Proglacial lakes and the retreat pattern of the southwest Laurentide Ice Sheet across Alberta, Canada

Quaternary Science Reviews 225 (2019) 106034

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Proglacial lakes and the retreat pattern of the southwest Laurentide Ice Sheet across Alberta, Canada Daniel J. Utting*, Nigel Atkinson Alberta Geological Survey, Alberta Energy Regulator, Edmonton, Alberta, Canada

a r t i c l e i n f o

a b s t r a c t

Article history: Received 20 June 2019 Received in revised form 17 October 2019 Accepted 21 October 2019 Available online 1 November 2019

The extent of proglacial lakes following the initial separation of the southwest Laurentide Ice Sheet from the Cordilleran Ice Sheet and its eastward retreat from the Canadian Rocky Mountains has been reconstructed across regions of Alberta at a range of scales. However, to date, no studies have integrated all available geological information to produce a province-wide deglacial reconstruction that considers the evolution of proglacial lakes as components of the ice-marginal system. In this paper, we utilize a geologically constrained shoreline projection method with a high resolution digital elevation model to reconstruct the evolution of the ice-marginal system along the southwest LIS during the last deglaciation. This method provides new details on the configuration, volume, drainage history and routing of ~240 proglacial lakes as they migrated across Alberta and establishes a succession of paleogeographic reconstructions that can place geological evidence of regional and local ice-flow reorganizations into a spatiotemporal context. These reconstructions demonstrate that although the evolution of proglacial lakes was largely driven by the topography of the emerging landscape and the configuration of the ice margin, positive feedbacks in the ice-marginal system, particularly where margins transitioned from terrestrial to subaqueous settings played a major role in deglacial ice dynamics. Narrow, ribbon-shaped lakes that paralleled the ice margin induced relatively minor changes in style and rate of deglaciation, whereas the evolution of progressively larger, longer-lived lakes extending obliquely to the ice margin promoted surging and subsequent rapid retreat. Crown Copyright © 2019 Published by Elsevier Ltd. All rights reserved.

1. Introduction Proglacial lake evolution during the last deglaciation played a significant role in ice sheet dynamics (e.g. Jansson, 2003; Stokes and Clark, 2004; Lovell et al., 2012; Carrivick and Tweed, 2013), climate (Teller, 2004), hydrology and marine circulation (Soulet et al., 2013), sediment flux and landscape evolution (Larsen et al., 2011; Evans et al., 2012; Winsemann et al., 2016). A major improvement in understanding the pattern and behaviour of the numerous proglacial lakes along the southern Canadian Interior Plains has been to consider the evolution of the whole proglacial system at the sub-ice sheet scale (Teller, 2004). Part of this improvement was based on the recognition that these proglacial lakes were integral to the paleoglaciology of the southern Laurentide Ice Sheet (LIS), since deglaciation proceeded down a reverse topographic gradient, thereby impounding spatially and temporally

* Corresponding author. E-mail address: [email protected] (D.J. Utting). https://doi.org/10.1016/j.quascirev.2019.106034 0277-3791/Crown Copyright © 2019 Published by Elsevier Ltd. All rights reserved.

transgressive proglacial lakes within topographic basins abutting the retreating ice margin (Klassen, 1989; Teller and Kehew, 1994, Fig. 1). In these settings, ice margins would retreat faster due to calving and thermo-mechanical erosion, resulting in increased flow rates and ice divide drawdown due to increased sliding associated with higher basal water pressures (Teller, 2004; Carrivick and Quincey, 2014). The role of proglacial lakes has been emphasized in reconstructions of the deglacial history of the southwest LIS, following its separation from the Cordilleran Ice Sheet (CIS) east of the Albertan portion of the Canadian Rocky Mountains and Northern Foothills. These reconstructions have shown that deglaciation was characterized by complex reorganizations in flow  Cofaigh et al., 2010; pattern and dynamics (Ross et al., 2009; O Margold et al., 2015, 2018; Atkinson et al., 2016, 2018; Utting et al., 2016), in part attributed to the potential of proglacial lakes to trigger rapid ice flow (Evans, 2000; Stokes and Clark, 2004; Evans et al., 2012; Hickin et al., 2015; Atkinson et al., 2016). However, in contrast to reconstructions of marine-terminating ice streams of

2

D.J. Utting, N. Atkinson / Quaternary Science Reviews 225 (2019) 106034

Fig. 1. Study area, including an outline of Glacial Lake McConnell and Agassiz (c.f. Dyke, 2004) and radiocarbon ages (Dyke, 2004). General flow directions of major drainage basins indicated by dashed arrows. Note the general southwest to northeast descent of the topographic gradient. CLAS: Clearwater-lower Athabasca spillway.

D.J. Utting, N. Atkinson / Quaternary Science Reviews 225 (2019) 106034

Fig. 2. Resolution of LiDAR DEM coverage used in this study and location of previous proglacial lake reconstruction projects. Areas outside of the 2 m LiDAR coverage were examined using 15 m LiDAR data; volume for Glacial Lake McConnell calculated using GMTED2010 data. The study area extents of Christiansen (1979) and Lemmen et al. (1994) continue to the east and north, respectively.

the northwest LIS, which can be placed in the paleogeographic context of radiocarbon dated marine limits (Dyke et al., 2003; Stokes et al., 2009; Maclean et al., 2015; Lakeman et al., 2018), linking the dynamics and deglacial configuration of the southwest LIS is hampered by limited chronological control. Therefore, despite reconstructions of the southwest LIS improving our understanding of the regional paleoglaciology, the spatial and temporal relationship between flow reorganizations and the evolving paleogeography of the ice margin remains unclear, and to date, few studies have addressed their temporal evolution relative to the rate and pattern of ice retreat (cf. Dyke and Prest, 1987). Considering there are few radiocarbon ages that constrain the marginal chronology of the southwest LIS (Fig. 1), reconstructing the evolution of proglacial lakes at the sub-ice sheet scale can provide the paleogeographic context to constrain the relative timing of regional deglaciation and better integrate evidence of ice streaming with patterns of retreat. Proglacial lake reconstructions along the southwest LIS are available at the sub-ice sheet scale in some regions of Alberta (e.g. St-Onge, 1972; Jackson, 1980; Mathews, 1980; Moran, 1986; Paterson, 1996; Fisher et al., 2009). However, information relating to the evolution of the proglacial system across the entire province is lacking, other than at the icesheet scale (Dyke and Prest, 1987; Dyke, 2004). Therefore, in this paper, we present a geologically constrained outlet projection method to reconstruct the extent, configuration, elevation and relative chronology of proglacial lakes along the retreating margin

3

Fig. 3. A and B.Conceptual model of a reverse-gradient retreating ice sheet impounding multi-stage proglacial lakes. A: In Phase 1, a proglacial lake has formed in a pre-existing basin or river-valley and drains through outlet 1. The outlet channel may have followed a topographic low, or initially formed by water flowing along the ice front, thereby delineating a former marginal position. Headward erosion would deepen the channel, lowering the lake level until further ice retreat enabled the lake to migrate into a lower basin (B: Phase 2) which subsequently drained through a new outlet (Outlet 2). Landforms related to each phase of lake could include an outlet channel, delta, incised delta, shorelines, moraines and streamlined bedforms, as well as glaciolacustrine sediments.

of the southwest LIS, which includes the Alberta portion of the desuture zone between the LIS and CIS. These proglacial lake reconstructions incorporate previous research, as well as new mapping using LiDAR bare-earth digital elevations models (DEMs), and extend from the Late Wisconsinan limit of the southwest LIS to Glacial Lake McConnell, a large proglacial lake that covered northeast Alberta and extended northwards across parts of Northwest Territories and Yukon (Fig. 1; Lemmen et al., 1994; Dyke et al., 2003). In addition to establishing the sequence of deglaciation and correlating evidence of ice streaming with paleogeography of the retreating ice margin, an improved sub-ice sheet scale understanding of the evolution of the proglacial system will help address some outstanding issues. These include insight into the linkage between rapid retreat of western LIS and lacustrine calving (Margold et al., 2018), as well as identifying the sources, routing and mechanisms of meltwater discharge from the southern LaurentideCordilleran desuture zone (c.f. Wickert et al., 2013). Moreover, to date, no systematic analyses have been conducted that utilize

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D.J. Utting, N. Atkinson / Quaternary Science Reviews 225 (2019) 106034

Fig. 4. LiDAR bare-earth image examples of outlet channels used to reconstruct proglacial lakes in this study. (a) Pass Creek, west-central Alberta. The base of channel (800 m asl) was used for reconstructing the lowest level of Glacial Lake Fox Creek, while the higher terraces flanking the channel (857 m asl) correspond with Glacial Lake Upper Pass Creek. (b) Red Deer River area. (c) Southern Alberta. Arcuate ice marginal channels drained proglacial lakes from the wes. Arrows indicate drainage direction.

D.J. Utting, N. Atkinson / Quaternary Science Reviews 225 (2019) 106034

5

Table 1 Characteristics of proglacial lakes in Alberta. Relative Glacial Lake Name Order

Height Area (m a.s.l.) (km2)

Overlap with mapped glaciolacustrine (km2)

Volume (km3)

Present day drainage basin

Reference

100 200 300 400 500

Piney Ridge Indian Farm Whiskey Gap Twin Cardston

1370 1335 1290 1225 1195

200.52 96.13 289.84 194.66 1360.84

200.3 96.1 289.5 155.5 1359.6

9.08 5.15 15.32 6.11 58.60

Gulf Gulf Gulf Gulf Gulf

600 700 700 800 900 900 1000 1100 1200 1300 1400 1500 1600 1700 1900

Mackie Brocket Layton South Mokowan Mokowan Chokio Magrath Makowan Lower McLeod Chain Lakes Emerson Crocodile Nanton Mosquito Lethbridge

1125 1105 1105 1070 1060 1060 1040 1010 1000 1340 1355 1190 1190 1190 940

92.35 218.61 314.77 242.55 299.37 82.27 170.92 213.74 1670.67 71.73 75.72 64.06 8.54 33.18 1194.10

92.2 218.5 314.4 242.3 299.1 82.2 170.7 213.5 1668.8 71.7 75.7 64.0 8.5 33.2 1192.4

2.23 8.34 12.32 8.27 11.71 2.26 6.20 7.80 73.50 3.07 2.38 3.61 0.34 1.07 34.55

Gulf of Mexico Gulf of Mexico Gulf of Mexico Gulf of Mexico Gulf of Mexico Gulf of Mexico Gulf of Mexico Gulf of Mexico Gulf of Mexico Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Gulf of Mexico

2000 2100 2100 2200 2300 2300 2400 2400 2500 2600

Pakowki Barrier Meadow Moose Jumpingpound Irwin Bragg Mt Quirk Stimson Millarville

880 1545 1560 1545 1490 1310 1480 1470 1305 1300

471.60 43.54 11.64 6.86 12.87 24.62 5.79 19.62 68.59 191.32

469.7 43.6 11.6 6.9 12.9 24.6 5.8 19.6 68.6 191.3

7.87 3.21 0.16 0.21 0.24 0.81 0.11 0.63 2.94 7.96

Gulf of Mexico Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay

2700 2800 2900 3000 3000 3000 3100 3100 3100 3200 3300

Black Diamond Calgary Sarcee Stage Stormy Boggy Water Valley East Stormy Calgary Cullen Stage Bighill Springs Calgary Cullen Stage Moran 4 Calgary Glenmore Stage (Bow Valley) Calgary Glenmore Stage (Elbow Valley) Calgary Glenmore Stage 2 (Elbow Valley) West Nose Creek Stony Creek Water Street Dewinton Taber Chin Stage Crossfield Mossleigh Highriver Irvine Taber Forty-Mile Stage Gleichen Cartier Cremona Dogpound Dogpound (2) Madden Elkton Bergen Sundre Westcott Snakes Head Westwardhoo Beiseker

1242 1210 1360 1310 1310 1350 1185 1260 1185 1175 1150

234.94 546.97 31.15 9.46 45.16 66.75 169.67 104.24 102.61 79.06 40.91

234.9 546.9 31.2 9.5 45.2 66.8 169.0 104.3 102.0 79.1 41.0

9.75 17.28 1.24 0.07 1.44 2.74 5.01 3.49 4.36 2.20 1.84

Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson

This Study This Study This Study This Study Bretz (1943); Horberg (1952); Jackson et al. (2008); Alley and Harris (1975) GL Caldwell?) This Study This Study This Study This Study This Study This Study Bretz (1943); Horberg (1952) This Study Bretz (1943); Horberg (1952); Paterson (1996) This Study This Study This Study This Study This Study Bretz (1943); Horberg (1952); Vreeken (1989); Paterson (1996) Westgate (1968); Paterson (1996); Kulig (1996) Walker (1971); Evans et al. (1998) This Study This Study This Study This Study This Study This Study This Study Moran (1986, Stage 1), (Chain Lakes Clays and Silts of Jackson, 1980?) Moran (1986, Stage 2) Moran (1986, Stage 3) This Study This Study This Study This Study Moran (1986; Stage 4) This Study Moran (1986; Stage 4) Moran (1986) Moran (1986; Stage 5)

1150

40.88

41.0

1.22

Hudson Bay

Moran (1986, Stage 5)

1135

29.72

29.7

0.73

Hudson Bay

Moran (1986)

1210 1230 1220 1125 875 1082 1000 1035 770 825 890 1265 1170 1160 1160 1150 1150 1205 1170 1080 1120 1060 895

118.89 24.47 22.50 167.27 276.78 104.21 517.63 273.80 568.68 1300.48 673.16 42.45 10.13 28.01 8.25 39.98 19.68 50.73 153.63 80.96 99.55 52.03 1373.45

118.7 24.5 22.5 167.2 275.9 104.2 517.2 273.7 566.1 1297.0 672.2 42.5 10.1 28.0 8.0 40.0 19.7 50.7 153.7 81.0 99.6 52.0 1372.0

3.41 0.68 0.33 4.31 4.81 1.83 19.86 5.87 18.78 31.64 18.98 1.23 0.21 0.58 0.15 0.92 0.42 1.27 4.59 1.18 2.37 0.84 59.33

Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson

This Study This Study This Study This Study Bretz (1943); Horberg (1952), Paterson (1996) Moran (1986) This Study This Study This Study Bretz (1943); Horberg (1952), Paterson (1996) Paterson (1996) This Study This Study This Study This Study This Study This Study This Study This Study This Study This Study This Study Paterson (1996)

3300 3400 3500 3500 3500 3600 3700 3700 3750 3800 3900 3900 4000 4100 4200 4200 4200 4200 4400 4400 4500 4600 4700 4800 4900

of of of of of

Mexico Mexico Mexico Mexico Mexico

Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay

Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay

(continued on next page)

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Table 1 (continued ) Relative Glacial Lake Name Order

Height Area (m a.s.l.) (km2)

Overlap with mapped glaciolacustrine (km2)

Volume (km3)

Present day drainage basin

Reference

5000 5100 5100 5100 5200 5400 5400 5500 5500 5600 5900 6000 6100

770 755 815 775 760 755 770 750 760 715 1155 1105 1375

487.05 1531.02 770.21 191.11 1809.31 2631.38 256.83 1633.19 404.21 11096.25 30.46 57.25 31.85

486.0 1526.6 769.1 190.7 1801.9 2620.3 255.9 1628.8 403.4 1635.1 8.3 38.0 31.9

6.66 14.68 29.06 1.58 69.96 84.90 5.38 54.06 8.43 510.91 0.47 1.04 3.68

Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson

Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay

This Study Paterson (1996) Stalker (1973); Paterson (1996) This Study This Study Westgate (1968); Paterson (1996) This Study Bretz (1943); Horberg (1952); Paterson (1996) This Study Paterson (1996) This Study Boydell (1978) Utting et al. (2016)

1060 1325 1250 1125 1100 1020 970 1405 1385 1395 1365 1260 1300 1215 1280 1200 1265 1260 1250 1170 1255 1245 1155 1230 1230 1225 1205 1200 1120 1200 1085 1180 1100 1090 1170 1075 1140 1120 1172 1110 1126 1155 1100 1075 1070 1045 1080 1075 1075 1015 980 1047 1035 1035 965 1015 980 890 960 1040

105.21 119.02 158.98 33.93 31.63 71.59 334.26 16.90 38.09 6.03 7.42 21.52 29.33 34.25 42.12 93.76 15.70 39.37 30.06 38.26 17.87 34.57 68.72 19.02 31.94 22.68 11.53 17.50 52.15 13.97 57.23 3.09 45.98 45.41 6.89 77.56 86.59 8.17 62.03 90.70 45.79 55.49 11.93 33.19 303.07 391.82 14.41 17.59 77.98 714.36 427.92 137.50 154.47 13.25 2075.22 400.23 1893.07 397.97 1534.09 584.36

105.2 13.7 17.4 12.3 7.0 28.8 157.5 0.0 0.0 0.6 0.0 0.0 5.4 1.4 19.4 7.6 0.0 39.4 0.0 2.3 0.0 22.1 29.6 0.0 19.6 0.0 7.7 8.9 14.7 14.0 57.3 3.1 46.0 45.4 6.9 77.6 86.7 8.2 62.1 90.8 45.8 29.6 11.9 33.2 303.2 392.0 14.4 17.6 78.0 714.6 428.1 137.6 154.6 13.3 2074.9 400.5 1893.8 397.8 1535.2 584.8

2.14 9.85 10.67 1.01 0.80 0.97 2.61 0.89 1.25 0.09 0.05 0.52 0.54 1.23 0.85 2.17 0.67 0.56 1.03 0.63 0.63 0.63 1.60 0.39 0.49 0.94 0.11 0.21 0.91 0.12 1.40 0.04 0.77 0.96 0.07 1.60 2.15 0.08 2.55 2.22 1.48 1.14 0.16 0.76 10.54 12.90 0.21 0.49 1.47 28.24 11.20 3.28 3.28 0.16 27.05 10.44 53.59 10.45 76.93 25.93

Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson Hudson

Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay Bay

Boydell (1978) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) This Study Boydell (1978) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) This Study Utting et al. (2016) Utting et al. (2016) Utting et al. (2016) This Study Utting et al. (2016) Utting et al. (2016) Stalker (1960, his Fig. 9) St-Onge (1972) St-Onge (1972)

6200 6300 6400 6500 6600 6700 6800 6900 7000 7000 7100 7200 7300 7400 7500 7600 7700 7800 7900 8000 8100 8200 8300 8400 8500 8600 8700 8700 8800 8900 9000 9100 9200 9200 9300 9400 9500 9600 9700 9800 9900 9900 1010 1010 1010 1020 1030 1040 1050 1060 1060 1070 1080 1090 1100 1110 1120 1130 1140 1150

Scandia Tilly Drumheller San Fran Rutherford Medicine Hat Jenner Bassano Duchess Empress Upper Crammond Crammond I Upper North Saskatchewan Crammond II North Saskatchewan North Saskatchewan North Saskatchewan North Saskatchewan Butte Caroline Upper Brazeau Upper Brazeau Upper Pembina Upper Pembina Brazeau Pembina Brazeau Pembina Brazeau Hanlan Pembina Upper Dismal Brazeau Hanlan-Pembina Pembina Brazeau Hanlan-Pembina Pembina Hanlan North Pembina Pembina Brazeau North Pembina Brazeau North Pembina Brazeau Brazeau North Pembina Brazeau Hanlan North Pembina North Brazeau Hanlan North Brazeau Brazeau North Pembina North Brazeau Brazeau Brazeau North Pembina Dismal Hanlan O’Chiese Lasthill Hanlan Hanlan Dismal Innisfail Edson (1015 m phase) Brazeau Forks Red Deer Edson Miette

D.J. Utting, N. Atkinson / Quaternary Science Reviews 225 (2019) 106034

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Table 1 (continued ) Relative Glacial Lake Name Order

Height Area (m a.s.l.) (km2)

Overlap with mapped glaciolacustrine (km2)

Volume (km3)

Present day drainage basin

Reference

1160 1170 1100 1260 1090 1090 1120

Buck Smokey (uncertain) Gough Malmo Sullivan Monitor Bulwark

930 900 845 825 818 772 750

1611.93 255.83 427.00 161.33 561.76 626.24 280.98

1612.6 256.0 426.3 161.2 560.6 624.0 280.3

47.62 9.90 6.40 3.03 7.35 10.05 4.47

Hudson Hudson Hudson Hudson Hudson Hudson Hudson

Bay Bay Bay Bay Bay Bay Bay

Utting et al. (2016) This Study Stalker (1960) Stalker (1960; Fig 13); Holter (1973; Fig. 14) This Study This Study This Study

1270 1270 1110 1260 1270 1280 1290

Samson Buffalo Dragon Poplar Sunnybrook Telfordville Upper Pass Creek

795 790 690 850 825 795 857

283.60 255.73 495.63 2504.91 2875.36 2396.91 29915.54

283.3 255.5 493.5 944.5 1164.6 1105.6 0.0

4.56 2.90 8.23 24.57 77.40 53.10 3826.24

Hudson Hudson Hudson Hudson Hudson Hudson Hudson

Bay Bay Bay Bay Bay Bay Bay

1300 1300 1310 1310 1310 1320 1320 1330 1340 1340 1350 1350 1360 1360

Square (2) Square New Norway Buford Gainford Majeau Telford Edmonton Gwyn Stage Vermillion Minburn Whitford Westlock Birch Mathews Phase 3 (modified) Fox Creek

1010 1010 755 755 755 730 735 710 680 675 670 700 670 800

122.02 97.86 569.94 938.76 1648.23 2132.10 447.49 2653.80 390.22 422.29 741.97 4024.79 34.12 20623.95

14.9 14.9 377.0 377.0 862.5 733.4 342.6 1580.7 11.7 164.1 378.6 1007.8 0.0 0.0

5.03 6.10 5.05 25.06 44.00 81.03 9.43 62.67 4.15 8.52 18.60 120.28 0.53 2121.60

Mackenzie Mackenzie Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Hudson Bay Mackenzie

This Study Stalker (1960, his Fig. 15) This Study St-Onge (1972; his phase 2) This Study This Study St-Onge (1972; Iosegun I); Mathews (1980; Bessborough), Hickin et al. (2015; Phase 3) This Study This Study This Study This Study This Study This Study Bayrock and Hughes (1962) Bayrock and Hughes (1962) This Study This Study This Study This study This Study Mathews (1980, Phase3)

800

11068.19

7219.3

545.69

Hudson Bay

Clayhurst (Mathews, mod by Hickin) Mt Lake Lesser Slave Edmonton Wostok Stage Milligan Adskwatim Upper Notikewin Hotchkiss Hotchkiss Hotchkiss Peace Nampa Stage Edmonton Limestone Stage Mannville Vermilion Marwayne Calling North Milligan North Milligan Foulwater Peace St Onge Phase 7 (modified) Conklin Adskwatim Algar Whitemud Notikewin Hotchkiss Meikle Meikle Hotchkiss Meikle Hotchkiss-Meikle Hotchkiss North Milligan Lovet Botha Halveston Fontas Fontas

724

15751.69

0.0

161.02

Mackenzie

St-Onge (1972; Iosegun II); Mathews (1980, Phase 3); Hickin et al. (2015) Mathews (1980); Hickin et al. (2015)

675 670 655 820 797 870 975 940 910 635 635

3370.90 12885.67 681.08 627.83 143.99 444.80 9.72 26.11 105.77 16851.14 601.31

1446.1 4540.1 315.1 0.0 0.0 51.7 9.7 26.1 105.8 8312.2 234.1

131.91 106.57 17.16 34.94 2.53 23.06 0.18 0.81 4.34 318.34 8.74

Hudson Bay Hudson Bay Hudson Bay Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Hudson Bay Hudson Bay

Slomka and Utting (2017) St-Onge (1972; Phase 6) This Study This Study This Study This Study This Study This Study This Study Slomka and Utting (2017) This Study

630 620 620 640 810 785 780 625

144.19 424.44 108.28 406.83 18.48 23.60 17.66 13150.33

0.8 21.2 0.0 184.0 18.5 23.6 17.7 6226.3

3.02 6.33 2.38 14.52 0.36 0.54 0.32 271.97

Hudson Bay Hudson Bay Hudson Bay Hudson Bay Mackenzie Mackenzie Mackenzie Hudson Bay

This Study This Study This Study This Study This Study This Study This Study St-Onge (1972); Stage 1 of Leslie and Fenton (2001)

655 770 630 715 760 862 860 850 825 820 785 765 850 760 760 800 890 845

1295.52 235.24 23265.72 413.78 1157.09 66.50 12.07 99.24 70.63 81.81 118.67 95.15 8.35 161.49 319.19 174.86 27.75 31.01

71.1 78.9 23247.7 51.7 275.2 66.5 12.1 99.3 70.7 81.9 118.8 95.2 8.4 0.0 319.4 175.0 27.8 31.0

52.22 2.94 375.67 13.20 55.54 2.32 0.30 4.06 21.96 3.14 3.26 2.80 0.10 5.96 14.53 7.47 0.90 4.69

Hudson Bay Mackenzie Hudson Bay Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie

This Study This Study This Study Mathews (1980) Mathews (1980) This Study This Study This Study This Study This Study

1360 1370 1380 1380 1380 1390 1400 1410 1420 1430 1440 1530 1450 1460 1470 1480 1490 1530 1520 1510 1450 1540 1550 1560 1570 1580 1590 1600 1610 1620 1630 1640 1650 1500 1680 1690 1700 1710 1720

This Study This Study This This This This

Study Study Study Study (continued on next page)

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Table 1 (continued ) Relative Glacial Lake Name Order

Height Area (m a.s.l.) (km2)

Overlap with mapped glaciolacustrine (km2)

Volume (km3)

Present day drainage basin

Reference

1730 1740 1760 1770 1780 1790 1800 1810 1820 1830 1840 1850 1850 1860 1870 1870 1880 1890 1900 1870 1900 1930 1910 1950 1910 1930 1920 1940 1990 1660 1800 TOTAL

830 750 750 735 745 735 718 595 710 660 640 560 590 595 670 700 530 520 590 465 415 560 405 512 380 330 380 330 300 760 960

65.4 95.4 0.0 177.4 196.9 67.0 640.8 4501.3 923.2 109.4 355.0 688.2 8315.3 13905.9 64.5 149.4 225.1 6271.6 13505.7 0.0 2411.0 11277.1 4610.4 0.0 3704.6 2190.3 5879.0 5350.7 0.0 0.0 554.3

19.69 1.54 3.62 29.48 7.59 1.56 24.45 277.16 33.33 2.87 10.48 13.33 361.76 274.94 6.74 52.08 6.93 307.86 546.48 0.51 110.85 581.35 166.23 1897.23 79.24 72.67 240.66 248.70 n/d 0.22 10.72 15871.90

Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Mackenzie Hudson Bay

This Study This Study Mathews (1980) This Study This Study This Study This Study Slomka and Utting (2017) This Study This Study This Study This Study This Study Mathews (1980) This Study This Study This Study Mathews (1980); Stage 2 of Leslie and Fenton (2001) This Study This Study This study Fisher et al. (2009) Mathews (1980) Fisher et al. (2009) Mathews (1980; Phase 6) This Study Cameron (1922) This study polygon from Dyke (2004) Mathews (1980) This Study

North Milligan North Milligan Naylor Trading Post Meikle Hotchkiss Meikle Peace Deadwood Stage Hotchkiss Hotchkiss Meikle La Biche Wabasca Cadotte Stage Peace Indian Creek Stage Waniandy Vader Haro Peace Rainbow Stage Wabasca Senex Stage Kimea Peace Keg River Stage McMurray Peace Chinchaga Stage Churchill Hay Zama Stage Peace La Crete Stage Hay Slavey Stage Peace John D’Or Stage McConnell Meikle Carmangay

65.39 95.33 112.11 786.61 196.76 66.94 640.32 7344.21 922.53 109.28 354.73 1118.19 15742.46 13895.43 202.35 1194.60 291.86 8659.21 23573.38 707.05 2762.24 24287.30 5708.22 33540.72 6604.23 3557.82 9376.86 7528.06 220984.75 107.40 554.92 620601.79

detailed geological evidence to discriminate whether glaciolacustrine sediments which extend across ~20% of the Albertan portion of the southwest LIS originated proglacially or if some deposits relate to subglacial lakes (c.f. Livingstone et al., 2013, 2016). Such analyses may help improve understanding of the coupling between the storage and release of subglacial meltwater and the triggering of accelerated ice flow. 2. Study area The topographic gradient of Alberta descends northeastwards from 2500 m above sea level (asl) along the Front Ranges of the Rocky Mountains, to 250 m asl on the low relief plains at the edge of the Precambrian Shield (Klassen, 1989, Fig. 1). Physiographically, the Canadian Rocky Mountains and Foothills of southwest Alberta are fringed by the Alberta Plains, which comprise an undulating, gently sloping surface that descends northeastwards from 1100 to 450 m asl. Locally this surface is interspersed with two northwestward-orientated tracts of dissected uplands that parallel the mountain front and separate the region into the Eastern and Western Alberta Plains. The physiography of northern Alberta is characterized by broad, low-relief lowlands that descend from 750 m asl in the west to 250 m asl in the northeast. These lowlands are separated by regional uplands that rise to summits ranging from 650 to 930 m asl. Collectively, these uplands represent erosional remnants (monadnocks) within a broad, regional planation surface. Higher order fluvial systems drain northwards in the northern Alberta to the Arctic Ocean via the Mackenzie River, eastwards towards Hudson Bay in central Alberta and southwards to the Gulf of Mexico in southern Alberta (Fig. 1). During Late Wisconsinan glaciation of Alberta, the LIS and CIS coalesced along the Foothills (Dyke and Prest, 1987; Dyke, 2004; Hickin et al., 2015; Atkinson et al., 2016; Utting et al., 2016), to the

east of the mountains (Fig. 1). The maximum southern extent of LIS extended into the United States. During regional deglaciation, the southwest LIS retreated down the drainage gradient, resulting in the formation of a succession of proglacial lakes within topographic basins impounded by the margin (Dyke, 2004; Teller, 2004). This pattern of deglaciation accounts for the widespread distribution of glaciolacustrine sediments, which cover approximately one fifth of the province (Fenton et al., 2013). As the southwest LIS continued to retreat, the extent and depth of these reverse-gradient, ice-dammed lakes were controlled by the position of the ice margin relative to topography of abutting basins. This control, which evolved spatially and temporally based on the paleogeography of the proglacial system, determined lake configurations as well as the elevation and routing of their outlet channels along the ice margin or across drainage divides (cf. Teller and Kehew, 1994; Utting et al., 2016; Slomka and Utting, 2017).

3. Previous work Tyrrell (1895) first noted geomorphic and sedimentological remnants of proglacial lakes in northern Alberta, proposing that several high lake levels likely related to deglaciation, in particular “hyper-Athabasca Lake” and “hyper-Churchill Lake”. Cameron (1922) expanded on this work, reconstructing the first proglacial lake margins in the region, and identified lakes that had existed in the Peace and Hay River areas, as well as the Clearwater-LowerAthabasca spillway (Fig. 1). Cameron (1922) proposed that this spillway related to an easterly drainage system, although subsequent work has recognized that it forms part of the northwest outlet of Glacial Lake Agassiz, which drained westward along the Clearwater and Athabasca rivers, en route to Glacial Lake McConnell (Smith and Fisher, 1993, Fig. 1). Proglacial lake margins were drawn in west-central Alberta by St-Onge (1972), which were

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refined by Mathews (1980) and included an expanded reconstruction across areas to the northwest (Fig. 2). These proglacial lakes extents were further improved by Hickin et al. (2015) who presented a glacioisostatically adjusted reconstructions of Glacial Lake Peace across parts of northwest Alberta. The first reconstruction of proglacial lakes in southern Alberta was completed by Bretz (1943), who mapped a number of lake outlines and outlets. This reconstruction did not include the Calgary area, where multiple lake phases were subsequently mapped by Moran (1986). Updates for the Red Deer and Drumheller areas were presented by Stalker (1960, 1973), and the Bassano area by Paterson (1996). The headwaters of the Oldman River above Glacial Lake Cardston were examined by Alley and Harris (1975) and Jackson et al. (2008). Utting et al. (2016) reconstructed glacial lakes along the suture zone of between the LIS and CIS in west-central Alberta (Fig. 2), which included updates to areas previously examined by Roed (1975) and Boydell (1978). Christiansen (1979) reconstructed deglaciation and proglacial lakes in Saskatchewan, portraying their extent within the parts of eastern Alberta. The only other reconstruction of proglacial lakes across Alberta is at the ice-sheet scale (Dyke et al., 2003) and focussed on specific time slices rather than the extent and configuration of individual lakes. Other researchers have discussed proglacial lakes in smaller areas (e.g. Horberg, 1952; Bayrock and Hughes, 1962; Westgate, 1968; Walker, 1971; Roed, 1975; Vreeken, 1989; Kulig, 1996; Evans et al., 1998) or described the distribution of glaciolacustrine sediments rather than the extent and relative age of the associated glacial lakes (e.g. Fenton et al., 2013 and references therein).

4. Methods This paper provides generalized reconstructions of proglacial lake extents based on the elevation of their outlet channels, with the exception of Glacial Lake McConnell, which extended outside of Alberta and is therefore adapted from previous studies (Lemmen et al., 1994). These reconstructions utilize a conceptual model (Fig. 3), based on the assumption that during deglaciation, the ice sheet impounded basins or river valleys that would accumulate subaerial and meltwater drainage and form proglacial lakes, which would drain through outlets. Once incised, headward erosion of the outlet channel would progressively lower lake level, until the downslope retreat of the ice margin exposed lower elevation basins, which subsequently emptied along a lower outlet (Outlet 2; Fig. 3B). Landforms related to each lake phase could include an outlet channel, delta, incised delta, shorelines and moraines, as well as glaciolacustrine sediments. However, these features did not necessarily form for every lake phase, due to limited sediment supply, perennial lake ice, restricted fetch and the transient nature of these lakes. As well some of the features may have been eroded by post-glacial processes. Identification of outlet channels is the first step for reconstructing glacial lakes. Meltwater either flowed through an outlet channel at a topographic low, crossing a drainage divide (Fig. 4A and B), or along the ice front (Fig. 4C). In places, meltwater flowed through zones of stagnant ice along the margin. Outlet channels were identified from previously mapped undifferentiated meltwater channels (Atkinson et al., 2014a) based on topographic position at the lowest point of a drainage divide, morphology and their downward long-profile. Contemporary drainage along outlet

channels typically results in the formation of wetlands, lakes or misfit streams (Gravenor, 1956; Beaty, 1990). Although post-glacial infilling may lower the long-profiles of the channels, they still lack the undulating long-profile typically associated with subglacial channels (Greenwood et al., 2007). Previously identified subglacial channels such as the Tawattinaw tunnel valley (Sjogren, 1999) and the Sand River interlobate drainage channel (Andriashek and Fenton, 1989) have not been used to reconstruct lakes in this study. A potential for error exists in identifying the correct channel base relative to its contemporaneous lake level due to post-glacial incision or infilling. However, such errors are of limited significance considering the areal extent of lakes reconstructed in this study. Using this conceptual model, proglacial lakes were reconstructed based on the elevation of outlet channels that constrained the extent of horizontal water surfaces projected across the DEM (Utting et al., 2016). The DEMs used were LiDAR (2 m and 15 m resolution), or Global Multi-resolution Terrain Elevation Data 2010 (GMTED2010, 225 m resolution), based on the geographic coverage of these data (Fig. 2). Where basins do occur, the extent of an inferred lake is determined by identifying lower areas that must have been ice covered in order to dam a lake at the outlet channel elevation (Utting et al., 2016). In most cases this is obvious; however in a few cases interpretation must be made based on other geomorphic indications, such as moraines, identified by previous work (Atkinson et al., 2014a, b). Following the delineation of their outlines, the proglacial lakes were further characterized by determining volume and area (Table 1). These were calculated using the surface volume function in ArcGIS. This calculation is based on the modern surface topography, hence postglacial erosion or deposition can result in over- or under-estimates of lake volume, respectively. Another limitation of the volume calculation is that the surface slope of the ice margin is not known, therefore it was assumed to have a vertical face, potentially resulting in an underestimation of lake volume. As a new outlet channel formed the lake level would have lowered to the elevation of that outlet. In most cases, at this time the ice margin would also be near the outlet channel, resulting in a relatively small lake and thus the minimum volume for that outlet. As the ice margin retreated the volume of the lake would have increased, until the ice margin retreated to near the next possible lower drainage route e the maximum size of the lake. Lakes that did not start from a previous lake phase would have a minimum volume of zero; hence we have only calculated the maximum volume for each lake rather than minimum values. The effects of glacioisostatic adjustment have not been incorporated into these geologically based reconstructions, primarily because empirical information on postglacial rebound is only available in central Alberta, where differential unloading accounts for a northeasterly increase in paleoshoreline elevation of ~1.2 m/ km (Henderson, 1959; Atkinson, 2009). Consequently, this method of configuring proglacial lakes based on the projection of a horizontal plane defined by the elevation of an outlet will be most accurate near the outlet, but will become increasingly inaccurate in more distal settings, particularly in larger lakes that extended obliquely across the expected pattern of regional isobases based on their extrapolation from the northwestward part of the Lake Agassiz basin (Farrand, 1962; Rayburn and Teller, 2007). However, despite the absence of rebound data elsewhere in Alberta, many proglacial lakes extended parallel to the expected regional isobase

Fig. 5. Examples of shorelines in the study area. (a) Shorelines associated with Glacial Lake Algar, and Glacial Lake Wabasca Cadotte and Senex Stages between 630 and 580 m asl. (b) Glacial Lake McMurray shorelines fringing Muskeg Mountain up to 560 m asl; lower shorelines are associated with smaller lake formed after the incision of an outlet channel. (c) Shorelines ranging from 685 to 610 m asl, associated with Glacial Lakes Peace Nampa and Athabasca stages, and Glacial Lake Lesser Slave.

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Fig. 6. Reconstructed proglacial lakes in Alberta, categorized by volume.

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Fig. 7. Comparison between glacioisostatically corrected (Hickin et al., 2015; their phase 3) and uncorrected proglacial lake extent in west-central Alberta. Differences in proglacial lake reconstructions using these methods are negligible, apart from the eastern part of the lake, which in this study was constrained by the elevation of the Pass Creek outlet.

pattern (Dyke, 1996), minimizing the effects of differential unloading between contemporaneous lake basins. Moreover, these effects would not have influenced the relative timing of the outlets and pattern of ice retreat, so that this method remains valid for regional scale reconstructions, and has been similarly applied in other regions that lack rebound data (e.g. Jansson, 2003; Stokes and Clark, 2004). An estimate of the validity of this approach was derived by using a script in ArcGIS to compare the area of a reconstructed proglacial lake with the corresponding extent of mapped glaciolacustrine deposits, including deep-water, nearshore and littoral sediments and landforms (Fenton et al., 2013, Fig. 5). A 100% congruence between mapped glaciolacustrine deposits and predicted lake extents would, however, not be expected because of limitations of the scale and detail of previous mapping, and in some cases discontinuous glaciolacustrine deposits being grouped with other units. For reconstructed proglacial lakes that fall outside of mapped boundaries of Fenton et al. (2013), no percentage of overlap was calculated.

these configurations indicates that the glacioisostatically corrected projection underestimates lake extent proximal to the Pass Creek (Fig. 4A; 14), the likely outlet for this stage of Glacial Lake Peace, suggesting that the inferred glacioisostatic tilt, at least during this early lake stage was too small. Volumetrically, the largest proglacial lakes reconstructed in this study are glacial lakes McConnell (18,960 km3), Upper Pass Creek (3826 km3) and Peace Phase 3 (2121 km3; Fig. 6, Table 1). By area, the largest lakes include McConnell (221,000 km2), McMurray/ Meadow (33,541 km2) and Upper Pass Creek (29,915 km2). The smallest reconstructed lake was glacial lake North Pembina, with a volume and area of 0.04 km3and 3.1 km2 respectively. Maximum water depths ranged from 120 m within the glacial lake that inundated part of the North Saskatchewan River, to 7.7 m in Glacial Lake Boggy (Fig. 6). Assessing the validity of these proglacial lake reconstructions by comparing the predicted extents with the distribution of associated glaciolacustrine deposits reveals an overlap ranging from 0 to 100%, with an average of 33% (Table 1). Of the 34 predicted lakes with no corresponding deposits, most (n ¼ 22) are less than 100 km2. Of these smaller predicted lakes, the lack of corresponding sediment may simply have been missed during regional mapping, while other portions of lakes without mapped sediment may indicate sporadic sediment deposition, or glaciolacustrine sediment may be mantled by organic deposits, as occurs across large areas of northern Alberta. Elsewhere, the mapped distribution of glaciolacustrine sediment does not correspond with the predicted extent of some proglacial lakes. For example, glaciolacustrine sediments mapped in the Wainwright area (Bayrock, 1967, Fig. 1) likely formed subglacially (Jacobson, 2009), and those west of Medicine Hat (Fenton et al., 2013) have the geomorphic characteristics of streamlined till, based on our interpretation of the 15 m LiDAR DEM of the area. Although shorelines are only present in only a few areas, littoral features can be used to confirm the limits of glacial lakes Algar (Fig. 5A), McMurray (Fig. 5B), Lesser Slave (Fig. 5C), Peace Deadwood stage (Fig. 5C), Peace Nampa stage (Fig. 5C), Pass Creek, Fox Creek, and Hay. Patterns of deglaciation can be inferred from the relative sequence of reconstructed proglacial lakes and ice marginal landforms. The relative timing of proglacial lake migration across Alberta is based on the assumption that the LIS must have retreated in a northeastward, down-gradient direction to allow meltwater to accumulate along the ice margin. This method cannot distinguish where re-advances occurred, but does reveal the pattern of deglaciation following re-advances. 6. Paleogeographic reconstructions

5. Results

6.1. Southern Alberta

The surface areas, volumes, and average depths of 237 proglacial lakes were reconstructed across Alberta (Fig. 6; Table 1; Utting, 2019). The mean lake surface area is 1600 km2, the mean volume is 43 km3, and mean-depth is 31 m. In general, proglacial lake volumes increase from the mountain front (and LIS limit) to the northeast, except for those in the headwaters of the Peace River valley (Figs. 1 and 6). Although these reconstructions are not corrected for postglacial rebound, many lakes broadly parallel the expected regional isobases (Dyke, 1996), which minimises the effects of differential unloading. A comparison of the predicted configuration of Glacial Lake Peace based on the projected slope of delevelled shorelines from northeastern British Columbia to northwest Alberta (Hickin et al., 2015) and the elevation of outlets in west-central Alberta (this study) reveals negligible differences between either approach (Fig. 7). A notable difference between

The earliest proglacial lakes occurred in southern Alberta and were associated with the retreat of the southwest LIS from its LGM limit. Initial deglaciation was punctuated by readvances of local ice lobes that resulted in the formation of glacial lakes Westrup and Oldman (Alley and Harris, 1975; Jackson et al., 2008). Subsequent retreat of the LIS blocked the regional drainage, with proglacial lakes draining south into the Missouri drainage system (Fig. 8A). These lakes include Glacial Lake Cardston, which formed in an elongate (12  110 km) basin within the paleo-Oldman river valley that drained through the Lonely Valley channel (Bretz, 1943; Horberg, 1952; Jackson et al., 2008, Fig. 8B). As the LIS continued to retreat down the regional gradient, the level of Glacial Lake Cardston fell below the local height of land, resulting in multiple smaller lakes that subsequently decanted along the ice margin through the Middle Coulee channel (Fig. 8C). Ongoing deglaciation in southern

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Fig. 8. Deglacial pattern of southwest Alberta. CAIS: Central Alberta Ice Steam (Evans et al., 2008, 2014) Extent of Cordilleran ice not shown.

Alberta was characterized by the separation of the southwest LIS margin into two ice lobes associated with the Central Alberta and High Plains ice streams (CAIS and HPIS respectively; Shetsen, 1984; Evans et al., 2008). The resulting exposure of an interlobate basin enabled the accumulation of Glacial Lake McLeod (Bretz, 1943; Horberg, 1952; Paterson, 1996) which drained through a large, arcuate ice marginal meltwater channel (Fig. 4C), now demarcated by Kipp Coulee (Evans et al., 2014, Fig. 8D).

Further retreat of the CAIS along the eastern Alberta Plains was accompanied by the northward migration of proglacial lakes, culminating with the development of Glacial Lake Lethbridge (Fig. 8E) and its attendant ice marginal meltwater channel, Etzikom Coulee (Bretz, 1943; Horberg, 1952; Paterson, 1996, Fig. 4C). A succession of proglacial lakes formed along the retreating lobe of the CAIS, primarily Glacial Lake Taber, which initially drained via the Chin Coulee, and subsequently 40 Mile Coulee (Fig. 8F).

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Fig. 9. Deglacial pattern of the Bow Valley area/western Alberta. HPIS: High Plains Ice Stream (Evans et al., 2008, 2014). The location of F is shown on Figure A. Extent of Cordilleran ice not shown except in the Bow Valley. G.L.: Glacial Lake; MWC: Meltwater Channel; OC: Outlet Channel.

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However, because of the topographic gradient of the western Albertan Plains, no major proglacial lake formed in front of the HPIS until the margin retreated to the Calgary area. 6.2. Bow Valley area/western Alberta Progressive retreat and thinning of the LIS and CIS from their coalescent LGM configurations led to a series of proglacial lakes forming within the desuture zone between these ice sheets in Bow Valley area. Cordilleran ice retreating from the Barrier Lake area dammed a glacial lake in the Kananaskis valley (Fig. 9A) that drained at various elevations through a succession of ice marginal channels (Fig. 9F) and subsequently decanted via a series of small glacial lakes (Moose, Ranger and Quirk) into Glacial Lake Millarville (Fig. 9A). The glaciolacustrine sediments in the Kanananskis valley comprise laminated fines with dropstones, and are overlain by a thin till, suggesting sedimentation occurred in an ice-proximal setting, and was followed by an interval of ice marginal readvance (Evans et al., 1998). As Cordilleran and Laurentide ice margins continued to separate, meltwater from the Kananaskis valley flowed along the Sibbald meltwater channel directly into the first stage of Glacial Lake Black Diamond (Moran, 1986, Fig. 9B and F). The ongoing unblocking of the Bow River valley allowed meltwater from the Kananaskis valley to flow along the Chiniki meltwater channel (Fig. 9F) into the Sarcee stage of Glacial Lake Calgary (Fig. 9C). This lake was also supplied by meltwater flowing from the west through the Ghost meltwater channel, which is demarcated by a welldeveloped kame-delta (Smith, 1987; Fisher, 1999). During this interval, a succession of smaller proglacial lakes accumulated north of the Bow River, culminating with the establishment of the Bighill Springs meltwater channel (Fisher, 1999) which flowed into Glacial Lake Calgary Sarcee stage (Fig. 9C). By the establishment of stage 3 (Moran, 1986), Glacial Lake Calgary had split into the Bow and Elbow valleys, which were connected by the Cullen Creek outlet channel (Fig. 9D). Continued ice retreat maintained a similar lake configuration, with lakes draining through progressively lower outlets, until meltwater was captured by an elongate, ice parallel basin to form Glacial Lake Dewinton which drained to the southward along the ice margin. This ice configuration corresponds to the position of a succession of arcuate push moraines around Frank Lake, which relate to the marginal lobation of the High Plains ice stream (Stalker, 1973; Evans et al., 1999). Similarly, Moran (1986) identified readvances elsewhere in the area, which deposited till over glaciolacustrine sediment and blocked the recently established outlets, thereby allowing lakes to refill to similar configurations as depicted in Fig. 9D and E. Subsequent retreat repeated the earlier deglacial pattern (Moran, 1986) until the ice margin finally retreated to the north, opening the Bow River valley and allowing drainage to the east. 6.3. East-Central Alberta As deglaciation continued, the ice margin retreated northwards from Glacial Lake Taber (Fig. 8F), damming Glacial Lake Gleichen in a basin flanking the Bow River valley (Fig. 10A). The continued separation of HPIS and CAIS (Evans et al., 2008) formed a broad interlobate basin within which Glacial Lake Beiseker accumulated. This lake drained through the Tudor valley into Glacial Lake Chancellor (Stalker, 1973, Fig. 10B). As the western margin of the CAIS retreated further north, Glacial Lake Beiseker migrated to inundate now exposed reaches of the Red Deer River valley and form Glacial Lake Drumheller (Stalker, 1973). Glacial Lake Drumheller decanted along the Crawling valley (Stalker, 1973, Fig. 10C) into a series of smaller, relatively short-lived lakes at the lobate

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terminus of the CAIS. However, as this lobe retreated more extensively to the northeast, it exposed lower reaches of the Red Deer River valley, resulting in the adjacent lowlands being inundated by Glacial Lake Bassano (Fig. 10D) which drained towards the south (Paterson, 1996). Further retreat of the CAIS lobe opened the Red Deer River valley, initiating final drainage of Glacial Lake Bassano (Bryan et al., 1987). 6.4. West and Central Alberta As the Peace River lobe retreated from west-central Alberta (Fig. 11), Glacial Lake Fox Creek formed between the mountainfront and Swan Hills, which drained southwards through Pass Creek (Mathews, 1980; St-Onge, 1972; Hickin et al., 2015). Downcutting of this outlet lowered water levels, dividing the lake into two basins connected by a short channel (Mathews, 1980, Fig. 11A). During this stage, the northern basin may have drained to the northwest, while the southeast basin continued to drain through Pass Creek, likely supplying proglacial lakes evolving across central Alberta. These included Glacial Lake Edmonton (Bayrock and Berg, 1966, Fig. 11B and C, 12A), which resulted from the accumulation of meltwater and subaerial discharge along the Sturgeon and North Saskatchewan rivers which were blocked by a broad lobe of Laurentide ice across central Alberta. Glacial Lake Edmonton initially drained southwards through the Gwynne Channel into the Battle River (Fig. 12A), which together with pitted deltas ~30 km to the northwest (Bayrock and Hughes, 1962) constrain the extent of this lake. Maximum lake extent was established following retreat of the ice margin and exposure of a broad lowland to the northeast (Fig. 12A). It was only once ice retreated north of the Beaver Hills Upland (Fig. 12A) that the lake lowered and drained to the east (Figs. 11B and 12A). During this stage, the base of the Gwynne Channel was perched above the lake, resulting in subsequent drainage occurring along lower ice marginal channels to the east (Fig. 11B and C), until Laurentide ice unblocked the North Saskatchewan River, resulting in the final emptying of the lake. Glacial lakes Lesser Slave, Mt. Lake and Peace (Clayhurst stage) represent a series of large proglacial lakes formed during the retreat of lobate LIS margins from the Peace River Lowlands and the Lesser Slave Lake area (St-Onge, 1972; Mathews, 1980; Pawley and Atkinson, 2012; Slomka and Utting, 2017, Fig. 11B). Although topographically constrained within discrete basins, these lakes sequentially decanted southeastwards along the ice margin through a series of meltwater channels, such as along the Fish Creek Moraine (Slomka and Utting, 2017). However, as lobes of Laurentide ice continued to retreat from west-central Alberta, these lakes coalesced to form Glacial Lake Peace Nampa Stage, a ~550 km long, ribbon-shaped water body extending from the Peace River to Edmonton region (Fig. 11C). Once Laurentide ice had retreated from the Lesser Slave Lake area, low topography focussed drainage to the east, forming the Athabasca meltwater channel (Fig. 11D), now occupied by the Athabasca River. This established an effective drainage route for Glacial Lake Peace Nampa stage, which resulted in a rapid ~10 m decrease in water level (Table 1). Lake levels to the west subsequently stabilized, until the margin of the Peace River lobe retreated far enough north to open up a succession of lower outlets (Mathews, 1980; Atkinson and Paulen, 2010; Slomka and Utting, 2017). Contemporaneously with this stage of Glacial Lake Peace, the ice margin north and east of Lesser Slave Lake separated into three discrete lobes which blocked a number of topographic lows, thereby impounding Glacial Lake Algar (Fig. 11E). The upper stage of Glacial Lake Algar is constrained by its southern outlets along Amisk and Long Lakes (Fig. 11E). Flights of shorelines formed during successively lower lake stages (Fig. 5A)

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suggest the Athabasca River valley remained blocked by Laurentide ice for a sufficient duration to enable Glacial Lake Algar to extend ~250 km north across the Wabasca Lowlands and drain by progressive downcutting of southern outlets. The eastward retreat of a lobe of Laurentide ice facilitated the drainage of Glacial Lake Algar via a channel now occupied by the Beaver River, en route to the Kehiwin channel (Fig. 11E and F; 12) and ultimately the North Saskatchewan River to the south (Norris et al., 2019). The Kehiwin drainage resulted in the level of Glacial Lake Algar falling below the local height of land, separating the lake into Glacial Lake Wabasca Cadotte Stage in the north and Glacial Lake Beaver in the south (Fig. 11F).

6.5. Northwest Alberta Deglaciation of northwest Alberta was characterized by the retreat of the LIS margin across the broad lowlands flanking the Peace and Hay rivers (Fig. 1). Mathews (1980) showed the Indian Creek stage of Glacial Lake Peace drained southeast through the Iroquois Lakes channels (Fig. 11E). However, as the Peace River lobe continued to retreat further north, drainage of the Rainbow stage of Glacial Lake Peace shifted to the northwest, through the Rainbow Lake outlet channel (Fig. 13A, E). This shift caused Glacial Lake Peace to drop ~75 m (Table 1). The Rainbow Lake outlet channel was partially infilled by glaciofluvial sediments, either due to allogenic controls such as deltaic progradation into Glacial Lake Hay (Mathews, 1980; Smith et al., 2005) or autogenic variations such as decreased flow velocities due to widening of the channel. Subsequent retreat of the Peace River lobe resulted in its bifurcation around the Caribou Mountains, forming a southern lobe that continued to impound Glacial Lake Peace, and a northern lobe that impounded the Zama stage of Glacial Lake Hay. This lake submerged the Fort Nelson lowland, reaching the local height of land ~150 km to the west, where it drained along the Fontas outlet channel (Mathews, 1980, Fig. 13B). The decanting of Keg River stage of Glacial Lake Peace into Zama stage of Glacial Lake Hay is recorded by a series of arcuate meltwater channels that parallel and in some areas are captured by the Chinchaga River (Fig. 13C, E). A delta at the mouth of one of these channels records meltwater influx into Glacial Lake Hay (Levson et al., 2004). Based on the work of Mathews (1980) and the elevation of the floor of the Fontas outlet channel, strandlines identified by Paulen et al. (2005a,b) at 390 m asl likely relate to the onset of drainage along of this channel during the Zama stage of Glacial Lake Hay, which is portrayed on Fig. 13B at 380 m asl. Lower strandlines mapped by Paulen et al. (2005a,b) are attributed to a ~340 m asl post-glacial highstand of Lake Zama, rather than an ice-dammed phase (Fig. 13D). As the margin of the Hay lobe retreated further east, Glacial Lake Hay formed an expanded Slavey stage, although the persistence of the Peace lobe continued to impound Glacial Lake Peace, resulting in ongoing drainage through the Chinchaga River (Fig. 13C). As the Hay and Peace lobes withdrew around the Caribou Mountains, Glacial Lake Hay drained along the ice margin to the north, while drainage during the La Crete stage of Glacial Lake Peace incised the Meander River outlet (Fig. 13D). This drainage was captured by the Hay River, and continued northeast, potentially extending under the margin of the Hay lobe (Campbell, 2006). As Laurentide ice retreated east of Mt Watt, Glacial Lake Peace migrated across the Vermillion Lowland and continued to drain through the Meander River outlet (GLP John D’Or Stage). This late stage configuration of Glacial Lake Peace likely persisted until Laurentide ice retreated east of the Great Slave Plain, where it eventually coalesced with Glacial Lake McConnell.

6.6. Northeast Alberta Following the establishment of the Kehiwin channel (Figs. 11E and 12), the water level of Glacial Lake Algar (Fig. 14A) dropped, dividing it into two lakes. The northern lake, Glacial Lake Wabasca Cadotte Stage, likely drained to the southeast through the Athabasca River valley (Figs. 11F and 14B). Alternatively, as the Wabasca ice lobe retreated to the north, the Cadotte Stage of Glacial Lake Wabasca may have drained to the west into Glacial Lake Peace (Slomka and Utting, 2017). As the Wabasca and McMurray ice lobes continued to retreat, Glacial Lake Wabasca would have grown in extent and volume (Fig. 14C) until a drainage route was exposed at 560 m asl, allowing flow into Glacial Lake Peace La Crete Stage (Fig. 13D). This outlet likely persisted until ice had retreated east of the Birch and Muskeg mountains, which impounded Glacial Lake McMurray (Fig. 14D). Shorelines on Muskeg Mountain at 560 m asl (Fig. 5B), suggest that Glacial Lake McMurray extended at least this far to the northeast (Fig. 14D). As the ice margin southeast of Muskeg Mountain retreated east, Glacial Lake McMurray likely coalesced with a lake to the southeast (Glacial Lake Meadow) forming Glacial Lake McMurray/Meadow (Fisher et al., 2009; Anderson, 2012, Fig. 14E). This lake likely continued to decant through the Wabasca River, until it finally drained into either Glacial Lake Agassiz to the southeast or Glacial Lake McConnell to the north. However, there are no obvious geomorphic indicators supporting either drainage route (Fisher et al., 2009). The Clearwater-lower Athabasca spillway (Figs. 1 and 13F) may have been caused by drainage of Glacial Lake McMurray/Meadow, or a subsequent drainage of Glacial Lake Agassiz (Fisher and Smith, 1994). 7. Discussion In this paper, we utilize an outlet projection method that integrates a range of new geological evidence to reconstruct the evolution of the ice-marginal system along the southwest LIS during the last deglaciation. This reconstruction provides a firstorder prediction of the area and volume of a succession of proglacial lakes that evolved during the initial separation of the LIS and CIS and reveals new details on their configuration, drainage history and routing as they migrated east of Alberta’s Rocky Mountains. However, because lakes operated as open systems throughout their history, the spatial and temporal effects of periodic drainage and refilling on total lake area and volume remain unknown. Therefore, the predictions presented in this paper represent minimum constraining values, and are intended to illustrate the relative differences between sizes of proglacial basins across Alberta during a single maximum lake extent. Collectively, these reconstructions provide insight into the contribution of proglacial lakes evolving in the southern portion of Laurentide-Cordilleran desuture zone to meltwater discharge into the Gulf of Mexico and Arctic Ocean via the Mississippi and Mackenzie drainage systems respectively. From this reconstruction, it is also possible to establish a temporal sequence of likely paleogeographic scenarios to constrain the relative chronology of ice retreat, which can be placed within the ice sheet scale retreat patterns based on the distribution of radiocarbon dates (Dyke et al., 2003). These scenarios in turn provide a spatiotemporal framework to integrate the evolution of proglacial lakes with geological evidence of large-scale complex flow reorganizations in the southwest LIS, particularly in areas where the proglacial system transitioned from terrestrial-to lake-terminating settings. Finally, by contrasting these paleogeographic scenarios with the mapped distribution of glaciolacustrine sediments across the Albertan portion of the southwest LIS (Fenton et al., 2013), some deposits are evidently inconsistent with predicted location of a

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Fig. 10. Deglacial pattern of east-central Alberta. CAIS: Central Alberta Ice Stream; HPIS: High Plains Ice Stream (Evans et al., 2008, 2014). G.L.: Glacial Lake.

proglacial lake, suggesting such areas may relate to the distribution of subglacial lakes or were incorrectly mapped. 7.1. Proglacial lake reconstructions The results of the outlet projections utilized in this paper indicate that the deglacial paleogeography of Alberta was characterized by the spatially and temporally complex evolution of ~240 proglacial lakes associated with the retreat of the southwest LIS (Fig. 6; Table 1). This reconstruction shows how the configuration, drainage history and routing of proglacial lakes across Alberta was in large part constrained by feedbacks between the topography of the emerging landscape and the configuration of the retreating ice margin. In order to provide an ice-sheet scale context for these proglacial lake reconstructions, we have used the radiocarbon-constrained deglacial paleogeography of the LIS presented by Dyke and Prest (1987) and Dyke et al. (2003), rather than the eolian dune luminescence chronology, which is only available from Alberta (Munyikwa et al., 2017). For the purposes of integrating evidence of changing ice sheet dynamics with the evolving configuration of proglacial lakes, both chronologies indicate comparable rates of ice retreat, despite the luminescence chronology indicating that the

onset of southwest LIS deglaciation proceeded ~1 ka earlier. The earliest proglacial lakes formed in southernmost Alberta at the onset of regional deglaciation between 14.5 and 14 ka BP (Dyke et al., 2003) and occupied 16 relatively small, short-lived elongate basins that supplied meltwater to the Mississippi River drainage system by a combination of down-gradient, ice-contact migration and outflow along spillways and proglacial channels extending between the Milk River Ridge and the Cypress Hills (Fig. 6). As deglaciation continued across the more open relief Alberta Plains north of the Milk River Ridge and the Cypress Hills, a succession of basins emerged obliquely from the ice margin between 14 and 11.5 ka BP (Dyke et al., 2003), and resulted in the accumulation of meltwater within 162 proglacial lakes (or associated lake stages) that drained eastward across the Alberta Plains, although due to the continued occupation of the incipient Hudson Bay drainage system by the LIS, this meltwater was instead discharged into the upper Mississippi. In the Lethbridge (Fig. 8) and Drumheller regions (Fig. 10), proglacial lakes drained along a suite of large arcuate meltwater channels eroded along the lobate margin of the southwest LIS. In settings where meltwater channels extended into topographic basins emerging at successively lower elevations along the ice margin, drainage would have been punctuated by the filling of these basins, which would sequentially decant eastwards across

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Fig. 11. Deglacial pattern of central Alberta. G.L.: Glacial Lake.

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Fig. 12. (a) Deglacial pattern of the Edmonton area. Drainage of Glacial Lake Edmonton Gwynne Stage was via the Gwynne Channel to the southeast, until ice retreated to the north of Beaver Hills Upland, allowing the lake to lower to the Whitford Stage, draining to the east. The location of this figure is shown on Fig. 11A. (b) The southeastward drainage of Glacial Lake Algar along the Beaver River, showing its southward deflection through a submarginal conduit by ice occupying uplands to the east. Submarginal drainage was captured by the Kehiwin Channel, which continued to flow south, en route to the North Saskatchewan River (not shown). The location of this figure is shown on Fig. 11F.

the Alberta Plains. Further north, deglaciation of the Central Alberta region was characterized by progressively larger and longer-lived proglacial lakes (Fig. 11). These lakes drained primarily by downgradient, ice-contact migration and outflow along channels extending away from the ice margin. During the final deglaciation of the Alberta Plains, high elevation lakes were impounded between the ice margin and the Swan Hills and Berland Uplands (Fig. 11a), which drained along outlets extending through low points in these uplands (Mathews, 1980; St-Onge, 1972; Hickin et al., 2015). Outlets such as Pass Creek decanted meltwater into a succession of basins extending southeastwards along the retreating ice margin and culminated with the submergence of a large proglacial basin in Edmonton area. However, as the ice margin vacated the gap between the Swan Hills and Pelican Mountains, a significant reconfiguration of lake extent occurred due to meltwater

draining south along the Athabasca valley and rapidly submerging basins to the south (Fig. 11b). The ice margin retreated beyond the Swan Hills and Pelican Mountains and on to the Northern Alberta Lowlands after 11.5 ka BP, and by the final deglaciation of Alberta at 9.5 ka BP (Dyke et al., 2003), 58 proglacial lakes (or associated stages) had inundated emergent basins across the region. Although minor lakes accumulated in the uplands of northwest Alberta that decanted through a series of small meltwater channels, the deglacial history of northern Alberta was dominated by the behaviour of proglacial lakes occupying the Peace River, Wabasca and McMurray lowlands. During the initial retreat of Laurentide ice lobes from Northern Alberta, these lakes were impounded within three broad basins that extended obliquely to the retreating margin. As ice lobes continued to retreat north and bifurcate around the Caribou

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Fig. 13. Deglacial pattern in northwest Alberta. Ice extent on Caribou Mountains not shown. The location of Figure E is shown on Figure A. G.L.: Glacial Lake; OC: Outlet Channel.

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Fig. 14. Deglacial pattern in northeast Alberta. Ice extent over the Birch Mountains is not shown. G.L.: Glacial Lake; CLAS: Clearwater-lower Athabasca spillway.

Mountains (Fig. 13A), Glacial Lake Peace overflowed along a series of well-developed spillways that extended into Glacial Lake Hay. The configuration of proglacial lakes during the final stages of deglaciation in the western and central regions of Northern Alberta was influenced by the combined effects of ice margins retreating

north, which opened the Hay River and enabled the drainage of Glacial Lake Hay, while in the east resulted in the convergence of Glacial Lake Wabasca with Glacial Lake Peace (Fig. 13D). During this late stage configuration, Glacial Lake Peace continued to drain through the Meander River outlet until the ice margin retreated

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east of the Great Slave Plain, where it coalesced with Glacial Lake McConnell. In northeast Alberta, retreating Laurentide ice impounded Glacial Lake McMurray, which initially migrated northeast due to down-gradient, ice-contact migration. Portions of the ice margin subsequently retreated eastwards, enabling Glacial Lake McMurray to coalesce with Glacial Lake Meadow, which had occupied adjacent lowlands in Saskatchewan. The coalescence of these lakes culminated with the formation Glacial Lake McMurray/ Meadow, which occupied a ribbon-shaped, southeast-orientated ice-contact basin spanning the Alberta-Saskatchewan border. This lake likely decanted into either Glacial Lake Agassiz to the southeast or through the Wabasca River into Glacial Lake McConnell to the north. 7.2. Implications for ice sheet dynamics In addition to establishing a spatiotemporal framework to constrain the configuration, drainage history and routing of proglacial lakes, these reconstructions now enable geologic evidence of the dynamics and patterns of ice retreat to be integrated with the deglacial paleogeography of the southwest LIS. The limited distribution of glaciolacustrine sediment-landform assemblages associated with proglacial lakes in southern Alberta (Atkinson et al., 2018) suggests that subaqueous processes exerted a minimal impact on the deglacial paleoglaciology of this portion of the southwest LIS. This likely relates to the geometry and duration of these lakes, which were insufficiently wide, deep or long-lived to promote widespread destabilization and reorganization of major flow units such as the Central Alberta and High plains ice streams. Subaqueous landsystems are more widespread on the Alberta Plains, where streamlined and grooved glaciolacustrine sediments demarcate up to 60 km long and 45 km wide fan-shaped flowlines splaying from the main flow-sets of Central Alberta and High plains ice streams (Evans et al., 2014). These sediment-landform assemblages record surging of subaqueous ice lobes, which readvanced into successively younger deglacial lakes, including Glacial Lakes McLeod, Lethbridge, Gleichen, Bassano, Beiseker and Drumheller. Flowlines based on mapped glaciolacustrine sediment-landform assemblages within the High Plains ice stream exhibit increasing complexity in west and central Alberta (Atkinson et al., 2014a, b). These include flowsets that contain up to 90 km wide, 40 km long tracts of ridge-furrow corrugations and sediment plowed grooves, which relate to drawdown and substantial readvances of the High Plains ice stream across the basins of Glacial Lakes Buck, Poplar, Sunnybrook, Edmonton, Majeau and Westlock. Contemporaneous, albeit more topographically restricted readvances also occurred along the western margin of the Central Alberta ice stream, where local ice lobes extended into Glacial Lakes Gough and New Norway. However, sediment-landform assemblages associated with extensive subaqueous ice marginal readvances are pronounced across the Northern Alberta Lowlands (Atkinson et al., 2018). These landsystems indicate that the evolution of progressively larger and longer-lived proglacial lakes exerted an increasingly significant influence on the paleoglaciology of the southern portion of the Laurentide-Cordilleran ice saddle (Hickin et al., 2015; Atkinson et al., 2016). For example, extensive (up to 150 km) readvances of the CIS and LIS have been mapped across the Peace River Lowlands of Alberta and British Columbia, and have been attributed to unimpeded expansions of rapidly flowing Cordilleran and Laurentide ice lobes which were drawn-down into Glacial Lake Peace (between Mathews and Clayhurst stages). Comparable subaqueous landsystems evolved during each of the subsequent stages of Glacial Lakes Peace, Wabasca and McMurray, where fan-shaped flowlines delineated by 40e75 km wide, and 50e100 km long tracts of closely spaced ridge-furrow corrugations and subaqueous

plough marks terminate at large moraines or arcuate belts of crevasse-fills ridges (Atkinson et al., 2018). These sedimentlandform assemblages indicate that retreat of the southwest LIS from the Northern Alberta Lowlands was punctuated by at least 7 major readvances associated with internal reorganizations in deglacial ice dynamics, ranging from local-scale, non-steady flows, to regional-scale surges. The preservation of landforms after each readvance suggests that the ice margin refloated during the establishment of subsequent lake stages, resulting in widespread subaqueous break-up (Hickin et al., 2015; Atkinson et al., 2016). These reconstructions demonstrate that although the evolution of proglacial lakes was largely driven by the topography of the emerging landscape and the configuration of the ice margin, positive feedbacks in the ice marginal system, particularly where margins transitioned from terrestrial to subaqueous settings played a major role in deglacial ice dynamics. These feedbacks are exemplified by the deglaciation of the Northern Alberta Lowlands, which were characterized by runaway effects in the ice marginal system due to the evolution of progressively larger, longer-lived lakes that extended obliquely to retreating margin. 7.3. Discriminating between proglacial and subglacial lakes Although this reconstruction delineates large, proglacial glacial lake systems, it does not necessarily account for the distribution of all glaciolacustrine sediments mapped across Alberta (cf. Fenton et al., 2013). For example, some sediments may have been deposited within dead-ice basins or in small valley floor lakes that were blocked by residual ice (Jansson, 2003). In other areas, sediments may have been deposited subglacially. This includes an approximately 3000 km2 upland surrounding Utikuma Lake in northcentral Alberta (Fig. 15). Although this upland is mantled by widespread, discontinuous glaciolacustrine sediments (Paulen et al., 2004), it is not predicted to have been inundated by a proglacial lake in our reconstruction. Instead, Livingstone et al. (2013) predicted that a subglacial lake was likely to have accumulated across these uplands, which is congruent with the glacial geomorphology of the area, which exhibits features diagnostic of subglacial meltwater drainage (Fig. 15; Livingstone et al., 2016). These findings provide new information to compare predicted distributions of pro- and subglacial lakes with sedimentological and geomorphic evidence to determine their origin and potential influence on the paleoglaciology of the southwest LIS. 8. Conclusions Based on the widespread distribution of glaciolacustrine sediments and associated landforms, the extent and drainage history of proglacial lakes during the last deglaciation have been reconstructed at a range of scales across northern, west-central and southern regions of Alberta (Tyrrell, 1895; Cameron, 1922; Bretz, 1943; Stalker, 1960, 1973; St-Onge, 1972; Alley and Harris, 1975; Roed, 1975; Boydell, 1978; Christiansen, 1979; Smith and Fisher, 1993; Mathews, 1980; Moran, 1986; Evans et al., 1998; Jackson et al., 2008; Hickin et al., 2015; Utting et al., 2016). However, to date, no studies have integrated all available geological information to produce a province-wide reconstruction that considers the configuration, drainage history and routing of proglacial lakes as important components of the ice-marginal system at the sub-ice sheet scale. Utilizing a geologically constrained outlet projection method, we have reconstructed ~240 proglacial lakes across Alberta. The evolution of these lakes was largely driven by the topography of the emerging landscape and the configuration of the retreating ice margin. As a result, proglacial lakes within the Alberta portion of

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Fig. 15. Extensive glaciolacustrine sediments mapped in the Utikuma uplands (Paulen et al., 2004) occupy areas above predicted proglacial lakes. These sediments may instead relate to subglacial lakes which were modelled to have a high likelihood of existing in this location (modified after Livingstone et al., 2013).

the LIS exhibit two distinct evolutionary patterns, which provide new insights into the pattern of ice retreat. In southern Alberta and the Alberta Plains, relatively small, short-lived elongate basins drained sequentially by a combination of down-gradient, ice-contact migration and decanting outflow along spillways and channels eroded along the lobate margin of the southwest LIS. Our predictions indicate that these proglacial lakes drained to the Gulf of Mexico via the upper Mississippi drainage system between 14.5 and 11.5 ka BP, following the retreat of ice north of the Cypress Hills and the opening of eastward drainage routes. In contrast, deglaciation of the Northern Alberta Lowlands, which included the widening desuture zone between the CIS and LIS along the mountain front, was dominated by the evolution of proglacial lakes impounded within three broad basins that extended obliquely to retreating ice margin. These drained primarily due to downgradient, ice-contact migration, although overflow spillways continued to supply proglacial lakes occupying lower basins to the north. During the final deglaciation of Alberta, these discrete proglacial lakes coalesced to form two larger meltwater bodies that extended beyond Alberta and into Saskatchewan and the Northwest Territories. Our predictions of proglacial lakes in Northern Alberta Lowlands, excluding components of Glacial Lakes McMurray/Meadow and McConnell, indicate this region supplied freshwater to the Arctic Ocean drainage system between 11.5 and 9.5 ka BP. These geologically based projections may benefit future modelling studies, because our reconstructions suggest that although rapid drainage events did occur in the southern LIS, meltwater flux across western Canada was modulated by the residence time of proglacial lakes on the landscape. Collectively, these reconstructions demonstrate that during the last deglaciation of Alberta (between 14.5 and 9.5 ka BP), proglacial lakes drained sequentially towards the south and east (Gulf of Mexico drainage) and then north (Arctic Ocean drainage) by a combination of downgradient, ice-contact migration and decanting outflow. Although

the far field effects of proglacial lakes in the Alberta may have been relatively minor, these reconstructions demonstrate the linkages between evolving geometry of emergent lake basins and the deglacial paleogeography of the ice margin, as well as feedbacks between the geometry and duration of proglacial lakes and paleoglaciology of the southwest LIS. Narrow, ribbon-shaped lakes that paralleled the ice margin induced relatively minor changes in style and rate of deglaciation, whereas the evolution of progressively larger, longer-lived lakes that extended obliquely to retreating ice margin promoted surging and subsequent rapid retreat, paired with drawdown of divides. Acknowledgements We appreciate discussions and input on this work with colleagues at the Alberta Geological Survey, including L. Andriashek, L. Atkinson, M. Fenton, G. Hartman, and J. Slomka. We would also like to thank Art Dyke and one anonymous reviewer for providing comments that improved this manuscript. References Alley, N.F., Harris, S.A., 1975. Pleistocene glacial lake sequences in the foothills, southwestern Alberta, Canada. Can. J. Earth Sci. 11, 1220e1235. Anderson, T.W., 2012. Evidence from Nipawin Bay in Frobisher Lake, Saskatchewan, for three highstand and three lowstand lake phases between 9 and 10 (10.1 and 11.5 cal) ka BP. Quat. Int. 260, 66e75. Andriashek, L.D., Fenton, M.M., 1989. Quaternary Stratigraphy and Surficial Geology of the Sand River Area 73L, vol. 57. Terrain Sciences Department and Alberta Geological Survey, Alberta Research Council Bulletin, pp. 1e154. Atkinson, N., 2009. Surficial Geology and Quaternary History of the High Prairie Area (NTS 83N/SE). Energy Resources Conservation Board, ERCB/AGS Open File Report 2009-07, pp. 1e34. Atkinson, N., Paulen, R.C., 2010. Surficial Geology and Quaternary History of the Cleardale Area. Energy Resources Conservation Board, ERCB/AGS Open File Report 2010-11, pp. 1e27. Atkinson, N., Utting, D.J., Pawley, S.M., 2014a. Glacial landforms of Alberta. Alberta Geol. Surv. Map 604, 1:1 000 000 scale. Atkinson, N., Utting, D.J., Pauley, S.M., 2014b. Landform signature of the Laurentide

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Milk River in southern Alberta: a classic underfit stream. Can. Geogr. 34, 171e174. Boydell, A.N., 1978. Multiple Glaciations in the Foothills, Rocky Mountain House Area, vol. 36. Alberta Research Council Bulletin, Alberta, pp. 1e35. Bretz, J.H., 1943. Keewatin end moraines in Alberta, Canada. Bull. Geol. Soc. Am. 54, 31e52. Bryan, R.B., Campbell, I.A., Yar, A., 1987. Postglacial geomorphic development of the dinosaur provincial park badlands, Alberta. Can. J. Earth Sci. 24, 135e146. Cameron, A.E., 1922. Post-glacial lakes in the Mackenzie River basin, north west Territories, Canada. J. Geol. 30, 337e353. Campbell, H., 2006. The Formation of the Glacial Lake Peace Meltwater Channel in Northern Alberta during the Retreat of the Laurentide Ice Sheet and the Potential for Aggregate Resources. University of New Brunswick, unpublished B.Sc. thesis. Carrivick, J., Tweed, F.S., 2013. Proglacial lakes: character, behaviour and geological importance. Quat. Sci. Rev. 78, 34e52. Carrivick, J., Quincey, D.J., 2014. Progressive increase in number and volume of icemarginal lakes on the western margin of the Greenland Ice Sheet. Glob. Planet. Chang. 116, 156e163. Christiansen, E.A., 1979. The Wisconsinan deglaciation of southern Saskatchewan and adjacent areas. Can. J. Earth Sci. 16, 913e938. Dyke, A.S., 1996. Preliminary Paleogeographic Maps of Glaciated North America. Geological Survey of Canada, Openfile, p. 3296. Dyke, A.S., 2004. An outline of North American deglaciation with emphasis on central and northern Canada. In: Ehlers, J., Gibbard, P.L. (Eds.), Quaternary Glaciations e Extent and Chronology. Part II: North America. Elsevier, Amsterdam, pp. 373e424. Dyke, A.S., Prest, V.K., 1987. Late Wisconsinan and Holocene Retreat of the Laurentide Ice Sheet; Geological Survey of Canada, “A” Series Map 1702A. Dyke, A.S., Moore, A., Robertson, L., 2003. Deglaciation of north America. Geol. Surv. Can. Open File 1574. Evans, D.J.A., 2000. Quaternary geology and geomorphology of the Dinosaur Provincial Park area and surrounding plains, Alberta, Canada: the identification of former glacial lobes, drainage diversions and meltwater flood tracks. Quat. Sci. Rev. 19, 931e958. Evans, D.J.A., Salt, K.E., Allen, C.S., 1998. Glacitectonized lake sediments, Barrier Lake, Kananaskis country, Canadian Rocky mountains. Can. J. Earth Sci. 36, 395e407. Evans, D.J.A., Lemmen, D.S., Rea, B.R., 1999. Glacial landsystems of the southwest Laurentide Ice Sheet: modern Icelandic analogues. J. Quat. Sci. 14, 673e691. Evans, D.J.A., Clark, C.D., Rea, B.R., 2008. Landform and sediment imprints of fast glacier flow in the southwest Laurentide Ice Sheet. J. Quat. Sci. 23, 249e272.  Cofaigh, C., Rea, B.R., 2012. Evans, D.J.A., Hiemstra, J.F., Boston, C.M., Leighton, I., O Till stratigraphy and sedimentology at the margins of terrestrially terminating ice streams: case study of the western Canadian prairies and high plains. Quat. Sci. 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