://doi.org/10.7202/014754ar

Introduction

The Great Lakes of North America consist of five major basins whose water surfaces comprise 32% of a total watershed area of 766 000 km² (The Great Lakes Environmental Atlas, 1995) (Fig. 1). The watershed forms the headwaters of the St. Lawrence River that drains to the Gulf of St. Lawrence and Atlantic Ocean. At the last glacial maximum, the watershed was completely covered by the southern flank of the Laurentide Ice Sheet about 18-21 ka BP (21.3-25.4 cal ka BP). Deglaciation of the basins occurred as the ice margin retreated generally in a northerly direction in a series of oscillations, first exposing the Erie basin about 15.5 ka BP (18.8 cal ka BP), and finally receding to the northern Superior basin about 9.5 ka BP (10.7 cal ka BP) (Table I; Dyke et al., 2003; Dyke, 2004). During this retreat, a series of proglacial lakes formed shorelines of different ages that are upwarped today towards the north-northeast in the direction of thicker and longer-lasting ice (Fig. 2). This deformation is the cumulative isostatic adjustment of Earth’s crust since formation of the shorelines as a result of removal of the former ice sheet load. This differential rebound proceeded throughout the period of ice retreat and postglacial time at decelerating rates of uplift. It is continuing today, as evidenced by tilting of the Great Lakes basins measured in long-term records of lake level gauges between southern and northern shores of the Great Lakes (CCGLBHHD, 1977; Tushingham, 1992; Mainville and Craymer, 2005).

Table I

**Sources of isobase and deglacial information**

The differential and continuing nature of rebound has raised indicators of former lake levels to differing elevations, and this makes their correlation difficult and uncertain, particularly for those that are at scattered locations within the basin. This difficulty is especially evident for lake-level indicators that are now below the surface of the present Great Lakes. To facilitate the reconstruction of original elevations of former lake-level indicators, a reference isostatic response surface with an age of 10.6 ka BP is constructed for the entire basin. This surface is used in conjunction with an exponential function to describe the vertical motion in the time domain of any location in the entire watershed of the Great Lakes. This motion is constrained by, and consistent with, the observed empirical data provided by well-known upwarped strandlines in the Great Lakes basins. With this reference response surface and the exponential expression for uplift, the original elevations of specific lake-level indicator sites and potential overflow outlets are reconstructed and compared. The paleogeography of individual basins and of the entire Great Lakes watershed may likewise be reconstructed for any desired age. Applications of the isostatic response surface and the exponential model of uplift are illustrated for: (1) digital reconstruction of Great Lakes paleogeography in a geographic information system (GIS) environment, and (2) the assessment of former lake-level indicators in relation to possible overflow outlets, leading to the conclusion that the Michigan, Huron and Georgian Bay basins, if not all the Great Lakes, were once hydrologically closed.

Figure 2

Maps of isobases of selected reference paleoshorelines in basins of the Great Lakes, showing the radiocarbon ages (ka BP) and trends of isobases with their present lowest and highest elevations. Sources of isobases given in Table I. C and D mark end points of a section to which isobases of the Iroquois, Algonquin and Nipissing lakes were projected for illustration in Figure 3. (A) Map of Laurentide maximum ice margin and isobases of lakes Wisconsin, Washburn, Minong, Algonquin, Glenwood II, Whittlesey and Iroquois. (B) Map of isobases for the Nipissing Great Lakes, and for an older phase of glacial Lake Wisconsin.

Cartes des isobases des anciennes lignes de rivage de référence de la région des Grands Lacs, où les âges au ¹⁴C (ka BP) et le soulèvement différentiel par rapport aux altitudes actuelles sont illustrées. Les sources bibliographiques sont présentées au tableau I. Les points C et D sont les limites du profil vertical projeté à la figure 3 des lacs Iroquois, Algonquin et Nippising. (A) Carte de l’extension glaciaire maximale de l’inlandsis laurentidien et des isobases des lacs Wisconsin, Washburn, Minong, Algonquin, Glenwood II, Whittlesey et Iroquois. (B) Carte des isobases de la phase Nipissing des Grands Lacs et pour une phase plus ancienne du lac Wisconsin.

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Uplift in the Time Domain

Isobases of Glacio-Isotatic Rebound

In the continental interior, isobases of glacio-isostatic rebound are usually defined by the elevations of a differentially uplifted shoreline of a former lake (Goldthwait, 1907, 1910; Leverett and Taylor, 1915; Hough, 1958; Walcott, 1972; Lewis and Anderson, 1989; Schaetzl et al., 2002). Sets of isobases selected for this study for basins within the Great Lakes watershed are illustrated in Figure 2A and 2B in which the lowest and highest elevations, trends of isobases, the name and uncalibrated radiocarbon age are shown for each lake. Although the Algonquin and Nipissing highstands were confluent in three of the basins about 10.6 ka BP (12.6 cal ka BP) and 5.0 ka BP (5.7 cal ka BP), respectively, most sets of isobases are confined to a single basin. These former, once-level lake surfaces are all warped upward in a north to northeasterly direction (Fig. 2A-B) as a result of differential glacio-isostatic recovery of Earth’s crust following deglaciation of the Laurentide Ice Sheet. The sources of isobase and related deglacial information for basins within and adjacent to the Great Lakes are listed in Table I.

Uplift as a Function of Time

Whereas isobases portray the present configuration of a rebounding surface, it is often desirable to describe the uplift of a given location through time, or to construct surfaces at intermediate times, especially when reference shorelines are widely spaced in time. Following Andrews (1970) and others, the exponential function is adopted to describe relative uplift (U_t) with time where time is expressed as age t (cal years BP) for a landscape previously loaded by an ice sheet. From Peltier (1994, 1998) we use:

where A and τ (tau) are parameters of the equation. A is a site-specific amplitude factor, and is evaluated as:

for a known relative uplift, age and relaxation time. Tau (τ) is the relaxation time or period in years for which decelerating uplift is reduced by 1/exp (1/2.7183 or 36.8%) in successive periods. It is evaluated by solving (Equation 1) at sites where U_t is known for at least two sets of isobases of different ages as shown below. As a first-order approximation, τ is assumed to be time invariant and similar in value throughout the Great Lakes region. Figure 3 shows the reasonably good fit of the exponential uplift curve to Great Lakes relative rebound data for the Nipissing, Algonquin and Iroquois shorelines between sites C and D for trial values of relaxation time between 3000 and 5000 years (see Fig. 2 for location and isobase values).

Evaluation of parameters (τ and A) in the relative uplift equation (1)

Relaxation time (τ) was determined on 20 specific short transects (Fig. 4) where relative uplifts, U₁, U₂, and their ages t₁, t₂, are known for two shorelines, for example the Algonquin and Nipissing isobases whose domains largely overlap one another. Then, on each transect:

As the amplitude factor is identical on a specific short transect, these equations were rearranged in terms of A and subtracted to yield a single equation which was solved for τ:

From Figure 4, a mean value of τ, rounded to 3700 ± 700 years, is used for computations of relative rebound in the Great Lakes basin. Similar values of 3400 years and 3500 ± 400 years, respectively, were obtained for the relaxation time of an exponential fit to relative sea-level changes in rapidly-uplifting James Bay south of Hudson Bay (Fig. 3A and 3C; Peltier, 1998), and for relaxation of glacial rebound in the Lake Winnipeg area, Manitoba (Lewis et al., 2000; Brooks et al., 2005).

The amplitude factor is then determined by evaluating Equation 2 using the known decay time τ:

Relative uplift (U_x) or shoreline slope (S_x) at other times

With τ and A known, relative uplift (U_x) or shoreline slope (S_x) for any given age t_x (cal years BP) in the same transect can be computed:

U_x can be either larger or smaller than U₁or U₂, and is the basis for removing the effects of rebound from a modern DEM for construction of the watershed paleogeography at a given time t_x, as described in a later section. Average shoreline slope (S_x) = relative uplift (U_x) / transect length (d) or:

Construction of a Multi-Basin Response Surface of Algonquin-Age Rebound About 10.6 Ka Bp (12.6 Cal Ka Bp)

Profiles perpendicular to isobases were selected on which an uplifted surface of Algonquin age could be constructed throughout the Great Lakes watershed (see locations of profiles on inset map of Fig. 5). Starting with area B on the Michigan-eastern Superior profile, the slope of the Elderdon phase shoreline of Lake Wisconsin (about 14 ka BP or 16.7 cal BP) was adjusted to the selected reference age, 10.6 ka BP (about 12.6 cal ka BP) using Equation 5 (Fig. 5). Then the reference profile was extended to the northeast using the Glenwood II shoreline. The uplift of the second isobase of the Glenwood II shore relative to its lowest isobase was adjusted for uplift using Equation 5. This procedure was repeated for each Glenwood II isobase until a complete shore profile, adjusted to the reference age, could be plotted. The adjusted Glenwood II profile was then shifted vertically to connect with the uplifted end of the Wisconsin profile. Next, the Algonquin profile, already at the selected reference age, was shifted vertically to connect with the end of the adjusted Glenwood II profile. The Minong isobases were then adjusted from an age of 9.5 ka BP (10.7 cal ka BP) to 10.6 ka BP (12.6 cal ka BP), and the adjusted shore profile was shifted vertically to connect with the end of the Algonquin profile. The continuous curve rising from zero to about 200 m at 800 km distance and beyond represents uplift on the Michigan-eastern Superior profile since 10.6 ka BP (12.6 cal ka BP). This rebound is relative to area B just beyond the Laurentide ice limit in the southwestern corner of the Great Lakes map area.

The adjusted Algonquin-age uplift values were transferred along connecting isobases to the other profiles (inset map on Fig. 5). After populating these profiles with Algonquin-age uplift values, contours were drawn throughout the region honouring the profile data and isobase trends. These contours constitute the reference response surface (Fig. 6) for the Great Lakes basin, representing isostatic rebound of the region from 10 600 BP (12 600 cal BP) to the present relative to area B. For a response surface at any other age, t_x, the contour lines remain the same, but each uplift contour value becomes Ut_x using Equation 5, i.e.

where A_contour is the amplitude factor of the contour, such that:

and U_10.6 ka is the value of the contour from the reference response surface.

Paleogeographic Reconstruction

The use of an isostatic response surface in paleogeographic reconstructions is described by Leverington et al. (2002) for Arctic Canada, and is similar to the use of relative-sea-level isobase maps as employed for reconstruction of Atlantic Canada paleogeography (Shaw et al., 2002). For paleogeographic reconstruction of the Great Lakes basin at a specific age, values for contours of an isostatic response surface at the specific age were first computed as outlined above in Equations 7 and 8. The new reference uplift contours, expressed as a vector data set of isobases, were transformed to a gridded data set or surface using the Triangulated Irregular Network (TIN) interpolation method. The interpolated values from this surface were subtracted from each corresponding pixel value of the modern Great Lakes digital elevation model (DEM) (Fig. 1) to generate a paleo-DEM for the desired age. Paleo-lake shorelines were determined by contouring the paleo-DEM within individual basins according to the elevation of their outlet sill or constriction, which is known from geological data. Paleo-DEM pixel values for areas within the shoreline contours were subtracted from the shoreline elevations to express paleo-lake water depths, which in turn were used to calculate lake area and volume. When present in the map area, an ice cover was superposed using information from the glacial geological literature (Table I) and from syntheses of deglaciation such as Barnett (1992) and Dyke et al. (2003). Twelve paleogeographic reconstructions, ranging from the Kirkfield Algonquin phase at 11.4 ka BP (13.3 cal ka BP) to the Nipissing Great Lakes at 5 ka BP (5.7 cal ka BP), were compiled using a GIS; images of two of these reconstructions are shown in Figure 7.

Application to Paleohydrological Modeling

The 12 GIS paleogeographic reconstructions based on the empirical model of isostatic adjustment were used to measure land and water areas, and lake volumes for studies of the paleohydrology and meltwater flow through the Great Lakes system (Moore et al., 2000). Variations in the basin attributes in the 11.4 to 5.0 ka BP (13.3 to 5.7 cal ka BP) period were substantial (Fig. 8). Maximum variation in individual basins under overflow conditions ranges from +72% to ‑95% for lake area, and from +200% to ‑97% for lake volume, compared to the present Great Lakes.

Reconstruction of Former Lake Levels and Discovery of Closed Lowstands, 8900-7800 CAL BP (7.9-7.0 KA BP)

Seventy-seven radiocarbon-dated, and two other, indicators of former lake levels in the Georgian Bay, Huron and Michigan basins (Table II and examples shown in Fig. 9) have been restored to their original elevations using a site uplift equation:

Here, the elevation at time t cal BP is equated to the present elevation (E_p) minus the uplift of the site since time t. This equation, based on Equations 1 or 5, is used to compare the relative altitudes of two or more lake-level indicators or outlets at various times during their isostatic adjustment.

Original elevations for the reported error limits of each dated lake-level indicator were computed using Equation 9 and are shown in Table II. Most indicators are plotted in Figure 10 for the Huron and Georgian Bay basins, and in Figure 11 for the Michigan basin for the interval 11.7 to 6.2 ka BP (13.5 to 7.1 cal ka BP). On these figures, pairs of symbols joined by tie lines indicate the age and reported error range (X axis) in the date of each water surface indicator (Table II). The original altitude of an indicator is the Y-axis value of the plotted symbol; the tilt of the tie line represents uplift during the reported error range of its age. See Figure 1 for locations of the numbered indicators. The lake level histories were interpreted from the restored elevations of the radiocarbon-dated geologic indicators of former water levels in Huron and Michigan basins. For the times when lakes in the Michigan and Huron basins were confluent (relatively high levels), water levels in the Michigan basin were transferred from the Huron basin diagram.

Figure 7

Paleogeographic maps showing reconstructions of the bathymetry and topography of the Great Lakes basin. (A) Main Lake Algonquin phase about 10.55 ± 0.1 ka BP (12.4-12.7 cal ka BP). At a late stage of Main Lake Algonquin, shown here, the lake had expanded to a possible maximum area in Superior (dashed line) and Huron basins by calving of icebergs from glacier margins in deep water. This lake overflowed via Port Huron to the Erie basin. Early Lake Erie, controlled by the Lyell-Johnson sill near the present Niagara Falls, overflowed via Niagara River to Early Lake Ontario which discharged via the emerging St. Lawrence River to Champlain Sea. (B) Mattawa highstand phase about 8.7-8.8 ka BP (9.6-9.9 cal ka BP), possibly a phase of high discharge when lake level in the upper Great Lakes basins was hydraulically dammed by constrictions downstream from North Bay at either the Rutherglen Moraine or the Rankin Constriction or both (Fig. 1). Discharge continued via Ottawa River to St. Lawrence River. There was no drainage from the upper to lower Great Lakes at this time; Lake Erie, Niagara River and Lake Ontario drained a separate watershed to the upper St. Lawrence River. Legend as for Figure 7A.

Cartes paléo-géographiques de la reconstitution de la bathymétrie et de la topographie du bassin des Grands Lacs. (A) Phase du lac Algonquin principal en 10,55 ± 0,1 ka BP (12,4-12,7 cal ka BP). Au cours du stade tardif de cette phase, le lac Algonquin a atteint son extension maximale dans les bassins du lac Supérieur (tireté) et Huron par le vêlage des icebergs issus des marges glaciaires en eau profonde. Ce lac se déversait dans le bassin du lac Érié par le Port Huron. Le lac Érié initial, dont le niveau était contrôlé par le seuil de Lyell-Johnson situé près des chutes actuelles du Niagara, se déversait dans le lac Ontario initial par la rivière Niagara, se déversant à son tour (le lac Ontario) dans le Saint-Laurent naissant par la Mer de Champlain. (B) Phase de haut niveau lacustre de Mattawa en 8,7-8,8 ka BP (9,6-9,9 cal ka BP), une phase de fort débit probablement liée à la fermeture hydraulique des Grands Lacs par un étranglement topographique situé en aval de North Bay provoqué par la moraine de Rutherglen, l’étranglement de Rankin, ou les deux. Le déversement dans le Saint-Laurent s’effectuait par la rivière des Outaouais. Il n’y avait pas de drainage entre l’amont et l’aval des Grands Lacs à cette époque : les lacs Érié et Ontario et la rivière Niagara drainaient un bassin-versant distinct vers le haut Saint-Laurent. La légende est la même qu’en A.

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The accuracy for estimating lake levels from empirical indicators is generally considered here to be at best about ±1 m based on survey error and variations in elevations of strandline bluffs, beaches, bars and spits (Schaetzl et al., 2002). The computed initial elevations for lake-level indicators are considered accurate within about 2 m in areas of good isobase control, and somewhat >2 m in areas where constraining isobases have been extrapolated.

A second set of data (Table III) comprising the uplift history of lowest-possible potential overflow outlets for the Huron and Michigan basins, was also computed in the same manner and plotted on the diagrams of Figures 10 and 11 as dotted bands and dashed lines. These constraints include sills at North Bay (dotted band), Dalles Rapids, Deane-Tovell Saddle, and the head of Mackinac River (Lewis and Anderson, 1989). The North Bay sill controlled the overflow level of the water in the combined Georgian Bay, Huron and Michigan basins during the Nipissing transgression (Karrow 1980; Monaghan et al., 1986; Colman et al., 1994a, 1994b). The Dalles Rapids sill controlled overflow levels of the Georgian Bay basin during earlier phases while the North Bay sill was isostatically depressed at lower elevations. The Lucas and Fitzwilliam channel sills were selected as representative of constraints on overflow water levels in the northern Huron basin by the Deane-Tovell Saddle between Bruce Peninsula and Manitoulin Island. The Michigan basin overflow water levels were controlled by a sill at the head of the now-submerged Mackinac River channel between Michigan and Huron basins beneath northern Lake Michigan and the Straits of Mackinac (Stanley, 1938b).

Interpretation of Huron and Georgian Bay Water-Level History

13 500-11 000 cal BP

In this synthesis, the interpreted water-level history is shown as a thick line in Figures 10A-B and 11, and begins with the transgression of the Kirkfield Algonquin lake level to the Main Algonquin level (sites 4, 2, 1c, 1b, 3) under influence of the isostatically-rising Kirkfield outlet (Fig. 1). After 12 600 cal BP (10 600 BP), lake levels fell, indicated by the initiation of gyttja sedimentation in isolated basins (sites 12, 13, 10), through the Post-Algonquin phases and below, evidenced by the onset of organic accumulation on Manitoulin Island (23) and in Georgian Bay (7b). Land emergence by this fall of lake levels through the Post-Algonquin phases is recorded further by additional, but somewhat delayed, organic sediment accumulation in small basins south of Georgian Bay (5, 6), near North Bay (24, 25) and on Manitoulin Island (8).

Figure 9

Types of evidence of former water levels in the Huron, Michigan, and Erie basins. Abandoned geomorphic shore features: (A) Former lake surfaces are inferred from erosional coastal features at the change in gradient between the gently-sloping shoreface and the steeply-sloping shorebluff, illustrated here for the Algonquin shore near Kirkfield, Ontario. (B) Storm beach ridges above Lake Huron on Manitoulin Island (air photo courtesy of Natural Resources Canada). (C) and (D) Beach ridges 53 m below Georgian Bay imaged by sidescan sonar and multibeam sonar, respectively (Blasco 2001). (E) Submerged beach and shoreface 21-28 m below Lake Erie (from Coakley and Lewis, 1985). Isolation basins, tree stumps, and unconformities: (F) Air photo (courtesy of Natural Resources Canada) view of a small basin isolated during the regression of a large lake, Manitoulin Island. A radiocarbon date of the contact zone between gray clastic large lake sediment and dark-coloured organic small-lake-sediment (photo courtesy of T.W. Anderson) provides the age at which the Great Lake surface passed below the sill elevation of the small isolated basin (Lewis, 1971). (G) In situ tree stump on lakefloor at entrance to Georgian Bay. (H) Subsurface reflections in a seismic profile indicate offshore unconformities and sequence boundaries caused by episodes of reduced lake level (Moore et al., 1994).

Indicateurs des anciens niveaux lacustres dans les bassins des lacs Huron, Michigan et Érié. Formes littorales abandonnées : (A) les surfaces des anciens lacs sont estimées à partir de formes d’érosion côtière qui sont situées à la rupture de pente entre la face de la côte en pente douce et le talus littoral en pente abrupte, comme le montre cet exemple du rivage du lac Algonquin près de Kirkfield, Ontario. (B) Crêtes de plage de tempête en bordure du lac Huron sur l’Île Manitoulin (courtoisie de Ressources Naturelles Canada). (C) et (D) Crêtes de plage à 53 m de profondeur dans la baie Géorgienne identifiées respectivement à l’aide du sonar latéral et d’un sonar multi-faisceaux (Blasco, 2001). (E) Plage et avant-côte submergées entre 21 et 28 m sous la surface du lac Érié (d’après Coakley et Lewis, 1985). Bassins isolés, souches d’arbres et discontinuités : (F) Photo aérienne d’un petit bassin isolé durant la régression d’un grand lac, Île Manitoulin (Source : Ressources Naturelles Canada). Une datation au radiocarbone de la zone de contact entre les sédiments clastiques gris du grand lac et les sédiments organiques foncés du petit lac (photo fournie par T.W. Anderson) fournit l’âge auquel le plan d’eau du grand lac est passé sous le niveau du seuil du petit bassin isolé (Lewis, 1971). (G) Souche d’arbre in situ au fond du lac à l’entrée de la baie Géorgienne. (H) Profil sismique montrant les discontinuités et les limites causées par des épisodes de bas niveau lacustre (Moore et al., 1994).

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Table II

**Indicators of former lake levels in the Michigan, Huron and Georgian Bay basins, and Georgian Bay-North Bay lowland**

Table III

**Present and early Holocene elevations1 of North Bay, Georgian Bay, Huron and Michigan sills (m above present sea level)**

11 000-10 000 cal BP

Relative lake levels continued to fall toward the level of driftwood deposited in Georgian Bay (31b). During the lowstand, trees grew in the Straits of Mackinac (28b), and organic sediments accumulated at relatively low-elevation at northern and southern Huron basin sites (29c, 30f). A subsequent abrupt lake-level rise and inundation is recorded by a shift to more aquatic species in Georgian Bay peat at site 7 (Lewis and Anderson, 1989). The lake level rise at about 10 800 cal BP (9500 BP) was possibly in response to the onset of outburst floods from upstream glacial Lake Agassiz through its eastern outlets to Superior basin (Clayton, 1983; Teller and Thorleifson, 1983; Farrand and Drexler, 1985; Lewis and Anderson, 1989; Breckenridge et al., 2004; Breckenridge and Johnson, 2005). The rise of lake level during high-discharge events may be related to constrictions in the Mattawa River valley at the Rutherglen Moraine (Harrison, 1972; Chapman, 1975) and in Ottawa River valley at the Rankin Constriction (Lewis and Anderson, 1989) (Fig. 1). The maximum of the lake-level rise possibly exceeded the altitude of the Sheguiandah site on Manitoulin Island (9c, 9d) where swamp organic sediment began accumulating just after the lake-level rise. Some sites with thick clay beneath gyttja (41, 33) suggest the waters of this inundation carried abundant fine-grained sediment in suspension. This pulse of higher water (Lake Mattawa) was apparently short-lived, and its decline may account for the emergence of small basins and onsets of gyttja sedimentation or peat accumulation at several sites (9c, 9d, 29c, 30b, 30c, 30d, 30f, 33, 34b, 38, 40, 41).The lake level drop was low enough to expose the eastern Huron basin lakebed for tree growth at site 66. Within a century or so, rising water level allowed for deposition of driftwood (64 then 32c), and terminated peat accumulation at slightly higher sites in southern Huron (39) and Georgian Bay (7a) basins.

10 000-9000 cal BP

The rising lake level that terminated peat accumulation (7a) continued to rise to a second high-level Lake Mattawa phase (about 9800 cal BP, 8700 BP) that inundated Pure Lake (50), and possibly culminated at a high enough level to induce sediment aggradation in Amable du Fond River (49) and to construct a barrier lagoon north of Lake Huron (43) (Fig. 10B). Thick clay deposits beneath gyttja in Pure Lake suggest that this water-level rise was also accompanied by a high influx of fine-grained suspended sediment. An extended period of moderately low water level following the Mattawa highstand is indicated by tree growth on the saddle between Huron and Georgian Bay basins (67, 68a, 68b), and by peaty marsh sediments in southwestern Huron basin (30a) and South Bay, Manitoulin Island (42a, 42b, 42c). Oddly, an erosional event and sequence boundary in the deep water sediments of Huron basin is not recorded at this time of low water level (Fig. 10; Moore et al., 1994; Rea et al., 1994a).

At about 9050 cal BP (8100 BP) water levels again rose, transforming sedimentation in South Bay from plant detritus to clastic silty clay (42a, 42b), and possibly overflowing the North Bay outlet where delta construction occurred at nearby Trout Mills (48). Water levels then rose abruptly to the last Mattawa highstand, likely as a result of overflow from large volumes of subglacial meltwater discharged during the Nakina ice advance over the Great Lakes-Hudson Bay divide, and recorded throughout the Superior basin as a set of 36-40 thick varves dated 9035 ± 170 cal BP or about 7950-8250 BP (Breckenridge et al., 2004). Although the maximum lake rise in Huron and Michigan basins is not known, it apparently fell short of the Sheguiandah archeological site on Manitoulin Island as it is not recorded in the peat sequence there between dated samples 9b, 9c and 9d (Anderson, 2002). Land emergence following this short-lived high Mattawa phase lake is documented by onsets of organic sedimentation at Wye Marsh, southern Georgian Bay (72), gyttja deposition in small basins northeast of Georgian Bay (44, 45, 46, 47), and tree growth at low elevations now under water in Huron basin (75a, 75c) and possibly in the Straits of Mackinac (28a).

Figure 10

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9000-7000 cal BP

Lake-level decline from the last high Mattawa phase continued down to levels defined by the Stanley unconformity (77) in Huron basin (Lewis et al., in press) and the Flowerpot beach (FB) in Georgian Bay basin (Fig. 10). This lowstand phase, which continued for several centuries, formed the Light Blue reflector and sequence boundary in deep water sediments (Moore et al., 1994; Rea et al., 1994a), and allowed tree growth on the Deane-Tovell saddle between Huron and Georgian Bay basins (65, 69, 70). Lake-level rise from this low phase by 7.5 ka BP (8.3 cal ka BP) is suggested by the chronology of the Light Blue reflector in northern Huron and Georgian Bay basins (Moore et al., 1994).

The final increase in lake levels from the low phases of late Lake Stanley and late Lake Hough in Huron and Georgian Bay basins, respectively, probably occurred about 8000 cal BP, late enough to allow tree growth in eastern (71) and western (52) Huron basin, but early enough to transgress the Rains Lake site (53b) in northwestern Huron basin. At this elevation, the water surface had risen to the level of the Nipissing beach at North Bay, indicating that the lake in Michigan-Huron-Georgian Bay basins was overflowing the North Bay outlet at full discharge. As this outlet uplifted faster than other parts of the lake basin, relative lake levels rose throughout most of the upper Great Lakes basins as the well-known Nipissing transgression, manifested in the present data by transgression of the Rains Lake (53b) and Smoky Hollow Lake (34a) sites (Fig. 10A).

Interpretation of Water-Level History of the Michigan Basin

Because of their connection via the Indian River lowland and Straits of Mackinac (Fig. 1), high water levels were always at common elevations in the Huron and Michigan basins, as at present, following retreat of ice from their northern regions (Eschman and Karrow, 1985; Hansel et al., 1985). Consequently, the history of lake-level variation above the Michigan basin sill defined in the Huron basin is applicable to the Michigan basin as shown in Figure 11. Only evidence of lowstands and other unique indicators of Michigan basin lake elevations are shown and discussed here.

Submerged in situ tree stumps in southern Lake Michigan (the Olson Forest of Chrzastowski et al., 1991) have been traditionally interpreted as having been drowned in the Nipissing transgression (Chrzastowski and Thompson, 1992, 1994; Colman et al., 1994b). However, in this analysis, the Olson tree stumps (sites 74a, 74b at 8.2-8.4 ka BP, 9.2-9.4 cal ka BP, in Figs. 1 and 11) appear tens of metres above the North Bay outlet, and could not be affected by the Nipissing transgression. A rise of water level to the final highstand of Lake Mattawa follows the Olson Forest after 8.2 ka BP (Fig. 11). This flooding event seems to have been the cause of forest drowning.

As in the Huron and Georgian Bay basins, an extreme decline in lake level following the last Mattawa highstand is suggested in the Michigan basin by the occurrence of shallow-water sediments and molluscan fauna in the deepwater environment of central Lake Michigan (51a, 51b, 51c) (Lewis et al., in press). A final recovery of lake level to an overflowing condition at the North Bay outlet probably occurred about 8.1-7.8 cal ka BP, allowing for tree growth in northern Michigan basin (54). From this time forward, relative lake level rose under control of the North Bay outlet as the Nipissing transgression.

Closed Lowstand Conditions

During the period 8.95-8.3 cal ka BP (8.05-7.4 ka BP) lake levels (late Stanley in Huron basin, late Hough in Georgian Bay basin, and late Chippewa in Michigan basin) were below the sill of the North Bay outlet, the lowest possible overflow outlet at the time (Figs. 10B‑11). Thus the late Stanley, late Chippewa, and late Hough phases were hydrologically closed lakes at their lowest level, and as such may have resulted from the impact of a severe dry climate in which evaporative water losses exceeded water inflows by precipitation and runoff. The inference of dry climate is supported by the presence of thecamoebians that indicate a more saline lake environment, and by climate transfer function analysis of pollen assemblages which suggests less precipitation and warmer temperatures at 7.7 ka BP than today in the Georgian Bay area (Blasco, 2001). Other lowstands around 9.3 and 9.8 ka BP (10.5 and 11.2 cal ka BP) may have also been closed briefly.

Possible shorelines for the dry climate-induced closed lowstands at 7.8 ka BP (8.6 cal ka BP) are illustrated in a paleogeographic reconstruction of the Huron, Erie and Ontario basins on Figure 12. The northern Huron lowstand shore is tied to the Stanley unconformity beneath northwestern Lake Huron (Hough 1962; Lewis et al., in press). The Georgian Bay lowstand shore is tied to the low-level Flowerpot beach in the entrance to Georgian Bay (Blasco, 2001). Water bodies in these basins and those beneath southern Lake Huron were isolated closed lowstands, well offshore from the present lake boundaries. The Erie lowstand shore is tied to a submerged beach in eastern Lake Erie (Fig. 9E; Coakley and Lewis, 1985; Lewis et al., 2004) at a lake level that would have extended into the central Lake Erie area before that sub-basin became mostly infilled with sediment (Sly and Lewis, 1972). The Ontario basin is assumed to have been impacted by the same dry climate that affected the other basins. Under these conditions, its lowstand shoreline was inferred at a level that made its basin area-to-lake area ratio equal to that of the Huron basin (Bengsston and Malm, 1997).

Discussion

Geophysical Models of Isostatic Adjustment

With knowledge or estimates of the elastic and viscous properties of Earth’s lithosphere and mantle, respectively, geophysical models compute the crustal isostatic response following a known or inferred history of ice sheet loading and pattern of deglaciation. Clark et al. (1994) used this approach to evaluate isostatic movements during deglaciation of the Great Lakes basin. They successfully illustrated the increasing amplitude of postglacial isostatic rebound towards the north and northeast in the direction of past thicker ice and ice retreat in accordance with the evidence of deformed paleo-lake shorelines. They applied an innovative approach to constrain and calibrate their model by using geological knowledge of large-lake drainage transfers from northern to southern outlets as basins were differentially tilted during the Algonquin and Nipissing phases. Lake-level history was derived by tracking the upward movement of overflow sills through time. Lake-level indicators were not tracked separately from the sills with the result that possible episodes of hydrologic closure could not be detected.

Gravitationally self-consistent models of glacio-isostatic adjustment on a global scale have been progressively developed and improved over the past few decades (Peltier, 1998). These models are constrained and calibrated to relative sea-level histories at ocean-continent boundaries. Model estimates of isostatic adjustment are produced for continental interiors, and one of these, the ICE‑3G model (Tushingham and Peltier, 1991), has been compared favourably with short-term evidence of tilting of the Great Lakes basins derived from trend analysis of lake-level gauge records (Tushingham, 1992). However, over the longer term, the geophysical model results are not in complete agreement with past events in the watershed. The relative performances of two geophysical models and the empirical model were assessed by comparing the elevations of the northern and southern outlets from the upper Great Lakes basins during the Algonquin and Nipissing drainage transfers which are known to occur at about 10.5 ka BP and 5 ka BP, respectively, from independent geological evidence (Karrow et al., 1975; Karrow, 1980; Monaghan et al., 1986). Outflows were transferred from the Kirkfield outlet to the Port Huron and possibly Chicago outlets during the Algonquin transfer (Fig. 7A), and from the North Bay to the same southern outlets during the Nipissing transfer. Estimates of the elevations of the overflow sills and timing for these drainage transfers when northern outlets rose isostatically above southern outlets were provided for the ICE‑4G glacio-isostatic model (Peltier, 1995, 1996) courtesy of W.R. Peltier. Differences between Port Huron and the other outlets which should have been near zero at the times of the drainage transfers are shown in Figure 13. The empirical model performs best with northern and southern sills being within two metres of each other during transfers. The Clark et al. (1994) model is similar, although the best agreement for the Algonquin transfer was obtained at 10 ka BP, rather than the expected 10.5 ka BP, and the Nipissing-age transfer at Chicago differed by 3 m. Variances in sill elevations during transfers for the ICE‑4G model are larger, and ranged from 3.5 to 10 m.

Near and beyond the maximum margins of the ice sheets, the geophysical models predict crustal uplift as a forebulge of modest relief compared with the amplitude of depression beneath the centre of the ice load. This has been demonstrated in ice-marginal Atlantic and Arctic coastal regions (Barnhardt et al., 1995; Dyke, 1998). The forebulge effect has not been recognized from empirical evidence in the continental region adjoining or south of the Great Lakes basin. However, ice marginal areas in this region could have undergone uplift and subsidence associated with the growth, migration and decay of a glacio-isostatic forebulge (Colman et al., 1994a), and some evidence for subsidence exists, for example, anomalies in modern tilting of the Erie and southern Michigan basins (Mainville and Craymer, 2005).

Benefits of geophysical models are their ability to compute both vertical and horizontal earth movements associated with the isostatic process, and to relate these estimates to an absolute datum such as present sea level. Confidence in predictions of past crustal movements in the Great Lakes region by geophysical models, especially deformation over several millennia, will be increased markedly when these models are calibrated and constrained by the empirical observations of glacio-isostatic adjustments within the same region.

Previous Recognition of Low Lake Levels in Huron Basin

An extreme lowering of lake level in the Huron and Michigan basins after the Algonquin phases was predicted by Stanley (1936) when the ice sheet margin receded from the isostatically-depressed lowland drainage route past North Bay, Ontario, to the Mattawa and Ottawa river valleys. Sediment unconformities, discovered by Hough (1955, 1962), confirmed lake-level lowstands in each of the Michigan and Huron basins; these lowstands were named Chippewa and Stanley, respectively. Hough (1962) inferred a delay in isostatic adjustment so that, in his model, the North Bay sill remained at its depressed Algonquin level, and Lake Stanley could be interpreted as an open overflowing water body (Lewis et al., in press).

Although a regional advance (Marquette readvance) of the Laurentide Ice Sheet about 500 km wide reached the southern coast of the Superior basin at 10 ka BP (11.5 cal ka BP) (Lowell et al., 1999), and ice retreat was slowed during the Younger Dryas in eastern Ontario and western Québec (Simard et al., 2003), there is no evidence of ice advance or of delayed isostatic recovery in the Georgian Bay-North Bay outlet region, about 300-500 km east of the Superior basin (Dyke et al., 2003). As a result, it is highly unlikely that rebound was delayed significantly and, at the time of late Lake Stanley (Fig. 10), the recovered outlet was above water level, as was also interpreted by Lewis et al. (in press). Closed lake conditions for late Lake Stanley are also consistent with the preliminary paleoecological findings based on thecamoebian analysis of somewhat higher salinity in late Lake Hough, the equivalent closed water body in the Georgian Bay basin, as reported by Blasco (2001).

Working Hypothesis for Climate Change and the Hydrologically Closed Lowstands

Major shifts in Great Lakes water levels have long been understood in terms of overflowing lakes. Lake elevations change as a result of shifting ice dams during glacial retreat or advance, outlet erosion, or by differential glacio-isostatic adjustment of the outlets (Eschman and Karrow, 1985; Hansel et al., 1985; Larsen, 1987; Barnett, 1992; Larson and Schaetzl, 2001). These mechanisms apply to the Algonquin highstand and the transgression to the Nipissing Great Lakes. The intervening Mattawa high phases (Lewis and Anderson, 1989) are also considered to be overflowing lakes, but with variable levels controlled by resistance at hydraulic constrictions downstream of North Bay to variable high-discharge flows from upstream Lake Agassiz or subglacial drainage.

The discovery of low lake levels below the lowest possible overflow outlet in their basins implies a controlling factor of excess evaporation (water loss) over precipitation and runoff (water supply) during a period of dry climate. These lowstands, the late Chippewa, late Stanley, and late Hough lakes in the Michigan, northern Huron, and Georgian Bay basins, respectively, are thought to reflect a phase of climatically-driven hydrologic closure that has not been recognized previously. The earlier pre‑9.5 ka BP (pre‑10.9 cal ka BP) lowstands, middle and early lakes Chippewa, Stanley and Hough, are inferred to be close in level to, or slightly below, their basin outlets, and thus were possibly also subject to enhanced evaporative losses of water. During these early, low lake phases, sustaining upstream inflows were not available, as Lake Agassiz discharge was diverted away from the Great Lakes by advances of ice in the Superior and Nipigon basins (Thorleifson and Kristjansson, 1993; Lewis et al., 1994). Dry air from the glacial atmospheric circulation over the nearby Laurentide Ice Sheet (Bryson and Wendlund, 1967; David, 1988; Anderson and Lewis, 2002; Wolfe et al., 2004) probably exerted continuous evaporative stress on nearby water surfaces, causing drawdown of these lakes during phases of reduced inflow.

By the time of the latest and longest-duration lowstands (late Chippewa, Stanley and Hough) the effects of glacial atmospheric circulation would have diminished greatly owing to the small size of the remaining Laurentide Ice Sheet, and to its more distant location relative to the Great Lakes basin. At this time (8 ka BP, 8.9 cal ka BP), meltwater drainage from the merged glacial lakes Agassiz-Ojibway began bypassing the upper Great Lakes basins directly into the Ottawa Valley (Veillette, 1994; Teller et al., 2002; Teller and Leverington, 2004). Thus, the Great Lakes watershed suddenly lost a significant source of water supply and was susceptible to the early Holocene dry climate (Edwards et al., 1996).

Atmospheric water supply today can be thought of as a function of the relative time spent over the Great Lakes basin by three major air masses over North America. These masses are the Arctic air from the north (dry and cold), the Pacific air from the west (dry and warm), and the Maritime Tropical air bringing moist, warm air north from the Gulf of Mexico (Bryson and Hare, 1974; Bradbury and Dean, 1993). Once glacial lake drainage began bypassing the Great Lakes, and as the Laurentide Ice Sheet downwasted over Hudson Bay, southward incursions of dry Arctic air, previously blocked by the high Laurentide Ice Sheet, were likely becoming increasingly frequent (Yu and Wright, 2001), and this dry air may have initiated or intensified draw down of lake levels to the late Chippewa, Stanley, and Hough lowstands. The closed lowstands may have been maintained by enhanced evaporation into increasingly strong flows of dry, warm Pacific air from the west. These flows are indicated by abundant evidence of vegetation shifts to drought-resistant plant taxa west of the Michigan basin (Baker et al., 1992; Wright et al., 2004). By 7 ka BP (7.8 cal ka BP), increasing incursions of the Maritime Tropical air mass were delivering sufficient precipitation, indicated by the appearance of mesic forest species in pollen diagrams of the Great Lakes region (Webb III et al., 1998) to convert the Michigan, Huron and Georgian Bay water bodies to open, overflowing lakes, as at present. These changes are consistent with paleovegetation maps which show a rapid northward migration through the Great Lakes region of the mixed-boreal forest biome boundary between 8.0 and 7.0 ka BP (8.9 and 7.8 cal ka BP) (Dyke et al., 2004).

The foregoing hypothesized climatic history is consistent with changes inferred for small Elk Lake in Minnesota, based on a comprehensive study of proxy climatic and limnological indicators (Bradbury and Dean, 1993). Similarly, this climatic history is supported by a coeval depletion in the ¹⁸O composition of precipitated carbonate in Deep Lake, Minnesota, attributed by Yu and Wright (2001) to the blocking of southern air masses by more frequent presence of dry Arctic air.

The phase of reduced and closed lakes at the onset of the present hydrologic regime offers a unique opportunity to evaluate the sensitivity of the Great Lakes system to high-amplitude climatic change. Understanding the sensitivity of the lakes to high-amplitude and long-duration change would be a distinct benefit in the light of the need to project and adapt to future changes under global warming which may drive the lakes below instrumentally-observed variability (Mortsch et al., 2000).

Isotopic Composition of Lake Water

As the evaporation process favours concentration of water molecules containing the heavier ¹⁸O isotope, previous findings of high concentrations of the lighter ¹⁶O isotope, similar to that of glacial meltwater, in fossil valves of Huron basin benthic ostracodes at about 7600 ka BP (Rea et al., 1994a, 1994b; Dettman et al., 1995) appear to contradict the coeval presence of evaporatively-driven lowstands as postulated in this paper and in Lewis et al. (in press). Although surface lake water isotopic composition undoubtedly became concentrated in ¹⁸O during seasonal periods of rapid evaporation, it apparently did not greatly influence bottom water composition, similar to the isotopic stratification found for glacial Lake Agassiz (Buhay, 1998; Birks et al., in press). Alternatively, adjustments in the chronology of isotopic events in the Huron and Michigan basins suggest that the low ¹⁸O inflow occurred prior to the lowstands (Breckenridge and Johnson, 2005; Breckenridge, in press), and may have remained as bottom water during the evaporative phase. Additional research is needed to resolve the origin of bottom water in the Huron and Michigan basins.

Summary

Postglacial isostatic adjustment in the Great Lakes region is described here in the time domain using an exponential decay expression constrained by the observed cumulative differential deformation of dominant paleo-lake strandlines in individual or groups of basins. Vertical earth movement throughout and following the last deglaciation into the middle Holocene is characterized as progressive, differential uplift relative to an area southwest of southern Lake Michigan, beyond the limit of the last glacial maximum. Rates and amplitudes of uplift increase towards the north-northeast in the direction of deglacial retreat and thicker ice. A mean relaxation time for the uplifting process of 3700 ± 700 years was obtained by averaging solutions for the decay time parameter at 20 transects throughout the basin where isobases (gradients) of two strandlines of different age were known. This relaxation time was used in the exponential expression to adjust the isobase gradients of dominant strandlines in the Great Lakes basins to 10.6 ka BP (12.6 cal ka BP), the approximate age of the well-known Algonquin phase.

Collectively, the Algonquin and adjusted isobases constitute a reference response surface for isostatic adjustment throughout the Great Lakes region. Interpolated values from this surface and the mean relaxation time were used in the exponential uplift expression to determine an ‘amplitude’ factor for uplift at any desired location. With these values and the exponential expression, uplift since any desired age could be computed. The original elevation of a site at any desired age could also be determined by subtracting the computed uplift from the present elevation of the site. This approach was used to transform values of pixels in a DEM for the present Great Lakes region to new DEMs at previous ages. A total of 12 paleogeographic reconstructions for the topography and bathymetry of the Great Lakes basins were prepared for ages between 11.4 and 5.0 ka BP (13.3 and 5.7 cal ka BP). These reconstructions showed that water surface areas ranged from +72% to ‑95%, and lake volumes from +200% to ‑97%, relative to the present lakes. Improvements in the estimation of glacio-isostatic effects can be expected from geophysical models when these are calibrated to the available observations of differential rebound in the Great Lakes basin.

The same empirical approach was applied to reconstruct the original elevations of dated indicators of former lake levels in the Michigan, Huron and Georgian Bay basins for the interval between 11.7 and 6.2 ka BP (13.5 and 7.1 cal ka BP). Original elevations for 79 dated indicators, comprising fossils from beach lagoon sediments, basal organic sediment from isolation basins, shallow-water fossils in unconformable zones within deepwater sediment, and submerged tree stumps in growth position, and others, were reconstructed to form the basis for interpreting lake-level history in these basins. In parallel fashion, the history of potential overflow sills was also reconstructed. Comparison of sill elevations with lake levels revealed a period (about 8.05 to 7.4 ka BP, or 8.95 to 8.3 cal ka BP) in which water surfaces were up to several tens of metres below lowest possible overflow outlets. The low lake levels are postulated to reflect the increased impact of early Holocene dry climate when upstream Agassiz overflow and/or subglacial floods were diverted around the Great Lakes basin, and water supply to the upper Great Lakes was reduced. Lake evaporation and draw down were likely enhanced by frequent incursions of dry cold Arctic air and, later, warm dry Pacific air, possibly related to atmospheric reorganization associated with the demise of the Laurentide Ice Sheet. This period of closed lakes, early in the present hydrological regime of the Great Lakes, offers an opportunity to probe and understand the sensitivity of the Great Lakes system to high-amplitude, long-duration climate change. Such information could improve confidence in projections of, and adaptations to, future levels of the Great Lakes under global warming.

Abstract

Résumé

Introduction