Catastrophic flooding from Glacial Lake Wisconsin

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Geomorphology 93 (2008) 384 – 397 www.elsevier.com/locate/geomorph

Catastrophic flooding from Glacial Lake Wisconsin Jordan A. Clayton a,⁎, James C. Knox b b

a Department of Geography, 260 UCB, University of Colorado, Boulder, CO 80309-0260, USA Department of Geography, 234 Science Hall, 550 N. Park St., University of Wisconsin, Madison, WI 53706, USA

Received 25 September 2006; received in revised form 12 March 2007; accepted 12 March 2007 Available online 23 March 2007

Abstract Glacial Lake Wisconsin was a large proglacial lake that formed along the southern margin of the Laurentide Ice Sheet during the Wisconsin glaciation. It was formed when ice of the Green Bay Lobe came into contact with the Baraboo Hills in southwestern Wisconsin and blocked the south-flowing Wisconsin River. During early glacial recession, the ice dam failed catastrophically and the lake drained in about a week. Despite early recognition of the former lake and the likelihood that it failed catastrophically, outflow rates during the failure have not been previously evaluated. Estimates based on step-backwater modeling indicate that peak discharge was between 3.6 and 5.3 × 104 m3/s in the lower Wisconsin River. As an alternate method, we used a previously derived empirical relationship between lake volume and peak discharge for dam-break events. From a digital elevation model altered to incorporate isostatic depression, we estimated the lake volume to be 87 km3 just prior to dam breach, suggesting that the flooding magnitude was as high as 1.5 × 105 m3/s at the outlet. Adjusting these results for downstream flood wave attenuation gives a discharge of around 4.4 × 104 m3/s in the lower reach, which closely matches the results of the step-backwater modeling. These estimates of discharge from the catastrophic failure of ice-marginal lakes improve our understanding of the processes that have produced the morphology and behavior of present-day upper Midwest river systems. © 2007 Elsevier B.V. All rights reserved. Keywords: Glacial Lake Wisconsin; Catastrophic flooding; Geomorphology; Paleohydrology

1. Introduction Proglacial lakes constitute geomorphically-important, yet ephemeral features of glacial landscapes. Limited to a brief existence in geologic time and subject to sudden, catastrophic drainage, these lakes challenge the notion that gradual or recurrent geologic processes maximize geomorphic change. Rather, their tendency to drain in sudden outbursts has resulted in an array of unique and conspicuous landforms which are now solely ⁎ Corresponding author. Tel.: +1 303 541 3041; fax: +1 303 447 2505. E-mail addresses: [email protected] (J.A. Clayton), [email protected] (J.C. Knox). 0169-555X/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.geomorph.2007.03.006

attributed to dam-break flooding. An integration of these extreme floods into models of drainage basin evolution improves our ability to understand the dynamics of landscape change in both glacial and fluvial systems. Glacial Lake Wisconsin was a proglacial lake that formed when meltwater from the Green Bay Lobe of the Laurentide Ice Sheet was ponded in front of the ice margin (Fig. 1). The late Wisconsin ice sheet abutted the Baraboo Hills to block the southeasterly flow of the Wisconsin River and other small drainages in westcentral Wisconsin. Early investigators (e.g., Warren, 1874; Chamberlin, 1883; Alden, 1918; Harloff, 1942) described evidence of the former lake and possible evidence for a large flood from its outlet (Bretz, 1950). More recently, Clayton and Attig (1989) authored a

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Fig. 1. Glacial Lake Wisconsin and its basins. The rectangle in the lower Wisconsin River shows the location of the reach used in the step-backwater modeling. Black lines (and the dashed lines in the inset) indicate ice margins. Modified from Clayton and Attig (1989).

comprehensive report on the lake, including time constraints and field mapping of the various stage indicators, and provided additional support for its catastrophic drainage. This paper builds upon these former investigations by offering the first detailed reconstructions of both the magnitude of the outflow from the lake and the lake's volume during peak stage. Results of this research improve our understanding of the types and magnitudes of former megaflood discharges that contributed significantly to formation of alluvial deposits and terrace and

valley morphologies that characterize many upper Midwest river systems (Matsch, 1983; Kehew and Lord, 1986; Knox, 1987; Teller, 1987; Becker, 1995). 2. Background and setting Evidence for Glacial Lake Wisconsin was recognized by many early researchers. Warren (1874) recognized lacustrine deposits in the main basin and conceived of a large glacial lake with a waterfall at the outlet that

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eventually drained the lake and formed the Wisconsin Dells. Chamberlin (1883), Martin (1916), and Thwaites (1928) identified additional lacustrine deposits and also attributed them to the large-scale ponding of glacial meltwater and the damming of the Wisconsin River. Alden (1918) was the first to refer to the lake as “Glacial Lake Wisconsin” and attributed discrete intervals of organic sediments and marsh deposits to water stage fluctuations related to periodic lake reaccumulation. Harloff (1942) noted that clay deposits in certain areas of the basin had been replaced by alluvial sands and gravels, which may represent evidence for stripping or truncation of the clay sediments during rapid, lakedraining flooding from the lake, even if that possibility was not examined by the author. Already known for his work on the reconstruction of the Glacial Lake Missoula flood, Bretz (1950) found continuous foresets of open-

work, coarse-sized, southward-dipping gravels with extraordinarily large crystalline erratics up to 1.5 m in diameter in the Alloa delta that formed in Glacial Lake Merrimac. He inferred from these deposits that Glacial Lake Wisconsin had been subjected to sudden, catastrophic drainage. Detailed Pleistocene geology maps for individual counties became more widespread in the 1980s, which led to a resurgent effort to explain Wisconsin's geologic history. The first comprehensive report including shoreline and outlet locations, the effects of isostatic rebound, and a drainage chronology was presented by Clayton and Attig (1989). They noted that the coarseness of Alloa delta deposits, anastomosing morphology of the Dells gorges, crossing of former drainage divides by modern Dells area streams, and scouring of main basin lake sediment suggested that the final drainage from the

Fig. 2. Time distance diagram for Glacial Lake Wisconsin. Drainage refers to the direction of flow from the basin occupied by the glacial lake. Phase refers to the Green Bay Lobe ice extent. (a) Before the late Wisconsin glaciation, the Wisconsin River flowed around the eastern edge of the Baraboo Hills. (b) The Green Bay Lobe blocked the former path of the Wisconsin River creating a large lake that may have temporarily drained through Devil's Lake gorge to the lower Wisconsin River. (c) The southern outlet of Devil's Lake gorge was blocked by the Johnstown moraine causing Glacial Lake Wisconsin drainage to flow out the NW outlets to the Black River. (d) The Dells outlet opened between the Main and Lewiston basins as the ice sheet wasted eastward. (e) The Alloa outlet opened as the ice sheet continued to waste eastward, resulting in the catastrophic drainage of the lake. (f) The Wisconsin River re-established a drainage route around the eastern end of the Baraboo Hills (modified from Clayton and Attig, 1989).

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lake was catastrophic. Knox et al. (1982) and Knox and Attig (1988) associated Baraboo quartzite erratics and large (1.4 to 1.8 m) westward-dipping foreset deposits in the lower Wisconsin River valley with the catastrophic drainage of Glacial Lake Wisconsin. The Green Bay Lobe probably reached its maximum extent by 23,000 cal YBP (Colgan, 1999; Dyke et al., 2002). By around 19,000 cal YBP or earlier, enough meltwater and river discharge had accumulated in front of the ice barrier to form a large lake (Clayton and Attig, 1989). Former extents of the lake and its subbasins' circumferences are indicated by relative elevations of lake outlets, shore-ice collapse trenches, shore terraces and deposits, and erosional benches. The Lewiston basin, which bordered the ice, and the main basin connected through the Johnstown moraine by the Dells outlet near the present-day city of Wisconsin Dells (Fig. 1). The outlet apparently formed through a low point in the terminal moraine during initial glacial wasting (Clayton and Attig, 1989). Nearby Glacial Lake Merrimac accumulated between the retreating Green Bay Lobe and valley train deposits that grade westward

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from the position of the Johnstown phase ice (Bretz, 1950; Clayton and Attig, 1990). Because of its 30-m lower elevations, Glacial Lake Merrimac is considered separate from Glacial Lake Wisconsin. Continuous drainage exited through the northwestern Babcock outlets to the Black River for most of Glacial Lake Wisconsin's existence until sometime between the commencement of ice wasting around 14,000 to 16,500 cal YBP and the formation of the Elderon moraine ∼ 13,000 cal YBP (Fig. 2) (Clayton and Attig, 1989). When the ice receded past its interface with the Baraboo Hills, the ice dam was breached; and a height difference of at least 33 m between the NW outlets and the lower, newly formed Alloa outlet (at around 264 m asl) led to the catastrophic drainage of the lake via Glacial Lake Merrimac (Bretz, 1950; Clayton and Attig, 1987; Colgan, 1996). As the lake drained, poorly cemented Cambrian sandstones just upstream from the Dells outlet were deeply incised to form the maze of narrow, anastomosing gorges that constitute the presentday Wisconsin Dells (Fig. 3). Some of these gorges cross former drainage divides, further supporting a rapid

Fig. 3. Flooding routes in the Wisconsin Dells region. The black line indicates the position of the Johnstown moraine, and flooding routes are shown by the arrows.

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scenario for lake drainage (Clayton and Attig, 1986; Attig et al., 2005). The flood deposited a substantial delta into downstream Glacial Lake Merrimac, with 12 m foreset heights; and farther downstream the lower Wisconsin River acted as a spillway for the outburst. At least 50 m of outwash had filled much of the lower valley before being incised by the erosive flooding and abandoned as terrace deposits (Knox, 1982). Also, pebble to bouldersize quartzite erratics found in the lower Wisconsin River basin provide strong evidence for the enormity of the flooding as they were found in the unglaciated Driftless Area underlain exclusively by sedimentary bedrock types (Mickelson et al., 1982; Knox, 1987; Knox and Attig, 1988). The largest of these clasts were probably ice-rafted during the floods; therefore hydraulic competency equations are probably not appropriate for estimating discharges associated with their transport. Still, the boulders provide approximations of former high-water stages that are useful in reconstructing former cross-sectional areas of the lower Wisconsin River channel. Generally speaking, the morphology of the lower Wisconsin River valley is rather rectangular, especially in the portion of the river below Ferry Bluff (Fig. 1), with wide valley bottoms that abut vertical cliffs. The floodplain substrate is ∼95% quartz sand and is extremely welldrained (Socha, 1984). Flood-transported clasts identified by the above researchers and Clayton (2000) showed weak, downstream fining; this trend was probably obscured by ice-rafting during the flood. A series of at least three westward-grading stairstep terraces parallel the main channel, particularly in wider reaches. In the reach from Lone Rock to Boscobel, the most pronounced terraces are found at 18 and 24 m above the modern river (Thwaites, 1928; Knox and Attig, 1988; Clayton and Attig, 1990); additional, less-developed terraces are also present. While sand dunes and loess deposits are widespread on the uppermost terrace, they are generally not found on lower terrace levels (Knox and Maher, 1974). The terraces represent the elevations of valley train deposits that were abandoned by progressive downcutting. The stairstep sequence may have been the result of single or multiple floods; several alternatives are explored below and in detail in Clayton (2000). A precise evaluation of their genetic history is problematic because of the paucity of dateable materials incorporated into the alluvium. 3. Estimation of flood discharge Given the uncertainties inherent in reconstructing a flood from the geologic record, we used two indepen-

dent methods to determine the magnitude of the catastrophic flooding and then compared the results (see Kershaw et al., 2005). Peak discharge was reconstructed (i) using the Army Corps of Engineers' Hydrologic Engineering Center, River Analysis System (HECRAS) step-backwater model to determine the flood profile downstream and (ii) from a reconstruction of the lake volume at the time of the flooding and an empirical relationship between peak discharge and lake volume developed for catastrophic lake-draining events. The approaches are presented separately in the following sections. 3.1. Step-backwater modeling: methods Step-backwater modeling of the flooding from Glacial Lake Wisconsin was performed for a 16 kmlong section of the Wisconsin River, located 51 km upstream from the river's mouth at the Mississippi River and 102 km downstream from the Alloa breach (see Fig. 1). The present-day river valley in this region is sandy and flat-bottomed, with an average valley width of 5 km. We delineated 13 cross sections, spaced around 700 to 2500 m apart, using 1:24,000 scale USGS topographic maps, among which four contained preserved high-water evidence associated with the flooding. A digital elevation map showing the cross section locations and the longitudinal profile through the reach is shown in Fig. 4, with two representative cross section profiles included for reference. This section of the river was chosen for several reasons: (i) it had the greatest density of flood stage indicators downstream from the Alloa outlet, such as ice-rafted quartzite boulders and exposed foreset beds; (ii) it was sufficiently far from the breach to satisfy the assumptions of the model, such as quasi-steady, one-dimensional flow set by the channel bed roughness; and (iii) the geometry of the reach has minimal irregularities, such as expansions or multiple channels. This procedure was complicated by ambiguous geomorphic conditions in the reach at the time of the dam-break flooding. The margins of the relict channel appear to be well-preserved, but the elevation of the channel bed is more difficult to ascertain. Significant downcutting almost certainly occurred during the flood event because the peak discharge would have been immense and sediments that comprise the valley train deposit are roughly gravel-size and have very low cohesion, but the extent of downcutting is unclear due to the lack of dateable material incorporated into the alluvium. In order to allow for a range of results and for simplicity, we modeled the flow using bed elevations

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Fig. 4. Digital elevation model of the reach of the lower Wisconsin River used for the step-backwater modeling. Flow direction is shown by the arrow, and cross sections are shown by the black lines. Water stage indicators were located at cross sections A (Mill Creek site), B (Bock site), C, and D (Coumbe site). Cross section profiles of B and D are given in the lower and upper-right inset diagrams. Both diagrams include the reconstructed water surface and bed elevations for the minimum and medium downcutting geometries, labeled “mincut” and “medcut”, respectively. The thick black line is the modern topography. The lower-left inset diagram gives the longitudinal profile of the modeled reach for the water surface, reconstructed bed geometries, and modern channel thalweg, and uses the same symbols as above.

that were (i) lower than the uppermost terrace level to account for downcutting during the flood; (ii) below the elevations of the high-water surfaces suggested by the position of crystalline erratics, truncated loess deposits, and the sedimentology of bedding structures; (iii) at or above the lowest terrace level where the bedding structures were located; and (iv) above the modern bed elevation of the Wisconsin River. Following these constraints, we selected the first and second terrace surfaces to correspond with two different modeling geometries, hereafter referred to as “medium downcutting” (roughly 15 m below the highest terrace) and “minimum downcutting” (roughly 8 m below the highest terrace), respectively (see Fig. 4). As these terraces were not present in all cross sections, the exact bed elevations were chosen to match at least one cross section's terrace elevation precisely. For example, the “Bock site”

(shown in the lower inset of Fig. 4) shows the match between the lowest terrace and the medium downcutting bed level. The bed elevations at other cross sections were interpolated from the gradient of the terraces through the reach, 0.00022 m/m, which also roughly matched the slope of high-water stage indicators in the area. We used the HEC-RAS step-backwater model to determine the flood discharge through the reach. HECRAS can be used to model steady, gradually varied flow, and can calculate discharge for mixed flow regimes (Corps, 1998). The model can be used to reconstruct former flows by adjusting the discharge iteratively to match the water surface profile evidence. A bed roughness value is required; Manning roughness values of 0.025 and 0.07 were applied to the main channel and overbank zones, respectively, estimated from perceived

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Table 1 Constraints on the water surface in the modeled reach

Less constrained

More constrained

WS (m asl)

XS a

Constraint a

217.3

B

222.5

A

222.6

A

236.0 219.0 222.0

C A D

223.0 225.9

B B

227.7

D

Minimum 5× extrapolated flood depth from foresets Low depth estimate from quartzite #2 from competency equations Low depth estimate from quartzite #1 from competency equations Uneroded loess Location of quartzite boulders Minimum 5× extrapolated flood depth from foresets Location of crystalline erratics Maximum 7× extrapolated flood depth from foresets Maximum 7× extrapolated flood depth from foresets

extrapolated water depth that was five times the height of these structures; thus 217.3 m asl represents the minimum water stage at this location. Also, estimates of competent depths of flow may be derived from clasts in imbricated fluvial deposits. Two large quartzite boulders (intermediate axis of 0.7 and 0.8 m) were found in situ at 219 m asl in the entrance to a tributary valley at cross section A (Fig. 5A). Although unlikely, these clasts may have been transported as bedload during the catastrophic flooding. Using the empirically derived hydraulic competency relations developed by Church (1978) and Costa (1983), we estimated a depth of around 3.5 m at this location. However, these boulders were probably

a See text and Fig. 4 for explanation of cross sections (XS) and constraints on the water surface elevation (WS).

analogues to illustrations given in Barnes (1967) and consistent with paleohydraulic reconstructions in nearby environments (e.g., Matsch, 1983; Lord and Kehew, 1987; Becker, 1995). Interpolated cross sections were added at 200-m intervals; these did not alter the model output per se, but were included to facilitate the calibration. We used a number of constraints for the water surface based on stage indicators found in the reach (Table 1). The field evidence was categorized by quality into “less” and “more constrained” scenarios. The “less constrained” scenario bracketed the probable minimum and maximum water surface elevations, serving mainly to create bounds on what was deemed possible. The “more constrained” scenario employed a narrower range of water stage indicators, representing the lower and upper bounds on what was deemed probable for the elevation of the water surface. Specifically, for the “less constrained” scenario, the water surface was never allowed to surpass 236 m asl at cross section C as this site has a deep, uneroded layer of loess that elsewhere was truncated; this elevation served as the maximum possible flood stage. For the minimum flood elevation, we extrapolated a probable water surface elevation from the height of 1.4 to 1.8-m-high, westward-dipping sedimentary structures found in cross section B (Knox and Attig, 1988). The height of uneroded foreset beds found on the beds of paleochannels may be interpreted to reflect between one-fifth to one-seventh of the water depths responsible for their formation (Allen, 1984; Knox, 1996). As it is possible that these structures were truncated during the flood (Carling, 1996), we used an

Fig. 5. Two examples of water stage indicators in the lower Wisconsin River: (A) large quartzite erratics were found just downstream from Muscoda, in cross section A; and (B) 2.5-m foreset beds were located at cross section D. Downstream is to the right.

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ice-rafted during the flooding; therefore, hydraulic competency equations would greatly overestimate the flow depth. Accordingly, we used the estimates of depth from competency equations for these boulders in the “less constrained” scenario and the exact elevation of the boulders as minimum stage estimates in the “more constrained” scenario. The “more constrained” scenario employed betterquality evidence for water stage and significantly narrowed the range of discharge estimates. A water level of 223 m asl at cross section B was indicated by the elevation of several angular crystalline clasts incorporated into the bed sediment, averaging 0.4 m in size (Knox and Attig, 1988, p. 510). The height of foreset beds found at this location by Knox and Attig (1988), described above, suggest that the peak stage was around 12.5 m above the level of the second terrace, giving a maximum water surface elevation of 225.9 m asl. Farther downstream at cross section D, 2.5-m-high sandy foreset beds were found (Fig. 5B), indicating flows that were locally around 15 m deep. 3.2. Step-backwater modeling: results Modeled discharges for the “less constrained” scenario ranged from 2.3 to 31.1 × 104 and 5.4 to 58.9 × 104 m3/s for the minimum and medium downcutting geometries, respectively. Modeled discharges for the “more constrained” scenario range from 3.6 to 5.3 × 104 m3/s for the minimum downcutting and 14.2 to 20.2 × 104 m3/s for the medium downcutting geometry. Table 2 lists these results and their associated constraints. These are the bounding values—discharge estimates based on other constraints fall between those given in Table 2. Because the relative quality of each water surface constraint could not be assessed within each flow scenario, the “more constrained” discharge estimates were averaged to provide a more general result; the corresponding values are 4.5 and 17.2 × 104 m3/s for the minimum and medium downcutting geometries, respectively. These immense estimates stem mainly from the large cross-sectional areas involved: flow was several kilometers wide for even the smallest discharges modeled, and water depths exceeded 6 m in most areas. The modeled water surface gradient was ∼ 0.0002 for both flow scenarios and steepened slightly from channel narrowing at the downstream end of the reach (Fig. 4). The model predicted reach-average velocities of 1.7 and 2.9 m/s for the minimum and medium downcutting geometries. Compared to some proglacial dam-break floods, these modeled velocities were rather low and probably reflect the low slope through the reach

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Table 2 Discharge (Q) estimates based on the different geometries and flow scenarios Geometry Scenario Mincut

Less constrained

2.3

B

31.1 3.6

C D

5.3

B

5.4

B

58.9 More 14.2 constrained 20.2

C B

More constrained

Medcut

Q XS Constraint (m3/s × 104)

Less constrained

B

Location of crystalline erratics Uneroded loess Minimum 5× extrapolated flood depth from foresets Maximum 7× extrapolated flood depth from foresets Minimum 5× extrapolated flood depth from foresets Uneroded loess Location of crystalline erratics Maximum 7× extrapolated flood depth from foresets

and the effects of flow attenuation from the large (100 km) distance to the breach location. Flow was subcritical in all modeled flows, and modeled reachaverage boundary shear stress was ∼ 10 N/m2 for the minimum downcutting and 23 N/m2 for the medium downcutting geometries. Additional results from the flow modeling are given in Clayton (2000). 3.3. Estimation of peak discharge from a reconstruction of lake volume This study used 30-m resolution U.S. Geological Survey digital elevation models (DEMs) for the state of Wisconsin, compiled by the Wisconsin Geological and Natural History Survey (WGNHS), and the ArcView GIS software package to derive an independent estimate of flood discharge based upon Glacial Lake Wisconsin's volume. Stage estimates corresponding with the Elderon phase ice position just prior to dam failure (Clayton and Attig, 1989, Figs. 12–17) were associated with the maximum lake area and provided the basis for volumetric reconstruction. These authors indicate that the lake stage in the western subbasins was between 286 and 293 m asl before the dam was breached. Elsewhere, Pastor et al. (1982) describe a layer of sand and sandstone cobbles between the bedrock and the Holocene soil at ∼280 m asl, which may represent a flood lag deposit slightly below the Elderon maximum lake stage. However, the present land surface does not properly reflect former differences in local elevation related to isostatic depression during the time of ice advance and maximum lake stage. Although crustal adjustment would have been widespread, the eastern half of the lake bordered the ice margin and would have been

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disproportionately depressed relative to the western half of the lake. During peak lake stage, crustal depression trended roughly 0.16 m/km downward to the NE (Clayton and Attig, 1989; T. Hooyer, WGNHS, personal communication, 2000), although slightly steeper gradients have been reported (Socha, 1984; Guetter et al., 1987). This suggests a relative difference of about 10 m between the eastern and western portions of the basin. To account for this effect, we suppressed the DEM elevations linearly by this gradient along an axis orthogonal to formerly reconstructed isobases for the Green Bay Lobe. The lake basin was then “filled” by selecting all DEM elevations at and below the reported water stage given by Clayton and Attig (1989) for the western portion of the basin, coloring the affected grid cells, and eliminating areas that would have been covered by the ice (Fig. 6). The individual cell distances from the generated water surface to the suppressed ground elevations were summed and then multiplied by the grid cell size to

produce a volume. A more detailed explanation of the procedure is given in Clayton (2000). This resulted in reconstructed lake volumes of 73, 87, and 102 km3 depending on whether the water stage was placed at 286, 289, or 293 m asl—these values encompass the full range of reconstructed water surface elevations given in Clayton and Attig (1989, Figs. 12–17). We elected to use the 289-m asl elevation because it most closely corresponded to the Elderon stage in their figures, but all three possibilities are shown in Fig. 6. The lake perimeter generated by this method overlapped quite well with field evidence for the lake's margins throughout the basin reported by Clayton and Attig (1989). A number of efforts have been made to regress peak discharge from lake volume in catastrophic lakedraining floods (e.g., Clague and Mathews, 1973; Costa, 1988; Walder and Costa, 1996; see discussion in Clague and Evans, 1997; Cenderelli, 2000; Herget, 2005). Of these, we opted to use the empirical relation of

Fig. 6. Glacial Lake Wisconsin, Elderon stage, as reconstructed from GIS modeling.

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Costa (1988, Eq. (19)) because it incorporated dam height and used a large sample size (n = 21) with an r2 of 0.79. The equation is given as 
0:61 Qp ¼ 3:8 H V =106

ð1Þ

where Qp is the peak discharge at the outlet, H is dam height in m, and V is lake volume in m3. We used an ice dam height of 40 m (Clayton and Attig, 1986), which corresponds with the margin of reconstructed ice surface profiles for the Green Bay Lobe developed by Winguth et al. (2004). Using Eq. (1), we predicted a peak discharge of 1.5 × 105 m3/s from the lake; other equations yielded similar results. 4. Discussion In order to compare the results from the step-backwater modeling with those obtained from the empirical relations between lake volume and peak discharge, the peak discharge estimates at the breach must be adjusted for downstream flood wave attenuation. Backwater spillage into tributary valleys, reduction in channel gradient, and cumulative roughness effects would have decreased the magnitude of the flood peak by the time floodwaters arrived at the reach used in the step-backwater calculations (Kershaw et al., 2005). Costa (1988, p. 449) developed an envelope curve to characterize downstream reduction in peak flow for dam burst floods in open valleys as Qx ¼ 100=10ð0:0052xÞ

ð2Þ

where x is the distance downstream from the breach in kilometers, and Qx is the proportion of the upstream peak discharge at location x. The modeled reach is 102 km downstream from the Alloa outlet; Eq. (2) predicts a 71% reduction in the magnitude of peak flow. Given a peak discharge of 1.5 × 105 m3/s at the Alloa outlet (above), the attenuated downstream flood peak would have been 4.4 × 104 m3/s. Interestingly, the discharge estimated by step-backwater modeling for the minimum downcutting geometry (4.5 × 104 m3/s) closely matches the estimate from this analysis; the overlap between discharge estimates is well within the expected error typically associated with paleodischarge reconstructions (Dury, 1985). Though perhaps coincidental, this suggests that the minimum downcutting geometry best characterizes the morphology of the lower Wisconsin River at the time of the flooding. Peak discharge estimates using the medium downcutting geometry are probably too large, considering that they are

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larger than those predicted from the lake volume, even before considering downstream attenuation. Moreover, Glacial Lake Wisconsin's multibasin configuration, in particular the separation of the main basin from the breach area, would have further attenuated the flood wave, resulting in a reduced outflow rate compared with other lakes used in the development of Eq. (1) (L. Clayton, WGNHS, personal communication, 2000). As such, we focused on the minimum downcutting geometry results for the remainder of this paper. The reconstruction of both lake volume and outflow rate allow for an estimation of the flood duration. Assuming constant discharge of 1.5 × 105 m3/s at the Alloa outlet, the 87 km3 lake would have drained in around 6 d. The flood duration probably exceeded this value, however, because the peak discharge typically constitutes a small proportion of the overall flood hydrograph in dam-break events (e.g., Mayo, 1989), and the multibasin configuration of the lake may have prolonged the hydrograph's rising limb. Still, this duration estimate is fairly consistent with estimates from other proglacial lake failures (e.g., Kehew, 1982; O'Connor and Baker, 1992; Menzies, 1995). Assumptions inherent in the reconstruction of the flood duration are detailed in Clayton (2000). Clayton and Attig (1987) envisioned the Alloa outlet to have been reblocked by readvances of the Green Bay Lobe, resulting in the refilling of the lake basin. As mentioned, Alden (1918) attributed discrete lacustrine sediment layers in the basin to lake reaccumulation. The stairstep-like terrace sequence observed in the lower Wisconsin River valley may have been generated from multiple floods—each successive flood eroding the bed and abandoning the former floodplain. If ice advances did re-dam the Wisconsin River, the lake could have refilled quickly from high ablation rates along the glacial margin in the late Pleistocene. Colgan (1996) estimated that during early periods of ice wasting, meltwater discharge along the ice margin of the Green Bay Lobe was 300 to 400 m3/m/d for the 3-mo summer ablation season. Approximately 200 km of ice perimeter contributed to the basin occupied by Glacial Lake Wisconsin. It therefore would have only taken around 15 yr to refill the lake to its former 87-km3 volume. However, an ice dam associated with late-glacial advances would probably not have been thick enough to withstand the pressures exerted by a lake refilled to the same volume. Although the ice surface profile was steep during the glacial maximum when the lake was initially filled (Colgan, 1999; Winguth et al., 2004), during early readvances the ice was probably thinner and more gradually sloping, as evidenced by an increasing

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influence of local topography on ice flow patterns, a transition from predominantly erosional to depositional basal processes, a lack of permafrost at the margin, and a reduction in the regional slope of end moraines (Colgan, 1996; Clayton et al., 1997). The thinner ice sheet and lack of a frozen bed would decrease the ice's ability to re-dam Glacial Lake Wisconsin because high-water levels would challenge the integrity of the thin ice, and high pressure water would be able to bypass the ice dam through a newly-thawed groundwater zone. Thus, if the lake were able to refill somewhat, it probably did not accumulate to former Elderon levels, even though it had adequate time to do so. This supports the hypothesis that, if there were multiple lake-draining floods, the initial flood would have been the largest. The scale of flooding from the drainage of Glacial Lake Wisconsin was quite large compared with the historical flow record. The modeled flow depths in the reach ranged from around 6 to 8 m, and widths from 3600 to 7100 m. For comparison, any flood in the range of 4.5 × 104 m3 /s dwarfs the 1938 peak flood of 2.3 × 103 m3/s from the 95-yr historical record for the Wisconsin River at Muscoda (USGS Gage 05407000), wherein the flow depth and width were approximately 3.7 and 350 m, respectively. However, similar-size floods have been observed elsewhere in historical times (see review in Cenderelli, 2000). In 1972, the Fraser River experienced flows of 1.4 × 104 m3/s near Chilliwack, British Columbia, with a mean depth and width of 8 and 530 m, respectively, and mean velocity of around 3 m/s (M. Church, personal communication, 2007). Glacial Lake George failed in Alaska in 1958, with a peak outflow discharge of around 1.0 × 104 m3/s (Costa, 1988). Also, Clayton and Attig (1989) reasoned that the peak outflow from Glacial Lake Wisconsin might have been roughly equivalent to the magnitude of the 1986 outburst flood at Russell Lake, Alaska, because, although smaller in volume, the gradient of Russell Lake's outlet is steeper than that for the Alloa outlet. Russell Lake stored ∼ 5.41 km3 of water behind its 700 × 150 × 26 m ice dam, and the lake drained in around 30 h with a discharge of ∼ 104 m3/s (Emery and Seitz, 1987; Mayo, 1989). For comparison, Glacial Lake Wisconsin is estimated herein to have stored ∼ 87 km3 of water and drained in around 5 to 10 d with an estimated average discharge around five times that associated with the Russell Lake flooding. Dam-break floods may strongly affect the downstream geomorphology of spillway channels (Baker and Komar, 1987; Costa, 1988; Carling, 1996; Benito, 1997; Cenderelli, 2000; Herget, 2005). We speculate that the stairstep-like terrace sequence in the lower Wisconsin

River valley sequence represents episodic downcutting associated with the initial flood from Glacial Lake Wisconsin and later, smaller events, such as subsequent floods from the same basin or from nearby proglacial lakes (e.g. Montgomery et al., 2004). For example, drainage from Glacial Lake Oshkosh, a comparatively small ice-marginal lake in south-central Wisconsin, may have discharged down the lower Wisconsin River valley after the initial flooding from Glacial Lake Wisconsin, and therefore could be responsible for the degradation of the valley floor to historic levels (Colgan, 1996; Attig et al., 2005). However, many characteristics common to megascale flooding were not found in the Wisconsin River valley. Examples include giant bed ripples, slackwater deposits, kolks, boulder piles, and oversized levees (Baker, 1974; Teller and Thorleifson, 1983; O'Connor and Baker, 1992). These features may have been poorly preserved or may not have been a component of the Glacial Lake Wisconsin flooding. Because the lower Wisconsin River valley temporarily remained the outlet for drainage from the Green Bay Lobe even after the failure of Glacial Lake Wisconsin, post-flood aggradation may have buried gravel dunes or other flood-generated landforms (e.g. Montgomery et al., 2004). Also, differences in floodplain paleovegetation between the Wisconsin River valley and the spillways associated with nearby failed lakes may offer some explanation for the contrasting landform assemblages observed. By the time of the Glacial Lake Wisconsin flooding (Fig. 2), it is likely that some spruce or other cold-weather species had replaced tundra vegetation in protected lowlands such as the Wisconsin River valley, as suggested from radiocarbon dates of spruce pollen found in Devil's Lake (Fig. 1) and the termination of loess deposition (Maher, 1982; Wright, 1987; Baker et al., 1996; Colgan, 1996; Mason and Knox, 1997; Colgan, 1999). If present, dense floodplain vegetation would have significantly reduced the boundary shear stresses associated with peak flood levels (Smith, 2004), and thereby reduced the potential to generate large-scale flood features outside the main channel. Finally, the extremely low gradient of the lower Wisconsin River valley resulted in reconstructed flow velocities well below those necessary to generate megaflood-type landforms (Benito, 1997). For example, the smaller of two lakes that failed catastrophically at the Tsangpo River in Tibet during the early Holocene had roughly the same volume as Glacial Lake Wisconsin, but the spillway gradient was 100 times steeper and probably produced floods capable of substantial bedrock incision in the valley downstream (Montgomery et al., 2004). In the Glacial Lake Wisconsin flooding, unit stream power

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would have been maximized at constrictions, such as at the Dells outlet (Fig. 1) where anastomosing channels were incised into poorly cemented sandstones (Clayton and Attig, 1986; Attig et al., 2005), but would have been below the threshold necessary to generate large-scale erosional features at downstream locations in the lower Wisconsin River valley (Benito, 1997). These and other factors help explain the observed variability in post dam-break flood landscapes (see Carling, 1996; Cenderelli, 2000). 5. Conclusions We have attempted to reconstruct the character of flooding associated with the catastrophic drainage of Glacial Lake Wisconsin—the magnitude of which has never been previously estimated despite the importance of the lake in Wisconsin's geologic history. Unlike many paleohydrologic reconstructions, the paucity of stage indicators in the lower Wisconsin River valley spillway and the lack of a preserved breach cross section render precise discharge estimates unattainable. However, our reconstruction suggests that the flooding from the lake was indeed catastrophic (peak discharge was on the order of 104 m3/s), and that the lake, which likely took around 15 yr to fill, may have drained in around a week. Field support for the extremity of the flood is evidenced by the (i) high topographic positions of the transported clasts and other flood stage indicators; (ii) enormous size of particles moved in the delta area; (iii) overall coarseness of transported sediments; (iv) spillway geomorphology of the lower Wisconsin River; (v) massively incised and anastomosing Dells gorges; (vi) simultaneous and unidirectional drainage from the lake's southwestern subbasins; and (vii) regional scouring of the lake basin's lacustrine veneer. Additional geomorphic features used to characterize the scale of catastrophic flooding from other proglacial lakes, e.g. giant bed ripples and kolks, were absent in this case; possible explanations include either the poor preservation of these features or insufficient stream power during peak flow to generate such landforms. The incorporation of proglacial lake floods into the developmental chronology of the upper Mississippi River basin provides a more realistic, process-based understanding of upper Midwest drainage development. During the late Pleistocene, the Mississippi River experienced multiple highly erosive floods that originated from rapid drainage of proglacial lakes (Willman and Frey, 1970; Wright, 1987; Becker, 1995; Knox, 1996). The geomorphology and hydrology of modern river systems of the Upper Mississippi Valley reflect and

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continue to be influenced by these relict flood events. Clearly, more information about the paleohydrology of these former floods provides significant improvement in our understanding of the behavior of present-day river systems in the region. Acknowledgements We would like to thank Vance Holliday, Dave Mickelson, Lee Clayton (no relation to first author), Jim O'Connor, Richard Waitt, Victor Baker, Michael Church, and an anonymous reviewer for their thoughtful revisions of earlier versions of this document. Thomas Hooyer generously shared unpublished research ideas and insights. Chip Hankley assisted with the DEMs and ArcView software. Michael Turaski and Martin Chy provided valuable field assistance. Furthermore, this work would not have been possible without the cooperation and trust of various land and gravel pit owners along the Wisconsin River. The investigation of paleoflood deposits and alluvial stratigraphy in the lower Wisconsin River valley was supported in part by National Science Foundation grants EAR-9206854 and ATM-0112614. References Alden, W.C., 1918. The Quaternary Geology of Southeastern Wisconsin with a Chapter on the Older Rock Formations. U.S. Geological Survey Professional Paper, vol. 106. Govt. Print. Off., Washington, D.C. Allen, J.R.L., 1984. Sedimentary Structures: Their Character and Physical Basis. Elsevier, New York. Attig, J.W., Hooyer, T.S., Mode, W.N., Clayton, L., 2005. Glacial Lakes Wisconsin and Oshkosk—two very different late-glacial ice-marginal lakes in Wisconsin. Geological Society of America, Abstracts with Programs 37 (5), 22. Baker, R.G., Bettis, E.A., Schwert, D.P., Horton, D.G., Chumbley, C.A., Gonzalez, L.A., Reagan, M.K., 1996. Holocene paleoenvironments of northeast Iowa. Ecological Monographs 66 (2), 203–234. Baker, V.R., 1974. Erosional forms and processes for the catastrophic Pleistocene Missoula floods in eastern Washington. In: Morisawa, M. (Ed.), Fluvial Geomorphology. Proceedings of the 4th Annual Geomorphology Symposia Series, Binghamton, New York, pp. 124–148. Baker, V.R., Komar, P.D., 1987. Cataclysmic flood processes and landforms. In: Graf, W.L. (Ed.), Geomorphic Systems of North America: Centennial Special Vol. 2. Geological Society of America, Boulder, CO, pp. 423–443. Barnes, H.H., 1967. Roughness characteristics of natural channels. U.S. Geological Survey Professional Paper, vol. 1849. Washington, D.C. Becker, W.M., 1995. Reconstruction of late-glacial discharges in the Upper Mississippi Valley. M.S. Thesis, University of Wisconsin, Madison, 100 pp. Benito, G., 1997. Energy expenditure and geomorphic work of the cataclysmic Missoula flooding in the Columbia River gorge, USA. Earth Surface Processes and Landforms 22, 457–472.

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