Megaflooding associated with glacial Lake Agassiz

Journal Pre-proof Megaflooding associated with glacial Lake Agassiz Timothy G. Fisher

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S0012-8252(18)30553-1

DOI:

https://doi.org/10.1016/j.earscirev.2019.102974

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EARTH 102974

To appear in: Received Date:

27 September 2018

Revised Date:

30 September 2019

Accepted Date:

3 October 2019

Please cite this article as: Fisher TG, Megaflooding associated with glacial Lake Agassiz, Earth-Science Reviews (2019), doi: https://doi.org/10.1016/j.earscirev.2019.102974

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Megaflooding Associated With Glacial Lake Agassiz

Timothy G. Fisher Department of Environmental Sciences, MS 604 University of Toledo, 2801 West Bancroft St., Toledo, OH

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43560-3390, USA

Abstract

Significant baseline and episodic megaflooding from glacial Lake Agassiz was routed to the

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south, east and north coasts of North America over the lake’s nearly 6000 year history. The five phases of lake-level change were controlled by which outlet was active, which in turn was

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controlled by ice margin position and glacioisostatic adjustment. The southern outlet is the oldest and best understood outlet, while successively younger outlets are progressively less understood. Eastern drainage synchronous with the Younger Dryas chronozone had been

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assumed by most studies, but a spillway to accommodate the ~90 m drop in lake level is yet to be described. With the south and east outlets unable to accommodate the necessary lake level drawdown, the northwest outlet has become the default outlet with Arctic Ocean core data

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supporting it. However, extant terrestrial data from the continent only provides data for a large flood from the Fort McMurray area closer to the end of the Younger Dryas, coinciding with the timing of the Preboreal Oscillation. The challenge remains to find agreement between the marine and terrestrial records. Uncontroversial is the geomorphic and sedimentologic evidence for flooding into and out of the lake consisting of large spillways, large boulders, and in places

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giant current ripples composed of boulders. Flood discharge estimates into and out of the lake range from 0.04 to ~1Sv, with final subglacial drainage into the Tyrell Sea estimated at ~5 Sv, sufficient to raise global sea level by 0.18 m and initiate the 8.2 ka stadial. Keywords Outlets, boulders, flood routing, spillway, Younger Dryas, Preboreal Oscillation

1.1 Introduction:

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Regional interest in glacial Lake Agassiz became international interest when Broecker et al. (1989) proposed that meltwater delivery to the Gulf of Mexico switched to the North Atlantic and triggered a slow-down in thermohaline circulation initiating the Younger Dryas stadial. Dubbed by many as the world’s largest glacial lake, Lake Agassiz evolved over nearly 6000 years through a series of configurations across the center of the North American continent (Fig. 1). Large volumes of meltwater entered the lake along its western and ice-contact sides and were episodically routed to the southern, eastern, and northern sides of North America by outlet-sill incision, deglaciation of lower outlets and glacioisostatic adjustment (GIA). However, an

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understanding of routing history is incomplete. The routing problem has been approached from two different directions—from marine basins looking inland for sources of meltwater, and from

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spillway channels delivering meltwater to the oceans. Oxygen isotope and other proxy data from marine sediment cores adjacent to the Mississippi, Saint Lawrence, and Mackenzie rivers record fluctuations in meltwater delivery, the timing of which is ideally paired with terrestrially-

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based reconstructions of ice-margin recession, lake evolution, and drainage. Complicating the problem is reliable chronology. Marine cores are hampered by carbonate reservoir corrections

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and identifying specific sources of the meltwater (i.e., water from Lake Agassiz or some other lake or meltwater source). Terrestrial records are hampered by radiocarbon ages giving only maximum or minimum ages of sediments and landforms—seldom the depositional or formation

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age. Optically stimulated luminescence (OSL) dating can provide ages of deposition, but suffers from larger errors than radiocarbon dating.

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There is ample evidence of megaflooding into and out of the lake consisting of large spillways and characteristic flood deposits. In particular, two very large drops in lake level have been proposed. The first released ~17,000 km3 near the start of the Younger Dryas (Breckenridge, 2015), and 163,000 km3 when the lake finally drained near the start of the 8.2 ka event (Clarke et al., 2004). After a brief overview of lake evolution, lake phases are discussed emphasizing

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water routing, and followed by the megaflooding history associated with various inlets and outlets. This review briefly summarizes pertinent Lake Agassiz meltwater routing history, and evidence of megaflooding from Lake Agassiz. Uncalibrated radiocarbon ages have the units 14C yr (or kyr) BP, calibrated radiocarbon ages use cal yr (or kyr) BP, and OSL or cosmogenic ages use ka.

1.2 Lake phases, levels and routing controlled by ice margin position

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Glacial Lake Agassiz formed once the Des Moines Lobe of the Laurentide Ice Sheet (LIS) receded north of the sub-continental drainage divide separating the Gulf of Mexico and Hudson Bay watersheds (Upham, 1895). Meltwater was trapped between the receding lobe and high topography of the Big Stone Moraine (Fig. 1A) by about 14 ka (Lepper et al., 2007). Opening or closing major outlets routing water to different oceans defines lake phases (Fig. 2). Volumes of water have been estimated for these phases (Leverington et al., 2000, 2002; Table 1). The first major lake phase is the Lockhart Phase (Fig. 2) when drainage was through the southern outlet. The start of the Moorhead Phase is thought to be synchronous with the start of the Younger

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Dryas chronozone (YD) when drainage shifted away from the southern outlet to some combination of eastern or northwestern outlet, and/or evaporation. The Moorhead Phase ended

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and the Emerson Phase began when flow returned to the southern outlet with lake levels at the Campbell Beach (cf. figures in Fisher et al., 2011). Because the Campbell Beach has been mapped to the southern and northwestern outlets, and projected to eastern outlets, it is possible

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that up to three outlets were active at once following the transgression at the end of the Moorhead Phase. The Nipigon Phase eastern (Kelvin) outlets finally captured drainage as the

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LIS continued its recession northwards, uncovering progressively lower outlets as recorded by the suite of beaches at successively lower elevations than the Campbell Beach (Fig. 2). During this phase water drained into Lake Kelvin and out through the Nipigon channels into the Lake

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Superior basin, through the Great Lakes basin and St. Lawrence River, and eventually into the North Atlantic Ocean (Fig. 1). With continued LIS recession glacial Lake Agassiz eventually merged with glacial Lake Ojibway (Ojibway Phase), its counterpart further east in the Hudson

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Bay lowlands. Drainage was south through the Kinojévis outlet into the Ottawa River and then St. Lawrence River. Final drainage of the lake occurred as the ice dam thinned and Lake Agassiz drained subglacially into the Tyrell Sea, the precursor of Hudson Bay. The final drainage is generally assumed to be responsible for the 8.2 ka cooling event (Barber et al., 1999; Clarke et al., 2004). Detailed paleogeography maps of the different lake stages and

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eastern routing of meltwater are available in Leverington et al. (2000, 2002), Leverington and Teller (2003), and Teller and Leverington (2004). 1.3 Lockhart Phase Flights of strandlines from the Herman dropping to the Tintah Beach (Fig. 2) developed as the lake expanded northwards during the Lockhart Phase through a combination of southern outlet incision and GIA. With the southern outlet at an isobase with the lowest rate of uplift for the lake basin, higher rates of rebound northwards resulted in strandlines warped upwards and an

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increasing number of strandlines for any named beach (Upham, 1895; Leverett, 1932; Johnston, 1946; Elson et al., 1967; Teller and Thorleifson, 1983; Teller, 2001; Fisher, 2005; Yang and Teller, 2012; Breckenridge, 2015). For example, there is a single Herman Beach graded to the southern outlet but more than a half dozen in northern North Dakota up to 70 m higher in elevation. Episodic outlet incision is recognized by the sequence of four strandlines graded to the southern outlet at successively lower elevations (Herman, Norcross, Upham, Tintah: Fisher, 2005) and terraces within the spillway (Fig. 3).

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Examples of strandlines are shown on Figure 4 from the east side of the basin, southeast of Grand Forks (Fig. 1) and on the west side of the basin, west of Fargo, ND. There is a striking

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difference in appearance of the strandlines between the two sites with seemingly many more strandlines on the west-facing slope (Fig. 4A). Black dots mark successively lower water planes in both DEMs and the number of dots between the Tintah Beach and highest strandline

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(Herman?) is instead similar on both maps. Strandlines below the labeled Norcross Beach west of Fargo are very subtle; resembling the Norcross strandlines described by McMillan and Teller

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(2012) in southern Manitoba that they suggest are storm beaches. The much larger and complex strandlines in Figure 4A offer greater opportunity for reconstructing paleo water planes and understanding littoral processes that reflects control by sediment supply, fetch, and aspect.

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The strandlines in Figure 4A are currently under study as they offer exceptional resolution of distinct water planes with large strandlines.

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Strandline chronology of the Herman through Tintah strandlines has been determined from primarily OSL dating because wood has not been found within them, and which is not from a lack of effort—presumably any wood was oxidized. While bones have been dated from the Campbell Beaches (e.g., Nielsen et al., 1984) most radiocarbon dates are from underlying basal organics (often peat) in wetlands dammed by beach ridges (e.g., Bjorck and Keister, 1983;

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Teller et al., 2000). Available OSL ages for the Lockhart Phase are shown on Figures 2, 4C. Those by Lepper et al. (2011, 2013) cluster tightly and are from the beach type area ~100 km from the southern outlet. Those dated by Teller et al. (2018) in southern Manitoba are weakly developed, contain poorly-sorted sediment, and are 400 km from the southern outlet where the higher number of beaches makes water plane assigned more difficult. Most errors associated with OSL beach ages are ± ~10% (Aitken, 1998). However, Lepper et al. (2011, 2013) and Liu et al. (2014) rely on a methodological approach in Lepper et al. (2011) that also reports just analytical error (~3%; Fig. 4C). A test of the veracity of the OSL age is shown in Figure 4C,

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where the geomorphic record of Lockhart strandlines records a regression and thus decreasing ages with decreasing elevation (Herman through Tintah strandlines). Independent support for this relationship is the decreasing gradient of lower elevation strandlines reflecting lower rates of GIA through time, a observation made by all who have mapped these strandlines, and most recently by Breckenridge (2015). Thus strandline age correlates with elevation, which is illustrated on Figure 4C where OSL ages decrease in age down slope. The OSL record within the gray panel (Fig. 4C) is the time window available for strandline formation. Its older bracket is deglaciation of the Big Stone Moraine and initiation of Lake Agassiz (Lepper et al., 2007) and

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the younger bracket is the oldest in-situ dated terrestrial organics from the Redwood Loop site (Fisher et al., 2008) that grew after lake level fell, signaling the end of the Lockhart Phase. The

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OSL ages from southwest of Fargo (Fig. 4C) are from sand deposited subaerially during the Moorhead Phase (Liu et al., 2014). Note that nearly all of the OSL ages fall within the gray panel, and the ages with their less conventional reported error assignment agree with the

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geomorphic strandline record, and radiocarbon dated stratigraphy.

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During the Lockhart Phase large influxes of meltwater from catastrophic drainage of ephemeral glacial lakes to the west of Lake Agassiz entered the lake through the Sheyenne, Pembina and Assiniboine rivers, and is further discussed in section 1.8. Terraces (a few mapped on Fig. 3)

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within the southern outlet spillway (also known as River Warren) may be associated with some of the strandlines and have suggested to some workers that outlet incision was in response to knickpoint migration (Matsch, 1983). Two sediment cores through lacustrine and gravel

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sediment into shale bedrock within the southern outlet record a minimum two episodes of outlet activity (Fisher, 2003). A thin lacustrine mud interbedded between gravel in core BVF-99 (Fig. 3) was interpreted to record the Moorhead Phase when the southern outlet was abandoned. 1.4 Moorhead Phase

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The Moorhead Phase began when lake level dropped below the southern outlet, and ended when a transgression (e.g., Teller, 2001) returned flow to the southern outlet (Fig. 2). The chronology and meltwater routing is uncertain where the lake drained at this time to cause the ~90 m drop in lake level (Breckenridge, 2015). Water levels fell from the Tintah Beach to a short distance north of Grand Forks, ND (Dilworth and Fisher, 2018). Drainage through eastern outlets into the Great Lakes basin following abandonment of the southern outlet was initially proposed by Upham (1895). However, few strandlines have been mapped on the mostly bare Canadian Shield bedrock, which has resulted in previous workers projecting strandlines from

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the main Agassiz basin eastward to low passes along the subcontinental drainage divide separating Hudson Bay and the Great Lakes basin. Historically, possible outlets have been described south and north of the Kaiasch Moraine (Elson, 1967; Zoltai, 1967; Teller and Thorleifson, 1983), here simply referred to as the Kam and Kelvin outlets respectively (Fig. 5). Detailed field investigations documenting any of the Kam outlets to the start of the Moorhead Phase have never been reported. The most recent projection of strandlines to any possible Kam outlets was by Breckenridge (2015), but the results were much the same as earlier work. Though Breckenridge emphasized that any early Moorhead flow would have begun at the

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lowest Tintah level. The outlets north of the Kaiasch Moraine have been associated with the younger Emerson Phase drainage, which is discussed in section 1.5. The original reconstruction

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of a > 40 m water level drop (Broecker et al., 1989) explained by eastward routing of Agassiz to the North Atlantic is problematic because nowhere along the drainage divide is there a col with enough relief to explain the drop in lake level. The low lake level reached during the Moorhead

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below any cols on the eastern drainage divide.

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Phase just north of Grand Forks, North Dakota (Fig. 1; Dilworth and Fisher, 2018) projects well

The drop in water level at the start of the Moorhead Phase is recorded by: 1) the transition to mud from gravel in lakes of the southern outlet, 2) the youngest age from a strandline

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abandoned when lake level fell, 3) the oldest ages from within the basin exposed when lake level fell, 4) deglacial ages where the new outlets opened, and 5) the age of freshwater signals in marine sediment. The youngest wood within basal gravel from the southern outlet spillway is

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12,550–12,720 cal yr BP (10,675±60 14C yr BP), which is a maximum age for when the southern was abandoned. The Tintan Beach is the lowest strandline graded to the southern outlet that is above the Campbell strandline and that records the end of the Moorhead Phase. Only two OSL ages are available for the Tintah Beach and they average 13.5 ± 1.1 ka (Lepper et al., 2011). The areal gap of strandlines between the Tintah and Campbell beaches was used by

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Breckenridge (2015) to argue for rapid lake level fall at this time. The Tintah Beach has not been mapped to any eastern outlet. High resolution LiDAR images (Fig. 4B) show possible weakly developed strandlines within this gap area, but require additional work to confirm. After lake level fell organic sediment could accumulate on the freshly exposed landscape by fluvial deposition and plant growth at elevations below the Tintah Beach. As discussed at length by Fisher and Lowell (2006) and Fisher et al. (2008), the oldest in situ terrestrial organic material is the best minimum estimate for lake level fall, but plant colonization time must be taken into account. The oldest root from a drowned forest is 12,540–12,740 cal yr BP (10,710±75 14C yr

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BP, ETH-32674) and the youngest in situ organics (12,250–11,760 cal yr BP 10,000±70 14C yr BP, ETH-32679) are ~1100 years younger. Based on the analysis of available radiocarbon dates (fig 8 in Fisher et al., 2008) lake level fall was ~12.7 cal kyr BP plus time for colonization (~100 years?), which gives an age close to the southern outlet chronology (Fig. 3; Fisher, 2003), and is overlapped by the OSL ages from the Tintah Beach, and oldest peat (12,150– 12,860 cal yr BP [10,700±140 14C yr BP, WAT-1910]) from the Rainy River Lowland (Bajc et al., 2000) in northwest Ontario (R on Fig. 1) that grew after lake level fell.

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Streams flowing across the exposed lake basin during the Moorhead Phase deposited fluvial sand of the Poplar River Formation (PRF) on lacustrine sediment of the Brenna Formation of

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Lockhart age. The PRF was buried by lacustrine sediment of the Sherak Formation (Ardnt,

1977) during the Moorhead transgression leading to reopening the southern outlet. From two sites just south of Fargo, ND, the PRF is expressed as a laterally continuous ribbon sand, with

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an average age of 13.1 ± 1.6 ka (Fig. 2) that may record a regressive sand deposited when lake level fell (Liu et al., 2014). Radiocarbon dates from the Poplar River Formation that are primarily

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channel sands elsewhere in the Fargo area are younger, beginning around 11,350 cal yr BP (Yansa and Ashworth, 2005).

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Before a new outlet can open to cause abandonment of the southern outlet, ice must deglaciate lower outlets. Minimum ages from lakes and cosmogenic ages on boulders were used by Lowell et al. (2005, 2009) to suggest that ice had not retreated far enough eastward to open lower

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spillways into the Great Lakes basin. More recently, Leydet et al. (2018) obtained older cosmogenic ages from boulders that suggest deglaciation was as early as 14.0 ± 0.4 ka to route meltwater eastwards before the start of the YD stadial. However, as most recently pointed out by Breckenridge (2015), none of the identified possible Kam spillway routes (Fig. 5) have the relief necessary to drop the lake ~90 m to cause the flood invoked by many earlier researchers.

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Deglaciation of eastern outlets may have initially dropped water levels at the start of the Moorhead Phase if the magnitude of the drop exceeded the vertical gap between the Tintah and Campbell beaches. A search for thick sedimentary sequences offshore of Thunder Bay, Ontario (Fig. 5) failed, suggesting to Voytek et al. (2012) that no significant megaflooding occurred in this direction. If initial drainage was to the east before the start of the YD and a connection to the North Atlantic Ocean through the Great Lakes basin was not available, routing could have instead been out the Lake Superior basin through the Brule spillway or Chicago outlet to the Mississippi River (B and C, respectively Fig. 1) if ice had retreated far enough before the

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Marquette readvance (Lowell et al., 1999). Such routing would not interrupt the meltwater signal in the Gulf of Mexico. Because any eastward drainage cannot explain the complete drop in lake level during the Moorhead Phase, drainage out of the northwest outlet has been proposed (Teller et al., 2005; Breckenridge, 2015; Leydet et al., 2018).

A northwest outlet was considered by Upham (1895), and Elson (1967) identified a series of channels in the upper Churchill River valley (Fig. 6), now termed the Wycherely channels mapped by Fisher and Souch (1998). Further to the northwest is the large Clearwater lower-

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Athabasca spillway (CLAS) that was initially interpreted as the northwest outlet for Lake Agassiz (Smith and Fisher, 1993; Fisher and Smith, 1994a, b), and is further described in section 1.11.

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Strandlines well above the head of the spillway were assumed to be a part of Lake Agassiz and labeled as Norcross. Since then, Fisher (2005) retracted the Norcross interpretation, primarily because the Norcross strandline does not extend north of southern Manitoba (Fig. 7) (cf. Fisher

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and Lowell, 2012). Instead the high strandlines above the head of the CLAS would now be associated with glacial Lake Churchill in the upper Churchill valley, a product of merging

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formerly mapped glacial Lake McMurray in Alberta and glacial Meadow Lake in Saskatchewan (Fisher et al., 2009), and drainage of glacial Lake Churchill is the source of water that eroded the CLAS. Glacial Lake Churchill strandlines in this region remain undated. Minimum deglacial

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ages of 11.3 ± 0.1 to 11.0 ± 0.2 cal kyr BP from the youngest moraine (Fort Hills Moraine) in the Fort McMurray area are a minimum limit for when the flood occurred (Fisher et al., 2009). If a hundred years are invoked for a lag time before plant colonization occurred, then this

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chronology overlaps with OSL ages (11.9 ± 1.0, 11.8 ± 1.0, 11.5 ± 0.7 ka) of sand conformably overlying the oldest flood gravel on the Mackenzie delta (Murton et al., 2010). Ages on wood associated with flood gravel in the Fort McMurray area, and in the resultant delta built into Lake McConnell (Smith and Fisher, 1993) are expected to be older than the flood because the Athabasca River can transport wood from an older landscape westward, and the oldest wood

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ages (cf. Fisher and Lowell, 2012) are Moorhead Phase aged. If this flood occurred at the start of the YD, then the older wood should predate the YD. While extant geomorphic and chronologic data from the Agassiz basin thus far do not support northwest drainage at the start of the YD (Fisher and Lowell, 2012), it has become the default outlet following initial drainage to the east (Brenkenridge, 2015; Leydet et al., 2018). Calculations demonstrating the feasibility of the Moorhead Phase to have resulted from enhanced evaporation as a closed basin (Lowell et al., 2013) have not been well received (Carlson and Clark, 2012; Teller, 2013; Breckenridge, 2015).

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The available data from the Fort McMurray area of minimum flooding ages of 11.3 ± 0.1 to 11.0 ± 0.2 ka BP overlaps with the Preboreal Oscillation (PBO) (Hald and Hagen, 1998). Fisher et al. (2002) suggested that drainage through the CLAS caused the PBO through increased output of freshwater and sea ice export to the North Atlantic Ocean through Fram Strait, and these ages overlap with reinterpreted (Carlson and Clarke, 2012) ages from the Mackenzie delta (Murton et al., 2010).

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Marine records have also been used to determine meltwater routing from Lake Agassiz. Data from the Gulf of Mexico record abandonment of the south outlet (Broecker et al., 1989; Flower

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et al., 2004; Williams et al., 2010; Wickert et al., 2013) which requires diversion of meltwater elsewhere. Eastern drainage through the Great Lakes basin is supported by marine proxy data for freshwater beyond the St. Lawrence River (Carlson et al., 2007; Levac et al., 2015) and

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evidence for freshwater excursions closer to land (Rayburn 2011a,b; Cronin et al., 2012). Once the St. Lawrence River opened following drainage of Lake St. Lawrence (Rodrigues and Vilks,

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1994; Levac et al., 2015) drainage from the Great Lakes is rerouted eastwards along with meltwater runoff from a large portion of the LIS. Data from the Arctic Ocean basin has also been used to support drainage from Agassiz northwards at the start of the YD (Hillaire-Marcel et al.,

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2013; Keigwen et al., 2018) to explain freshwater inputs. Modeling of meltwater runoff by Tarasov and Peltier (2005, 2006) suggested that meltwater runoff from the Keewatin Dome of the LIS alone could explain freshening in the Arctic Ocean, but was criticized by Carlson and

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Clarke (2012) for their enhancing precipitation during the YD. Condron and Winsor (2012) supported northern versus eastern drainage as freshwater delivery was more effect in their modeling results at weakening circulation in the North Atlantic Ocean. Alternative sources of meltwater beyond North America are from the Greenland Ice Sheet that is much closer to sites of deep-water formation (Jennings et al., 2006) and drainage from the Baltic Ice lake

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(Muschitiello et al., 2016).

The routing of meltwater at the start of the Moorhead Phase is poorly understood. Available chronology from the Lake Agassiz basin supports an abandonment of the southern outlet close to the start of the YD, but field data revealing when routing transitioned is not available. Evidence for meltwater excursions at the start of the YD from marine records off shore of the Saint Lawrence and Mackenzie Rivers compete for drainage from Lake Agassiz. The lack of field evidence for strandlines extending to the CLAS spillway at the start of the YD may reflect

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their erosion by younger lakes (Breckenridge, 2015) or a readvance. Additional work is required to evaluate the megaflooding history from the Fort McMurray area and sites further north where other large lakes once existed (e.g., Smith, 1992; Lemmen et al., 1994), collect radiocarbon ages from Lockhart-aged beaches, and to collect chronology data from spillways identified in northwest Ontario.

The strandline record during the Moorhead Phase is poorly understood. Because the lake transgressed over the lake basin at the end of the phase, preservation potential of strandlines is

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low (Teller, 2001; Breckenridge 2015). Fisher et al. (2008) dated plane bedded sand they considered a Moorhead Phase beach (Ojata) as it was buried by fine-grain lacustrine sediment.

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More recently Teller et al. (2018) suggested that the Gimili Beach in southern Manitoba (G on Fig. 1A) may also be a Moorhead Phase beach. The previously undated Gimili Beach had always been assumed to be a Nipigon Phase beach because of its lateral extent (50 km) and

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low slope, but the much older than expected OSL ages (Teller et al., 2018, Fig. 2) complicates this understanding. Older sand may be present in the beach that is not representative of the

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geomorphic landform, similar to older ages encountered elsewhere in other Lake Agassiz

1.5 Emerson Phase

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beaches (Lepper et al., 2013).

Much of the earlier literature assumed that water drained eastwards during the Moorhead Phase, and the Emerson Phase began when flow returned to the southern outlet and ended

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when drainage switched back to the east. Water level at this time is recorded by the Campbell Beach, which is divided into Upper and Lower beaches. The Campbell beaches grade to the southern outlet and are projected to the eastern Kelvin outlets (Figs. 5, 8) and the Wycherely channels as the northwestern outlet (Fig. 1). If an eastern or northwestern outlet was active during Moorhead time, then it was GIA driving the transgression that returned flow to the

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southern outlet.

The duration of the Emerson Phase is perhaps best estimated by the range of ages on the Campbell Beach. Flow returning to the southern outlet is recorded by gravel overlying lacustrine mud and bedrock within the spillway (Fisher, 2003). The youngest piece of wood from the gravel is 10,520–11,080 cal yr BP (9460 ± 70 14C yr BP, AA38302), and is the most constraining maximum age for reopening the southern outlet. The OSL chronology for the undifferentiated Campbell Beach is shown in Figure 2, with each age listed in Table 2. There is a substantial

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suite of radiocarbon ages on the Campbell Beach (Table 2), some of which were discussed in Lepper et al. (2013) and agree with the OSL ages. Three of the OSL ages (samples AB0710, AB0711, E2; all not plotted on Fig. 2) are older than most Campbell ages and may result from Moorhead-aged sand around which the Campbell Beach formed, thus the younger OSL ages are considered more limiting and the E1 sample from Manitoba (Teller et al., 2018) agrees with the Campbell ages near and at the southern outlet type area (Lepper et al., 2013). Wood ages from within beaches are considered maximum ages for those beach deposits. Where lagoons formed behind Campbell beaches, the youngest basal dates from the lagoons are good

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estimates for the Campbell Beach age. Many of these ages in Table 1 overlap in time with the

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OSL ages, and the duration of the Emerson Phase may be estimated from 10.7 to 10.0 ka.

Deglacial ages from the Kelvin outlets region that overlap in age with the Emerson Phase are from Lowell et al. (2005, 2009) and Teller et al. (2005) from minimum limiting ages in lakes near

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moraines, and cosmogenic ages from boulders. In northern Saskatchewan, Schreiner (1983), Fisher and Smith (1994a), and Rayburn and Teller (2007) mapped Campbell Beaches

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northwards from Nipawin, Saskatchewan along the eastern side of the Cub and Wapawekka Hills (Fig. 6). Well-developed strandlines and strandline fragments north and east of the Wapawekka Hills were only associated with the Lower Campbell strandline by Rayburn and

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Teller (2007). Further to the northwest Fisher and Souch (1998) mapped the Wycherely channels, and assumed they drained Lake Agassiz at the Campbell level to the northwest into a small regional lake dubbed Lake Wagtufro, which drained into the head of the Clearwater lower-

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Athabasca spillway (CLAS, Fig. 6). This reconstruction requires that the CLAS spillway had already formed. Radiocarbon dates and age assignments based on sedimentation rates from lakes adjacent to the Wycherley channels and Lake Wagtufro gave an age of 10.2–10.8 cal kyr BP for cessation of northwest drainage from lake Agassiz, which overlaps with OSL ages (Lepper et al., 2013) from the Campbell Beaches elsewhere in the lake basin (10.0–10.7 ± 0.9

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ka) and dates from boulders (Kelly et al., 2016) in the Kelvin outlets (11.1 ± 0.4–10.8 ± 0.2 ka). Whether flow first occurred to the east, northwest or south is unknown, but flow was eventually captured by the Kelvin outlets ending the Emerson Phase and starting the Nipigon Phase. Synchronous flow through three different outlets was possible for a limited time, similar to the three-outlet history of mid-Holocene Lake Nipissing in the upper Great Lakes basin (Thompson et al., 2011).

1.6 Nipigon Phase

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Outflow through the eastern Kelvin outlets was into a highstand of Lake Nipigon known as Lake Kelvin. Water then passed through the Nipigon channels into the Lake Superior basin (Fig. 5). Thick sequences of sediment within troughs of Lake Superior just offshore of Nipigon record high sediment loads carried through the Nipigon channels (Teller and Mahnic, 1988; Gary et al., 2012) as evidence for catastrophic flooding from Lake Agassiz.

Detailed paleogeographic maps of ice margin retreat and meltwater routing through successively lower outlets west (Figs. 5, 8), then north of Lake Nipigon are associated with the

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suite of 14 strandlines (Fig. 2) recording episodic regressive lake level fall in the southern basin of Lake Agassiz (Leverington and Teller, 2003). Opening each successively lower outlet has

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been postulated to result in a brief catastrophic flood until lake level stabilizes at a new sill. The resulting strandlines following the successive drop in lake level are all younger than the Campbell beaches—but the only age control is based on their tilt, becoming younger with lower

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gradients. Some age estimates have been proposed using this relationship (Yang and Teller, 2012). As ice recession continued northward, drainage was eventually diverted from the Great

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Lakes basin as Lake Agassiz merged with glacial Lake Ojibway when lake level transitioned between The Pas and Gimili levels (Fig. 2; Leverington and Teller, 2003). Using varve thickness records from the Lake Ojibway basin, Breckenridge et al. (2012) estimated this occurred at 9040

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cal yr BP.

1.7 Ojibway Phase and final drainage

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As Lake Agassiz was further expanding northwards against the receding ice sheet margin, further east proglacial Lake Ojibway was developing over northeastern Ontario and northwestern Quebec. With an initial stage known as glacial Lake Barlow (Vincent and Hardy, 1979) glacial Lake Ojibway formed between the receding ice margin and the drainage divide between Hudson Bay and the Atlantic Ocean and lasted for about 2100 years as determined

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from varve counts (Breckenridge et al., 2012). Overflow was south to the Ottawa River through the Kinojévis outlet (KinO, Fig. 1). Proglacial lakes Agassiz and Ojibway merged at about 9040 cal BP (Breckenridge et al., 2012), with the Lake Agassiz name preserved. Fluctuations in water level are recorded by scarps (Roy et al., 2015), and Roy et al. (2011) described a 0.6 m thick laminated unit with rounded clay balls and other clasts interbedded between lacustrine and marine sediments recording final drainage, building on earlier work by Skinner (1973). Additional analysis on this unit by Breckenridge et al. (2012) and Stroup et al. (2013) alternatively suggest a Lake Ojibway sedimentary deposit predating final drainage of the lake.

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The final drainage of Lake Agassiz is hypothesized as one or numerous subglacial channels forming beneath the thinning ice sheet over modern Hudson Bay, which resulted in catastrophic drainage in one or a few stages (Clarke et al., 2004). Geomorphic indicators for the flooding include meltwater channels kilometers in width, fields of giant current ripples and arcuate iceberg scours recording unidirectional flows (Josenhans and Zevenhuizen, 1990; Lajeunesse and St-Onge, 2008). Veillette (1994) worked in the Lake Ojibway basin and along with others, estimated final drainage at 8.0 14C kyr BP (~8.86 cal kyr BP) based on marine shells associated

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with the Tyrell Sea transgression following final drainage of Lake Agassiz. The estimate for flood timing from marine records by Barber et al. (1999) from carbonate marine shells in the

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sedimentary record of the flood in Hudson Strait and Hudson Bay is 8.45 cal kyr BP, and the 8.2 ka event isotopic record from the Greenland Ice Sheet may be associated with final drainage

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(see discussion in Thomas et al., 2007).

1.8 Floods into Lake Agassiz

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Catastrophic drainage events into Lake Agassiz from ice-marginal lakes hundreds of kilometers to the west are recorded by large spillways ending at shorelines, and by large fans deposited into the lake (Figs. 1, 9). Many of these events and characteristic erosional and depositional

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features are reviewed in Kehew et al. (2009) with peak flood discharge estimated at ~0.8 Sv (Table 3). The general sequence of events involves meltwater flowing along and away from ice margins within the Missouri drainage basin when ice was along the Missouri Coteau and

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receding northwards. The spillways west of the James spillway record some of the earliest flows (Fig. 9) around 13,800 cal yr BP (Brophy and Bluemle, 1983). On Figure 9 selected ice margins are shown as broad dashed white lines to illustrate how their position controlled formation of icedammed lakes, and with subsequent recession of the ice margin, how new drainage routes evolved. For example, when ice retreated northwards from the Heimdal ice margin, flow was

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initially through the James spillway until lower land was uncovered; flow was then captured and diverted down the Sheyenne River. Both rivers end at Lake Dakota (Fig. 9). Brophy and Bluemle (1983) describe initial flow in the Sheyenne River as typical braided river flow before the spillway developed, evidenced by high elevation terraces of outwash. With further ice margin recession water was trapped at the margin of the Souris Lobe in glacial Lake Souris, and to the east in Lake Minneoaukan. While the triggering mechanisms for what initiated lake drainage and spillway formation is uncertain, headwater migration of knickpoints to the lakes, melting of stagnant ice near the sill, and large inputs of meltwater from upstream have all been

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proposed (cf. Kehew and Lord, 1986; and Lord and Kehew, 1987). The James and Sheyenne spillways that enter Lake Dakota record high-magnitude flows before the Des Moines Lobe had retreated north of the Big Stone Moraine. Once Lake Agassiz began forming, sometime after the ice margin had receded north of the present day Sheyenne River, a large flood eroded the Sheyenne spillway and deposited the large Sheyenne fan into the southern basin of the lake. Drainage from ephemeral Lake Souris is often invoked to have caused this event (Kehew and Lord, 1986). The Sheyenne fan area is ~2000 km2 and consists of a wedge of sediment fining

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outwards from gravel and sand, to silt and clay (Brophy, 1967).

With continued ice recession northwards, Lake Souris transitions into glacial Lake Hind (Fig. 9).

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Detailed mapping and reconstruction of the lakes paleogeography can be found in Sun and

Teller (1997). Once the Assiniboine Lobe had retreated far enough northwards, drainage from Lake Hind was to the southeast through the Pembina spillway. The spillway morphology records

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large flows, although the Pembina Fan is much smaller than the Sheyenne Fan. The width of the spillway is larger than the Sheyenne’s, suggesting that much of the eroded sediment

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entering the lake was transported away by littoral processes. Further recession of the Assiniboine Lobe opened a short spillway (box 4, Fig, 9) and sediment was deposited in the Assiniboine reentrant. Continuing ice margin retreat increased the Assiniboine River discharge,

Assiniboine delta.

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with flow across the Lake Hind basin and into glacial Lake Agassiz contributing to growth of the

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The Assiniboine spillway extends much further north, along the western side of the hills that make up the Manitoba Escarpment (Fig. 7). Many spillways come off these headlands, and are tributaries to the Assiniboine River and spillway, indicating that lower elevation land to the east and north was ice covered when the Assiniboine delta was forming. The westward flowing spillways were abandoned once lobes of ice retreated from the reentrants along the Manitoba

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Escarpment, presumably in a south to north direction. Deltas from local, small ice-marginal lakes are present within the Swan River reentrant and along the eastern side of the Porcupine Hills and Duck Mountain (Fig. 7) (Nielsen, 1988). Short strandline segments associated with the local lakes have been mapped above the upper Campbell beach in this area (Nielsen, 1988). The highest Lake Agassiz beach north of Duck Mountain mapped in these reentrants is the Upper Campbell beach, indicating that when ice retreated from the Manitoba Escarpment lake level was below the Campbell beach during the Moorhead Phase. Further north, two large deltas developed where the Saskatchewan River entered proglacial Lake Saskatchewan (Fig.

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7). The Fort Á La Corne delta formed when Lake Saskatchewan drained southward through the Assiniboine spillway system, after Lake Assiniboine (Wolfe and Teller, 1995) had drained (Fig. 7). Set below and eastward of the Fort Á La Corne delta is the Nipawin delta. It formed at a level lower than glacial Lake Saskatchewan; Elson (1967) and Christiansen (1995) associate the delta with Lake Agassiz.

1.9 Southern outlet flows The southern outlet of Lake Agassiz (Figs. 1, 3) is the oldest and most studied outlet of the lake.

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At the Herman and Norcross level, flow was through the Cottonwood and central spillway (Fig. 3) and the Cottonwood channel is abandoned once lake level dropped to the Tintah level. The

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central spillway complex at the Tintah and Upham levels was very broad, up to 10 km width. Discharge estimates through the southern outlet range from 0.4 to 0.7 Sv based on a variety of techniques including boulder transport (Fig. 10A); see Fisher (2004) for more information. As

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the lake area increased through time, baseline discharge would have increased through the southern outlet. As discussed in section 1.8, ice retreat allows successively northern spillways

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to form, delivering catastrophic floods into the lake along its western side. While discharges have been estimated between 0.058 to 0.8 Sv, water levels remained at the Herman lake level when the Sheyenne and Pembina deltas were deposited. The reasoning for this is that the

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Herman Beach extends north from the southern outlet to the Pembina delta. If the southern outlet was incised to a lower lake level when a delta formed from a catastrophic flood, then a lower lake level would appear north of the delta. Elson (1967) noted that the highest Herman

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strandline ends near the Pembina delta, and strandlines are difficult to map across the Assiniboine delta. The Norcross is the highest strandline north of the delta. Thus, it is likely that some combination of increasing baseflow and catastrophic floods through the Pembina and Assiniboine spillways initiated incision of the southern outlet, causing lake level to fall to the Norcross level. The River Warren spillway continued to evolve across Minnesota (cf. Johnson et

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al., 1998), and now is occupied by the Little Minnesota River that eventually joins the Mississippi River (Fig. 1).

1.10 Eastern outlet flows Flow out the Kelvin outlets was estimated by Teller and Thorleifson (1983) using channel geometry and boulder dimensions to be 0.1–0.2 Sv over a few years. Kelly et al. (2016) calculated lower discharges for the Roaring River site, using conservative estimates of channel width (Table 1) and boulder dimensions (Fig. 10B). A maximum discharge from the Mundell

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Lake site within the Pillar outlet complex is 0.12–0.016 Sv. At this site large giant current ripples (Fig. 10C) were constructed from boulders up to 1.8 m b-axis. Preservation of these bedforms adjacent to the lake likely indicates a short-lived megaflood event followed by a much lower discharge steady flow. Past workers have discussed or implied such a flooding history controlled by the receding ice margin opening progressively lower outlets further to the north as the sub-continental drainage divide is uncovered (Elson, 1967; Zoltai, 1967; Teller and Thorleifson, 1983; Leverington and Teller, 2003; Kelly et al., 2016). Leverington and Teller (2003) estimated water level drops of 7 to 58 m and corresponding water volume releases of

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1900–8100 km3 as new outlets became established. With better topographic data and ice

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margin positions, more accurate estimate of paleodischarge can be determined.

The various Nipigon outlets (Fig. 5) and on-shore fan sediments into the Lake Superior basin were described by Teller and Mahnic (1988). Evidence for flooding include sandy turbidities up

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to 0.65 m thick and large trough crossbeds. Gary et al. (2012) used marine seismic methods offshore of Nipigon, and described deeply incised bedrock channels in the bay overlain by

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coarser-grained hummocky material interpreted as sediments deposited during high energy events, which are in turn overlain by finer-grained sediments reflective of non-catastrophic

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events.

1.11 Northwest and Ojibway outlet flows

The main spillway and flood deposits associated with the Clearwater lower-Athabasca Spillway

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(CLAS) were described in detail by Fisher and Smith (1993, 1994), Smith and Fisher (1993), and Fisher et al. (1995). A peak discharge estimate of ~2.4 Sv used the slope-area equation with mean velocities of 13.2 ms-1. Assumptions of duration assumed a draw down of 46 m and an areal extent of 350,000 km2 (Teller and Clayton, 1983), but the area would be considerably

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less for the much smaller glacial Lake Churchill.

Sediments deposited from megaflooding or catastrophic flows have been described from glacial Lake Agassiz and from many other lake drainages (cf. Martini et al., 2002; Burr et al., 2009) that are commonly associated with deglaciation. Highly erosive flows result in large channels or sets of channels eroded into bedrock (Fig. 11A) or glacigenic sediment (Fig. 11B). Sediment transport is high and flood deposits may be found within the spillway where velocities decrease from flow separation or expansion. Flood deposits are found in the CLAS where flow occurred in upper scoured zones and at its northern end where flow slowed and bifurcated around the Fort

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Hills (Fig. 6) (Fisher and Smith, 1993; Fisher et al., 1995) with examples described next. Often very-large boulders are transported or left behind as lags (Fig. 11D), while large piles of usually subangular to subrounded boulders record bedload transport, but they are too large to crush for most gravel operations (Fig. 11E). In areas of flow separation large sets of cross-beds develop (Fig. 11C) consisting of graded beds with open- and closed-work characteristics (Fig. 11H). Some clast-supported gravel is open-work (Fig. 11F) requiring that all other material remained in suspension, and elsewhere gravel clasts are supported by a matrix resulting in a bimodal particle size distribution (Fig. 11G, I). Intraformational clasts consisting of loosely consolidate oil

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shale bedrock, till and glaciolacustrine sediment range in size from pebbles and boulders (Fig.

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11J) to large rafts 15 m along an exposed axis.

The Wycherely channels (Fig. 6) near the headwater of the Churchill River were the northwest outlet active at the Campbell level (Fisher and Souch, 1998; Fisher, 2007). Routing from the

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Wycherely channels was into Lake Wagtufro and then through the CLAS. One channel was briefly visited on foot in 2003, with access by float plane, but was not surveyed in the field.

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Using topographic maps and aerial photographs (Fig. 6) hydrological variables are estimated at ~400 m width and ~15 m depth. Individual boulders up to a few meters diameter were observed as were deposits of boulders of ~1 m diameter. Assuming velocities of between 3 to 12 ms-1

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(methodology in Fisher, 2004), each channel may have experienced peak discharges of 0.018– 0.072 Sv. If all channels were active at once discharge could have been 0.08 to 0.29 Sv. Such a value is higher than Lake Agassiz baseline flow estimated by Teller et al. (2002) at 0.034 Sv for

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the Campbell level.

Veillette (1994) estimated paleoflows for the Ojibway Phase at different river levels downstream of the Kinojévis outlet in the Ottawa River valley. Water depths were estimated from fluvial erosional marks (e.g., potholes) cut into bedrock and the Manning equation was used to

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estimate discharge. Results ranged from 33,000–65,000 m3s-1, which are more conservative than an earlier estimate of 0.2 Sv from Lewis and Anderson (1989). 1.12 Final Drainage There is some uncertainty about the final drainage of Lake Agassiz, whether it was in two stages or one (Leverington et al., 2002; Teller et al., 2002; Clarke et al., 2004). Clarke et al. (2004) advocate for subglacial drainage that could accommodate an initial drop in water level from the Kinovéjis to Fidler levels and then final drainage. Leverington et al. (2002) calculated a

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total drainage volume of 163,000 km3, with an initial drainage of 113,100 km3 followed by 49,900 km3 from the Fidler level. Resulting discharges were about 5 Sv if drainage was in one stage, or 3.6 Sv and then 1.6 Sv if a two-stage drainage. Clarke et al. (2004) estimate a halfyear flood duration, with a discharge of about 5 Sv, and a maximum sea level rise of 0.19 m. The final drainage either in one or two parts is explained as the forcing mechanism for the 8.2 event recorded in the Greenland ice cores and elsewhere (Barber et al., 1999). 1.13 Summary

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The history of drainage from Lake Agassiz extends over a nearly 6000 year period as the lake evolved across the center of the North American continent. With water trapped at its

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southernmost extent at the sub-continental drainage divide, the large southern outlet was

episodically incised as recorded by numerous beaches graded to it, and boulder-armored terraces within it. Most recent discharge estimates for the southern outlet are up to 0.36 Sv.

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Large floods from deglacial lakes further west deposited large fans into Lake Agassiz with estimated discharges of up to 0.8 Sv. The receding ice margins controlled where and when

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lower outlets to the east and north could open. Uncertainty remains where the lake drained the first time the southern outlet was abandoned at the start of the Moorhead Phase, which is approximately coeval with the start of the YD. While the ice margin may have retreated far

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enough east to permit drainage into the Great Lakes basin, spillways with vertical relief to explain the ~90 m drop in elevation remain to be documented. Drainage out a northwest outlet is now viewed as the only other candidate with discharges estimated at ~2 Sv. However, only

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the much younger Campbell beaches have been mapped to a northwest outlet unless an earlier phase of Lake Agassiz extended to the head of the CLAS spillway and evidence for it has been removed from a readvance. When flow returned to the southern outlet during the Emerson Phase there is evidence that a northwest, eastern (Kelvin) and southern outlet were active simultaneously, or in close succession with the lake at the Campbell levels. Drawdown of water

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during the Nipigon Phase through successively lower Kelvin outlets resulted in paleodischarge estimates of 0.1 to 0.2 Sv that entered the Lake Superior basin. Lakes Agassiz and Ojibway merged with continued retreat of the ice margin, and outflow through the Kinojévis outlet into the Ottawa River system was estimated at 0.033–0.065 Sv. Final subglacial drainage of the lake occurred as ice thinned over Hudson Bay with drainage in one or two stages. With discharge estimated at ~5 Sv, sea level was raised 0.19 m and this megaflood freshened the surface of the North Atlantic and is often cited as the cause of the 8.2 ka event.

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Conflict of interest There is no conflict of interest with this manuscript.

Acknowledgements SRTM data for Figures 5 and 9 provided by OpenTopography.com. Thomas Valachovics, Brian

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Samsen, Tom Lowell, and two journal reviewers are thanked for their constructive comments.

TABLE 1 Select Volume and Area of Lake Stages

Lake Phase

Area (km2) Volume (km2)

Herman

Lockhart

134,000 10,900

Norcross

Lockhart

166,000 13,300

Tintah

Lockhart

184,000 15,700 117,000 10,800

Late Moorhead

185,000 19,700

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Early Moorhead

Emerson 263,000 22,700

McCauleyville

Nipigon

219,000 16,400

The Pas

Nipigon

151,000 4600

Kinojévis

Ojibway

841,000 163,000

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Upper Campbell

Ojibway

408,000 49,900

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Fidler

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Lake stage

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(modified from Leverington et al., 2000, 2002)

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Comment

10.0 ± 0.9 (± 0.2) ka 10.3 ± 0.9 (± 0.2) ka 10.7 ± 0.9 (± 0.3) ka 10.6 ± 0.9 (± 0.2) ka 10.7 ± 0.9 (± 0.2) ka 12.0 ± 1.0 (± 0.3) ka 13.6 ± 1.1 (± 0.3) ka 10.1 ± 0.7 11.7 ± 1.4

Undifferentiated Campbell Beach (CB) Undifferentiated CB Undifferentiated CB near Campbell, MN Undifferentiated CB near Campbell, MN Undifferentiated CB near Campbell, MN Undifferentiated CB near Campbell, MN Undifferentiated CB near Campbell, MN Upper Campbell (UC) Beach UC Beach

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Age1

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Sample ID dated material Age of sand deposition—OSL KG0505 fine sand KG0505 very-fine sand AB0704 fine sand KL0902 fine sand KL0903 fine sand AB0710 fine sand AB0711 fine sand E1 fine sand E2 fine sand

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Table 2. Campbell Beach Chronology

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Wood within beach sediment2 TO-4855 wood frag. 10,250–10,770 (9340±90) youngest age in beach sediment GSC-5357 wood 11,320–12,030 (10,100±90) transported wood in UC beach GSC-5330 wood 11,210–11,710 (9940±80) in sand of LC Beach GSC-4732 wood 11,210–11,710 (9940±80) in sand of LC Beach

1For

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Maximum ages from organics underlying beach or resultant lagoon unknown peat 10,250–10,790 (9350±100) peat immediately below beach sand TO-4856 charcoal 10,280–10,720 (9330±80) basal lagoon sediment TO-4873 organic frag 10,280–11,070 (9380±90) from silty clay in base of lagoon TO-4284 peat 10,500–11,100 (9460±90) base of lagoon GSC-5731 wood 11,110–12,120 (9990±160) underlying LC Beach (max age) BGS-1305 plant detritus 11,150–11,820 (9920±110) soil buried by offshore bar or spit BGS-1304 wood 10,650–11,720 (9750±170) beneath UC Beach WAT-1934 wood 10,490–11,220 (9530±140) beneath UC Beach AA-50803 wood 10,570–11,090 (9490±70) youngest age from valley fill beneath UC Beach

Reference Lepper et al., 2011 Lepper et al., 2011 Lepper et al., 2013 Lepper et al., 2013 Lepper et al., 2013 Lepper et al., 2013 Lepper et al., 2013 Teller et al., 2018 Teller et al., 2018 Teller et al., 2000 McNeely & Nielsen, 2000 McNeely & Nielsen, 2000 McNeely & Nielsen, 2000 Bjorck & Keister 1983 Teller et al., 2000 Teller et al., 2000 Teller et al., 2000 McNeely & Nielsen, 2000 Bajc et al. 2000 Bajc et al. 2000 Bajc et al. 2000 Fisher et al., 2008

OSL ages, propagated error shown with analytical error in parentheses, cf. Lepper et al. (2011), ages in parentheses

2 Radiocarbon

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TABLE 3 Reconstructed flows from spillways associated with Lake Agassiz Spillway

Paleohydrology

Souris Spillway

v = 2.9–11.7 ms-1, Q = .0.58–0.8 Sv ms-1,

6.4–30.4

Reference ms-1;

Lord & Kehew, 1987

Southern Outlet

v = 8.4

Q = 0.1–0.36 Sv

Fisher, 2004

Kelvin Outlets

Q = ~0.1–0.2 Sv

Kelvin Outlets

1900–8100 km3 released from 7–58 m drops in water level

Teller & Thorleifson, 1983 Teller & Leverington, 2004

Kelvin Outlets

0.012–0.024 Sv; 0.066–0.16 Sv

CLAS

Q = ~2.16 Sv

Wycherley channels

Q = 0.08–0.29 Sv

Kinojévis Outlet

Q = 0.03–0.06 Sv

Subglacial

Q = 0.5–5 Sv

Kelly et al., 2016

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Fisher et al., 2002 this paper

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Veillette, 1994

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Leverington et al., 2002; Clarke et al., 2004

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Figure Captions Figure 1. Overview of the maximum extent of Lake Agassiz occupied throughout its nearly 6000 year history. Panel A modified from Leverington and Teller (2003) and panel B modified from

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Dilworth and Fisher (2018).

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of ro -p re lP ur na Jo Figure 2. Lake Phase history of glacial Lake Agassiz with age control shown for the three oldest phases. Note change in elevation scale at 9.9 ka.

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Figure 3. Upper segment of the southern outlet spillway by the town of Browns Valley. Terraces armored with boulders are common along the spillway with a few examples mapped adjacent to

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the streamlined islands within the spillway. Core panel is modified from Fisher (2003).

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Figure 4. Strandline examples from glacial Lake Agassiz. See Figure 1A for location. A) West facing slope in Minnesota illustrating a high number of individual strandlines. Black dots represent distinct water planes, white dots possible strandlines within the Tintah gap area.

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Center of DEM is at 47.6671°N, 96.3375°W. B) Strandline examples along an east facing slope previously studied by Lepper et al. (2011) as the Wheatland transect. Black dots represent distinct water planes. Center of DEM is at 46.9188°N, 97.3859°W.

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Figure 5. Hillshaded DEM of the eastern outlet regions identifying major outlets. Flow from the Kelvin outlets was buffered by Lake Kelvin before passing through the Nipigon outlets into the

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Lake Superior basin.

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Figure 6. Hillshaded DEM of the northwestern extent of Lake Agassiz within the upper Churchill River basin. See Figure 1 for location. At the Campbell level flow was out a series of the four Wycherley channels (B) into a broad lowland occupied by modern Lakes Wasekamio, Turnor and Frobisher (Lake Wagtufro) and then into the already formed Clearwater lower-Athabasca spillway (CLAS) and eventually the Mackenzie River. (C) View (white arrow on B) of one channel looking east towards Wycherley Lake. Panel B is modified from Fisher and Souch

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(1998).

Figure 7. Hillshaded DEM of the Manitoba Escarpment on the east side of the hills and mountains. Until ice lobes retreat northwards west of the escarpment, drainage from the ice sheet and uplands was to the west and south into the Assiniboine River. On the eastern side of the escarpment, ice retreat eastwards and northwards causing new spillway channels to drain

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into the reentrants along the east side of escarpment. North of Duck Mountain only the

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Campbell Beach has been mapped.

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Figure 8. Oblique aerial photographs of two examples of channels associated with the Kaiashk spillway system, location shown in Figure 5. In A) the height of the cataract above the Devil’s Lake plunge pool is about 75 m. In B) Ottertooth Lake is 3 km long in this image and narrows to

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0.25 km wide.

Figure 9. Hillshaded DEM of the interior plains west of Lake Agassiz illustrating large spillway

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systems and controlling ice margins. Most ice margins from Bluemle (1988).

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of ro -p re lP ur na Jo Figure 10. Example of boulders associated with spillways. A) Boulders in the southern outlet spillway. The boulder closest to camera is 2.6 m high. B) Approximately 1 m diameter boulder in

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a Kelvin outlet, west of Lake Nipigon. C) Transverse giant current ripple composed of boulders up to 1.8 m diameter adjacent to Mundell Lake. Author at 1.75 m height is circled for scale. Flow

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direction was from right to left.

Figure 11. Examples of flood deposits from the Clearwater lower-Athabasca spillway. A) 5 m of

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imbricate and cross-bedded gravel overlying bedrock. Person is 1.6 m tall. B) Plane-bedded sand overlying glacial lacustrine rhythmites. Scale in cm and inches. C) Cross-bedded gravel, person is 1.95 m tall. D) 1 to 2.4 m diameter boulders, person is 1.6 m tall. E) Sub-rounded boulders, person is 1.6 m tall. F) Open-work gravel in center of image. G) Very-coarse sand to fine-gravel matrix supported pebbles and cobbles. H) Planar tabular cross-beds consisting of pebbly and open-work cobbles-to-boulder beds. Spade is 25 cm wide. I) Cobbles supported within a matrix of granules normally graded to coarse sand. Note black laminae of tar sand at top of image. J) Boulder-sized clast of diamicton within a boulder-gravel deposit.

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References

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