Ice streams of the Late Wisconsin Cordilleran Ice Sheet in western North America

Quaternary Science Reviews 179 (2018) 87e122

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Ice streams of the Late Wisconsin Cordilleran Ice Sheet in western North America Nick Eyles*, Lina Arbelaez Moreno, Shane Sookhan Department of Physical and Environmental Sciences, University of Toronto at Scarborough, 1265 Military Trail, Scarborough, ON M1C 1A4, Canada

a r t i c l e i n f o

a b s t r a c t

Article history: Received 20 July 2017 Received in revised form 12 October 2017 Accepted 18 October 2017 Available online 20 November 2017

The Late Wisconsin Cordilleran Ice Sheet (CIS) of western North America is thought to have reached its maximum extent (~2.5
106 km2) as late at c. 14.5 ka. Most (80%) of the ice sheet's bed consists of high mountains but its ‘core zone’ sited on plateaux of the Intermontane Belt of British Columbia and coterminous parts of the USA, shows broad swaths of subglacially-streamlined rock and sediment. Broad scale mapping from new digital imagery data identifies three subglacial bed types: 1) ‘hard beds’ of variably streamlined bedrock; 2) drumlinized ‘soft beds’ of deformation till reworked from antecedent sediment, and 3) ‘mixed beds’ of variably-streamlined bedrock protruding through drumlinized sediment. Drumlins on soft beds appear to be erosional features cut into till and antecedent sediments, and identify the catchment areas of paleo ice streams expressed downglacier as flow sets of megascale glacial lineations (MSGLs). ‘Grooved’ and ‘cloned’ drumlins appear to record the transition from drumlins to MSGLs. The location of paleo ice streams reflects topographic funneling of ice from plateau surfaces through outlet valleys and a soft bed that sustained fast flow; rock-cut MSGLs are also present locally on the floors of outlet valleys. CIS disintegrated in <1000 years shortly after c. 13.0 ka releasing very large volumes of meltwater and sediment to the Pacific coast. Abrupt deglaciation may reflect unsustainable calving of marine-based ice streams along the glacio-isostatically depressed coast; large deep ‘fiord lakes’ in the ice sheet's interior may have played an analogous role. Mapping of the broad scale distribution of bed types across the Cordilleran Ice Sheet provides key information for paleoglaciological modelling and also for understanding the beds of modern ice masses such as the Greenland Ice Sheet which is of a comparable topographic setting. © 2017 Elsevier Ltd. All rights reserved.

Keywords: Cordilleran Ice Sheet Drumlins Megascale glacial lineations Paleo ice streams

1. Introduction In western North America the Late Wisconsin (Fraser Glaciation) Cordilleran Ice Sheet (abbreviated as CIS throughout this paper) (Figs. 1 and 2) is thought to have reached its maximum extent at ~14.5 ka (Fulton et al., 2004) though age data are overwhelmingly from the former ice sheet's southernand most accessible margin and recent climate models suggest a more complex history (Menounos et al., 2017). If applicable to the entire ice sheet this timing is considerably later than the adjacent Laurentide Ice Sheet (LIS) which reached its maximum at c. 18e20,000 ka, and is also out of phase with hemispheric climate trends (Seguinot et al., 2016). The timing of the subsequent deglaciation of CIS is better constrained and occurred early and was extremely rapid (completed by

* Corresponding author. E-mail address: [email protected] (N. Eyles). https://doi.org/10.1016/j.quascirev.2017.10.027 0277-3791/© 2017 Elsevier Ltd. All rights reserved.

c. 10.0 ka) when compared to the larger ice sheet to the east. The Late Wisconsin CIS was approximately 2.5
106 km2 in area (i.e., approximately 50% larger than the modern Greenland Ice Sheet; 1.7
106 km2) but was possibly smaller than ancestral preWisconsin Cordilleran ice sheets (Hidy et al., 2013) reflecting possible long term tectono-topographic changes in western North America (e.g., Chamberlain et al., 2012). The likelihood of fast flowing ice streams being present within the Late Wisconsin CIS has been suggested (Booth et al., 2004; Clague, 1984; Eyles et al., 1990; Hicock and Fuller, 1995; Evans, 1996; Marshall et al., 1996; Mathews, 1991; Stumpf et al., 2000) but to date, there has been no systematic ice sheet-wide search for the characteristic subglacial footprint of ice streams in the form of megascale glacial lineations (MSGLs; see Campo et al., 2017; Clark, 1993; Clark and Stokes, 2013; MacLean et al., 2017; Stokes, 2011; Stokes and Clark, 2002). This is in contrast to the situation pertaining to the much larger LIS, where some 110 paleo ice streams have been so far been identified from mapped flow sets of MSGLs (e.g., Margold et al., 2015; Stokes et al.,

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2017) underscoring the need for a wide ranging geomorphological assessment of its bed types, their geology and spatial variation. 1.1. Purpose of this paper The principal objective of the present contribution is to show how combined geological and geomorphological studies can advance paleoglaciological reconstructions of CIS by identifying the principal subglacial terrain types left by the ice sheet. Subglacial bedforms and their underlying composition provide first-order data on the mode(s) of flow and thermal regime of ancient ice sheets. In particular, important information regarding the broad range of former ice flow velocities can be determined from analysis of subglacial bedforms such as drumlins, known to be present over larger areas of the bed of the CIS (Armstrong and Tipper, 1948) and which are now considered to be formed under slow to moderate ice flow velocities perhaps somewhere between approximately 100 and 400 m yr1 (e.g., Ely et al., 2016a). By contrast, highly elongated streamlined ridges and grooves (MSGLs) record the trunks of paleo ice streams that by analogy with modern ice sheets were flowing at much greater velocities as much as 10 km each year (Margold et al., 2015). The paper commences with an outline of the physical setting, basal topography and known history of CIS briefly emphasizing major data gaps in understanding its structure and glaciology. This is then followed by a review of the digital topographic databases and methods used in this study to produce digital images for mapping the geomorphology and geology of the ice sheet's subglacial footprint which clearly identify numerous flow sets of MSGLs left by paleo ice streams. A discussion section explores the many geological factors that may have initiated ice streaming and the possible relationship between fast flowing ice streams along the glacio isostatically depressed Pacific margin of the ice sheet, and its rapid collapse. An additional motivation for this work was to generate understanding of the various bed types present across an entire paleo-ice sheet which might then serve as an analog for modern ice sheets of comparable size and geologic setting, such as for example the current Greenland Ice Sheet where there is considerable discussion (and uncertainty) of the nature of possible bed types and their glaciological role. New digital imagery date now allow assessment of subglacial sediments and landforms across broad swaths of the beds of former ice sheets, which is an essential first step in integrating glacial geology and glaciological data for glaciological modelling. 2. Physical setting and history of the Cordilleran Ice Sheet Fig. 1. Simplified morphogeologic regions of the Cordillera of western North America after Gabrielse et al. (1992), Holland (1964) and Mathews (1986) with superposed very generalized outline of the Cordilleran Ice Sheet (after Fulton et al., 2004) at its maximum Late Wisconsin extent at about 14.5 ka. The core zone of the ice sheet primarily lay across relatively low relief plateau and lowlands of the Intermontane Belt. Note this study does not include the far northwestern extremity of the ice sheet developed over the Wrangell and St. Elias Mountains (see Fig. 2).

2016), and in regard to other late Pleistocene ice sheets (e.g., Putkinen et al., 2017; Ottesen et al., 2005). While much work has described and dated glacial deposits left by the CIS (see Booth et al., 2004; Easterbrook, 2003; Easterbrook et al., 2003; Fulton et al., 2004; Stumpf, 2003 for summaries) the lack of fundamental information regarding its paleoglaciology and drainage is a barrier to physical modelling of its growth and rapid decay. It has recently been said that the Cordilleran Ice Sheet is the least understood among Pleistocene ice sheets in terms of its extent, volume, and dynamics (Seguinot, 2014; Seguinot et al., 2016; Montelli et al.,

The Late Wisconsin Cordilleran Ice Sheet, named by T.C. Chamberlin in 1895, extended across a topographically and geologically diverse area of high alpine mountains, highlands, plateaux uplands and lowlands, and deeply incised valleys and lake basins, and extended to the Pacific coastal margin which is heavily indented by fiords and offshore troughs. This topographic variation gives rise to a correspondingly wide range of subglacial and proglacial environments and deposits that accumulated through the course of any one glacial cycle resulting in a complex stratigraphic record. Much of the primary glacial record too, has been lost by postglacial reworking by rivers along confined valleys and by marine processes. The size of the ice sheet together with large areas of rugged inaccessible terrain and dense forest cover, constrains geomorphological fieldwork and thus glaciological modelling on an ice sheet scale; as this study demonstrates, this constraint can be overcome in part, by using recently available digital elevation data and imagery.

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The varied topography across the Cordillera reflects a complex geologic history resulting from tectonic activity along the active margin of the North American plate (the NAp). The bed of the ice sheet extends across five principal morphogeological belts that make up the Cordillera (Figs. 1 and 2). These belts reflect a protracted history of crustal accretion and strike-slip movement of fartravelled terranes and resulting eastward-directed thrusting that commenced when the NAp began to move west after the break-up of Pangea some 200 million years ago. From west to east, these belts are the Insular Belt, the Coast Belt dominated by the Coast Ranges, the Intermontane Belt comprising the Interior, Stikine and Yukon plateau and the Skeena Ranges, the Omineca Belt composed of the Cassiar, Omineca, Columbia (Purcell, Selkirk, Cariboo), Monashee and Selwyn mountains and the Shuswap and Okanagan highlands, and lastly, the Foreland Belt located along the eastern extremity of the ice sheet dominated by the high Rocky, Mackenzie and Franklin mountains (Gabrielse et al., 1992). Not surprisingly the regional topographic setting is very closely reflected in the broader paleoglaciological structure of CIS whose north-south elongated central ‘core zone’ formed across relatively low relief plateaux of the Intermontane Belt and adjacent lowlands (Figs. 1 and 2). The core zone is rimmed to the west by the high Coast Ranges (<4000 m above sea level; masl) and to the east by the high (<3500 masl) mountains of the Omineca and Foreland belts (e.g., Fulton et al., 2004; Jackson and Clague, 1991; Stumpf et al., 2000). The basal topography of CIS with its central core zone rimmed by high mountains closely resembles the bed of the Greenland Ice Sheet (Bamber et al., 2013a; Rippin, 2013) such that significant parallels can be drawn between the two ice sheets (see Discussion below). CIS was possibly as much as ~2 km thick at the time of its maximum volume during the Late Wisconsin Vashon Stade of the Fraser Glaciation in the Pacific Cordillera some 14,500 years ago (also known as the Pinedale Glaciation in the central Rocky Mountains of the USA; Fullerton et al., 2004). It was about 1000 km wide when its eastern margin expanded out from the Front Ranges of the Rocky Mountains into the Interior Plains where it either locally abutted the Laurentide Ice Sheet (LIS) or terminated as terrestrial piedmont lobes or valley glaciers (e.g., Bednarski and Smith, 2007; Bobrowsky and Rutter, 1992; Fullerton et al., 2004; Jackson et al., 1991; McCuaig and Roberts, 2002; Robert, 1991; Holme et al., 2000; Stroeven et al., 2010, 2014 and refs therein). The most vigorous accumulation areas lay in the west close to sources of Pacific moisture along the Coast Ranges and Skeena Mountains and these fed numerous outlet glaciers that flowed through narrow valleys and fiords to terminate within offshore troughs (e.g., Chatham Strait, Dixon Entrance; Figs. 2 and 3). Ice from the high Coast Ranges and Skeena Mountains also flowed eastward into the interior and coalesced with ice forming across the large north-south elongated ‘core zone’ of the ice sheet in the lower elevation (<1400 masl) Intermontane Belt. Geological estimates suggest ice thicknesses of about 600 m in the core zone (e.g., Huntley and Broster, 1996) and possibly much thicker (Seguinot et al., 2016). The Intermontane Belt consists of large areas of moderate relief (~200 m) plateau surfaces formed by the successive eruption of Tertiary flood basalts that now comprise the Nechako, Chilcotin and Cariboo plateau, which in British Columbia are collectively known as the Interior Plateau (Figs. 1 and 2) (Mathews, 1989) and as the Columbia Plateau in adjacent Washington State and its northward continuation, north of the Columbia River, as the Omak Plateau (Carlson and Hart, 1988; Hooper et al., 2007). In a very important contribution, Armstrong and Tipper (1948) mapped very broad swaths of drumlinized sediment and subglacially streamlined rock across the core zone of the Interior Plateau in British Columbia indicating sustained flow from the Coast Ranges

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toward the interior and the Nechako Lowlands (see also Arnold et al., 2016; Sacco, 2012; Stumpf et al., 2000) (Fig. 2) and identified a very strong topographic control on ice flow borne out by recent studies of ice flow directions (Arnold and Hickin, 2017). The changing location of ice sheds as CIS thickened and thinned resulted in abrupt and complex variations in ice flow direction through time (see Andrews et al., 2012; Fulton et al., 2004). A major and apparently long-lived ice divide near Williams Lake in the Fraser River valley separated north easterly from southerly flowing ice (Armstrong and Tipper, 1948, Figs. 2 and 3). The latter moved southwards across the Thompson Plateau and was then topographically forced through narrow structurally-controlled outlet valleys that drain south into what is now the states of Washington and Montana in the northern USA (Richmond, 1986). These outlet glaciers terminated as several large piedmont lobes (from west to east the Okanogan, Sanpoil, Columbia River, Colville, Little Spokane, Priest River, Lake Pend Oreille, Bull River, Thompson River and Flathead lobes; Fig. 4). The largest ice lobe (Okanogan) left a prominent end moraine (Withrow Moraine in Washington State) on glacially-scoured flood basalts of the Omak Plateau. Southflowing ice lobes dammed a series of deep interconnected lakes (e.g., glacial Lake Missoula) (Fig. 4) that drained catastrophically and repeatedly down the Columbia River valley to the Pacific (see Hanson et al., 2012 and refs therein). Along the eastern margin of the ice sheet, several smaller piedmont lobes extended east from the Rocky Mountains and the main body of CIS out onto the plains most notably immediately east of present day Glacier National Park area in western Montana. The largest of these was the Two Medicine Glacier which was a piedmont lobe about 50 km wide and extended 55 km east from the front of the Rocky Mountains (Alden, 1932; Carrara, 1989; Fullerton et al., 2004, Fig. 3). A wealth of information on the timing of the maximum extent and deglaciation of CIS has been generated from local or regional scale stratigraphic work principally along the ice sheet's southern margins (e.g., Booth, 1987; Booth et al., 2004; Fulton et al., 2004; Richmond, 1986; and refs therein). The timing of the maximum extent of the ice sheet, at least in the south, is markedly out-ofphase with that of LIS and hemispheric paleoclimate (Seguinot et al., 2016). The CIS attained its maximum volume much later in the last glacial cycle (c. 14.5 ka) compared to the LIS (20e18.0 ka) and in addition, the ice sheet had abruptly disappeared by 10.0 ka with the exception of ice remaining on high mountain massifs (Carrara, 1989; Dyke, 2004; Dyke et al., 2002, 2003; Fulton et al., 2004; Kaufman and Manley, 2004; Lakeman et al., 2008). 2.1. This study The dominant subglacial bedform occurring across the former core zone of CIS are drumlins which occur in their many thousands on basalt plateaux of the Intermontane Belt (e.g., Armstrong and Tipper, 1948). New digital imagery now reveals additional information on the morphology of these bedforms, in particular the common presence of flow sets of drumlins that are transitional downglacier to megascale glacial lineations (MSGLs). These were initially identified on LiDAR data sets by Eyles et al. (2016) on the bed of the Puget Lobe in the southwest sector of CIS and record the onset zone and downstream trunk of a large paleo ice stream in the Strait of Georgia (Fig. 5). This discovery provided first-order geomorphological confirmation of the results of much earlier numerical modelling of fast ice flow (~600 m yr1) of a Puget Sound Lobe (Brown et al., 1987). The same distinctive subglacial geomorphic footprint of paleo ice streams in the form of well-developed MSGLs was also recognized on the bed of the Okanogan Lobe near Withrow (Fig. 6) and that of the Flathead Lobe near Kalispell

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Fig. 2. Physiography of the bed of the Cordilleran Ice Sheet (maximum extent outlined). Note the presence of large elevated interior ranges (Skeena Range) and interlinked plateaux of the Intermontane Belt (Fig. 1) surrounded by high mountain belts to the west (Coast Ranges) and east (Rocky Mountains, Liard, Omineca, Cassiar, Columbia Ranges, etc.).

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Fig. 3. Generalized ice flow directions (A) based principally on Carrara et al. (2007), Bednarski (2015) and Arnold et al. (2016) showing outflows from the core zone of the Cordilleran Ice Sheet developed across the Intermontane Belt (outlined). ‘R’ identifies a large area of Rogen moraine under an ice divide located near Williams Lake, British Columbia (see Fig. 21). B: Index map showing locations of representative examples (with figure numbers) of subglacial bed types formed below the Cordilleran Ice Sheet. The maximum extent of ice sheet is outlined.

(the Eureka Drumlin Field) (Fig. 7) indicating that some of the piedmont lobes within the southern sector of CIS (Fig. 4) were fed by topographically-controlled ice streams. This finding was the prime motivation for a more comprehensive assessment of the number of paleo ice streams and their glaciological importance across the entirety of the Cordilleran Ice Sheet. This is now feasible given the recent availability of digital topographic databases which are key to geomorphological mapping of the ice sheet's bed in remote and otherwise inaccessible locations. It is emphasized that the study reported herein does not include the full footprint of CIS at its maximum extent since it excludes ice

covers on its northern flank that developed over high mountain massifs such as the Wrangell Mountains and the St. Elias Range adjacent to the Gulf of Alaska coastal margin, and further west along the Alaskan Peninsula. These ice masses were largely separate from the main body of the CIS (e.g., Ely et al., 2016b) (Figs. 1e3) where the local terrestrial record of glaciation is dominated not by subglacial deposits (e.g., tills) but by glaciolacustrine, glaciomarine and glaciofluvial sediments along the floors of confined valleys reflecting the destruction of primary glacial sediments and landforms by meltwater and sediment gravity flow (see Bennett et al., 2002; Eyles, 1987; Powell and Molnia, 1989). Much glacial

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Fig. 4. Map of the southern margin of CIS based principally on Richmond (1986) and Booth et al. (2004). Megascale glacial lineations occur on the beds of the Puget Lobe (Fig. 5), Okanogan (Fig. 6A and B, Fig. 12) and Flathead lobes (Fig. 7) together with the Two Medicine piedmont lobe (Fig. 24) indicating that these lobes are the termination of ice streams (Fig. 20).

sediment completely bypassed interior valleys to accumulate offshore in the Gulf of Alaska basin (e.g., the <5 km thick Yakataga Formation) (Eyles, C.H. 1987; Eyles and Vossler, 1992; Lagoe et al., 1993; Montelli et al., 2017; see Discussion below). A major limitation of this study is that there is very little information available regarding the submarine geomorphology of most of the Pacific Coast continental shelf margin of the ice sheet which hampers assessment of the full population of former paleo ice streams that flowed through coastal fiords and offshore troughs (see Swartz et al., 2015; Montelli et al., 2017). Similarly, in the ice sheet's interior geomorphic evidence for the presence of some ice streams has undoubtedly been destroyed by lateglacial and postglacial fluvial activity in glacially-overdeepened valleys now flooded by deep ‘fiord lakes’ (e.g., Adams, Shuswap, Okanagan, Pend’Oreille) akin to their marine counterparts along the Pacific Coast (see below). Consequently, this paper presents only an initial, provisional assessment of the number of ice streams within CIS. 3. Databases and methodology The primary databases used for mapping the subglacial geomorphology of the CIS consist of a series of Digital Elevation Models (DEM) compiled from data downloaded from Natural Resources Canada and the United States Geological Survey (USGS) (Fig. 8). Analysis and manipulation of these data sets allows identification and mapping of glacial landforms and their grouping within a restricted number of recurring subglacial terrain types across the former bed of CIS; the basic methodology has been outlined by Bennett and Glasser (2009, p. 348). Follow-up ground surveys were conducted across the core zone of the ice sheet in central and southern British Columbia and adjacent parts of the northern USA to confirm geomorphological interpretations and to determine the geology of the ice sheet's bed using traditional field sedimentological methods at outcrops exposing the interiors of various subglacial bedforms. The analytical methodology followed in this study and the various steps involved in manipulation of digital data and analysis

of resulting images is based off of work outlined in Yu et al. (2015) and Sookhan et al. (2016) and is depicted schematically in Fig. 8. Initial DEM data were retrieved from the Natural Resources Canada GeoGratis server for British Columbia and Yukon. These data are composed of 240 bare earth DEM tiles of 2
2 with 20 m resolution from the Canadian Digital Elevation Model (CDEM) compiled by the Earth Sciences Sector, Mapping Information Branch, Centre for Topographic Information - Sherbrooke of Natural Resources Canada. Data were also generated from pre-existing Canadian Digital Elevation Data (CDED) digitized from National Topographic Database maps at 1:50 000 scale published between 1986 and 1987. The tiles have a vertical accuracy ranging from 0.5 m to 48.6 m Root Mean Square Error (RMSE) and an average horizontal accuracy ranging from 2.5 m to 35 m RMSE. For DEM data for the states of Washington and Montana, the USGS National Map Download server was used and comprised 45 seamless bare earth DEM tiles of 1
1 degree from the National Elevation Data (NED) of 1/3 arcsecond (12 m) resolution. Data were produced by the USGS by merging elevation data from different projects and tiles have a relative vertical accuracy ranging from 1 to 6 m RMSE. Certainly, many subglacial bedforms are smaller in height and relief amplitude than this but the present study is not primarily concerned with morphometric mapping of single bedforms but rather regional assessment of bedform types to identify former ice streams. All topographic data were processed using ArcGIS 10.0. DEM tiles were stitched together using a Mosaic to New Raster Data Management Tool and then projected using Canada Albers Equal Area Conic projection in order to calculate areas and lengths of streamlined terrain accurately given that the data spans more than one UTM zone (Z9-12). Statistics and Pyramids were calculated to allow for visualization of such a large dataset at various scales. A perceptually uniform colour palette with a smooth transition from dark blue to green to yellow was used to work with the DEM in order for changes in elevation to be accurately displayed as a continuous change in colour (or shade, in black and white printing). It also emphasizes the difference between the highs on the peaks of

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Fig. 5. Megascale glacial lineations on the bed of the Georgia-Puget Ice Stream (# 1: Fig. 20), West of Case Inlet (A) and South of Hood Canal (B). Ice flow direction to southwest. Note widespread presence of ‘cloned’ drumlins where parent bedform has been partially dissected to form more elongate MSGLs (see Fig. 23) marking acceleration of ice flow velocity to streaming flows.

the mountains and the valley floors where the potential for ice streaming is greatest. A hill-shaded DEM was then constructed using the spatial analyst tool at four times the vertical exaggeration and default values of azimuth (315 ) and of altitude (45 ). This hillshaded DEM was then overlain by the coloured DEM in order to create the appearance of a 3D surface to aid identification of landforms and terrains (Fig. 8). Hill-shaded DEMs were divided into two sets of quadrants of different size; 900 km2 quadrants at 1:125 000 scale, and smaller 100 km2 quadrants at 1: 50 000 scale in order to identify broader glacial terrain and landform types respectively. In this process, the relatively small 100 km2 quadrants were used to identify and map individual glacial bedforms smaller than 3 km in length; this

information was then scaled-up and used to identify and map the predominant subglacial terrain type(s) within the larger 900 km2 quadrants. At all stages of the process, supplemental data were collected by visual inspection of other satellite imagery and geological reports to confirm or modify identification of bedforms and terrain types. Additional glacial geological data needed to aid identification of subglacial bed types and their geology were derived from surficial geology maps for Yukon, British Columbia and Washington State which were retrieved from the Yukon Geological Survey, British Columbia Geological Survey, Natural Resources Canada, State agencies in the US and the United States Geological Survey. An additional shape file pertaining to Quaternary geology was

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Fig. 5. (continued).

constructed primarily from Fulton (1995) using vector data from the Geological Survey of Canada, to show the broad range of Pleistocene sediment types present across the bed of the ice sheet; further geological and geomorphic data were used from many other published sources (Table 1). 3.1. Terminology The primary objective of this study was to identify and map paleo-ice streams within CIS by reference to diagnostic subglacial landforms such as corridors of MSGLs which are widely agreed to be the footprint of fast flowing ice streams (e.g., Campo et al., 2017; Clark and Stokes, 2013; Ely et al., 2016a; MacLean et al., 2017; Margold et al., 2015; Putkinen et al., 2017 and many refs therein). The upstream parts of the ice stream's catchment (above onset zones) is commonly dominated by drumlins suggesting more sluggish flowing ice; these bedforms are conventionally separated from MSGLs on the basis of elongation ratios (where ER ¼ bedform length/width ratio) calculated in this study by visual inspection of the most elongate landform in any one 100 km2 sized quadrant. Thus streamlined bedforms with ERs ranging from 2 to 10 are classified as drumlins, and as megaridges in those cases where their elongation ratio is greater (e.g., Stokes and Clark, 1999; Clark et al., 2009; Ely et al., 2016a; Spagnolo et al., 2014); the ER of MSGLs mapped in this study is sometimes in excess of 20. It is stressed that this simple differentiation is arbitrary and artificially divides what is a bedform continuum; it is used purely for assessment purposes of the principal subglacial bedforms types across a large area of the former bed of CIS; much more detailed work is now clearly required on identification of the transitional bedforms between drumlins and MSGLs and also to establish appropriate boundary ice flow velocities for subglacial bedforms in terms of a bedform hierarchy. In regard to mapping dominant bed types across a large area, some simplification of existing published map data was required. Glacial deposits and landforms are inherently complex spatially

and glacial geologists simplify the process by employing a ‘land system approach’; the terms ‘depositional system’ or ‘facies model’ are essentially synonymous. The fundamental underlying premise is that landforms are genetically related to underlying sediments since they accumulated coevally during the same depositional event (e.g., a single glacial cycle) (see Benn and Evans, 2010; Evans, 2013; Eyles et al., 1983 and refs therein). The concept has proved very useful as an organizing framework and a number of glacial land systems are now recognized, such as the drumlinized ‘subglacial land system’ that consists of drumlinized terrain underlain by subglacially-deposited till and perhaps associated with sluggish flowing ice (~100e400 m yr1?) (Ely et al., 2016a; Eyles, 1983; Eyles and Eyles, 2010; Spagnolo et al., 2016; Stokes et al., 2013). Another is the ‘ice stream land system’ which emphasizes the presence of MSGLs formed under faster flowing ice that may have velocities of as much as 10 km yr-1 (Benn and Evans, 2010; Clark and Stokes, 2013; Eyles and Doughty, 2016; Otteson et al., 2016; Ross et al., 2009, 2011; Stokes, 2011). Associated ice stream geomorphic features are large end moraines built largely of subglacial till that was advected to the ice margin, and shear moraines formed where fast flowing ice abuts more sluggish flowing zones (e.g., Evans et al., 2008, 2014). While a land system approach is very useful (see Benn and Evans, 2010; Evans, 2013 for comprehensive reviews) the assumptions underpinning the method are not always appropriate in areas of thick sediment covers or multiple glaciations where older glacial and non-glacial deposits are preserved and partially reworked by younger ice masses. This consideration is felt to be especially relevant to the present study given the widespread occurrence of thick pre Late Wisconsin deposits in topographic lows across the Intermontane Belt below the core zone of CIS. In this area, it is apparent that many drumlins (and genetically-related MSGLs) appear to be primarily erosional in origin because they are cut across underlying materials that can include both sediment and rock (see below and Eyles et al., 2016; Krabbendam et al., 2016 and

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Fig. 6. Rock-cut drumlins and megascale glacial lineations (A) on the hard bed of the Okanogan Ice Stream (#3; Fig. 20) composed of flood basalts of the Omak Plateau north of the Columbia River (see Fig. 3B and Fig. 4 for location). Ice flow direction to south. B: Megascale glacial lineations cut into rock.

refs therein). In such cases the subglacial bed and its bedforms (drumlins, MSGLs) can be recognized as a first-order glaciallystreamlined unconformity cut subglacially across geneticallyunrelated and non-coeval sediments (and/or rock). At the same time, there is growing recognition that the internal composition of drumlins unlikely to be genetically related in any simple fashion to the processes that formed the bedforms (Stokes, 2011; Stokes et al., 2016). Consequently, this paper simply recognizes a variety of ‘subglacial terrains’ that discriminate former subglacial bed types of markedly different geomorphology, rheology and thus paleo ice flow of CIS. It is emphasized that recognition of paleo ice streams in the present study is based solely on the presence or absence of megascale glacial lineations; other landforms known to be associated with MSGLs (such as large end moraines, shear moraines; Stokes, 2011) are not common across the bed of CIS possibly as a result of ice streams terminating in areas of confined topography or large water bodies where primary landforms were reworked by meltwaters during rapid collapse and melt of the ice sheet. This aspect of preservational bias is discussed further below.

4. Results The largest portion of the former bed of CIS mapped in this study (~1.2
106 km2) consists of linear belts of high-standing mountains and deep confined valley networks cut into a wide range of bedrock types and geologic structures. High mountains surround the core zone of the ice sheet across the Intermontane Belt creating an elevated rim (Figs. 1 and 2). The highest parts of this mountain rim were only (and very briefly) ice-covered at times of maximum ice sheet volume with many exposed nunataks possibly covered by thin cold-based ice. As a consequence, much of the former bed of the ice sheet has been exposed to long-term nival and periglacial processes together with mass wasting processes, especially landsliding. This is reflected in the long distance supra- and englacial transport of coarse-grained talus as medial and lateral moraine debris by outlet glaciers (e.g., Church and Ryder, 2010; Jackson et al., 1997). Deeply-cut narrow valleys with often oversteepened side slopes, are commonly controlled by faults and other structures, and are often part of well-defined dendritic tributary

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Fig. 6. (continued).

systems draining to major trunk outlet valleys where the effects of glacial erosion and deposition is much greater. Smoothed, glacially abraded valley floors at lower elevations often resemble a ‘half pipe’ with an abrupt upslope boundary akin to glacial ‘trim lines’ seen in modern glaciated valleys, with steeper fluvially-dissected slopes upslope. High elevation cirque basins testify to the importance of freeze-thaw processes marginal to perennial snow patches, glacierettes and cirque glaciers (e.g., Lian and Hicock, 2010) and the delivery of coarse debris downslope to valley side alluvial fans and ultimately large trunk rivers. This paper focusses exclusively on the core zone of the ice sheet across the comparatively low relief Intermontane Belt (Figs. 1 and 2) where the effects of subglacial abrasion are widely expressed

in the form of streamlined bedrock and/or sediment (Armstrong and Tipper, 1948, Figs. 8 and 9). Processing, manipulation and examination of the topographic and geomorphic data sets described above, combined with geological field work and assessment of glacial geologic mapping by National, Federal, State and Provincial geological surveys (Table 1), permits identification of three fundamental types of subglacial terrain (referred to here as ‘bed types’) based primarily on the relative areal extent of rock and till. These are; 1) streamlined ‘hard’ beds of exposed rock; 2) so-called ‘soft’ beds composed primarily of till, and 3) an intermediate (‘mixed’) bed type with areas of exposed rock protruding as knobs of various sizes through a till/sediment cover. Each of these is reviewed below.

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Fig. 7. Megascale glacial lineations and drumlins on sediment and rounded rock uplands of a ‘mixed bed’ west of Whitefish Lake on the bed of the Rocky Mountain Trench Ice Stream (#4; Fig. 20) which terminated as the Flathead Lobe (Fig. 3B and Fig. 4). Ice flow direction to south.

4.1. Subglacial terrain types 4.1.1. Hard bed subglacial terrain The geomorphology of subglacially-streamlined hard beds is
es, crag and tails, rock drumlins and dominated by roche moutonne by kilometre-long parallel ‘fluted’ bedrock ridges with a sparse or non-existent sediment cover commonly consisting of coarsegrained, often bouldery till (Figs. 6, 10 and 11). The origins of fluted ridges cut into bedrock were discussed at length by Krabbendam et al. (2016) and given the time required for their formation, may be long lived bedforms that evolve over several glacial cycles. Nonetheless, given their presence under modern ice streams (e.g., Livingstone et al., 2012, 2016) it is reasonable to

assume that they have the same paleoglaciological significance as megascale glacial lineations (MSGLs) found on soft till beds and were also cut below fast flowing (streaming) ice. Such rock-cut MSGLs are present across flood basalts of the Omak Plateau on the northern margins of the Columbia Plateau (Fig. 6) and on volcaniclastic strata exposed along the floors of major outlet valleys such as the Flathead, Nass and Skeena valleys (Figs. 6, 10 and Fig. 11A, B). These rock-cut landforms are argued to identify the trunks of fast flowing paleo ice streams that were able to abrade their underlying hard beds as a consequence of their high flow velocity (e.g., Krabbendam et al., 2016). The so-called ‘fluted bedrock ridges’ mapped by Stumpf et al. (2000) (MSGLs; Fig. 6A) along the floor of the Babine Lake Valley can now be recognized as

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Fig. 8. Methodological steps used in this study to produce and analyse DEMs for identification of the footprint of paleo ice streams within CIS identified by megascale glacial lineations.

being part of an extensive regional flow set of rock-cut megascale glacial lineations defining the bed of the newly identified Babine Ice Stream (#9: Fig. 10B). Across the Interior Plateau, Columbia Plateau and Omak Plateau subglacially-streamlined bedrock surfaces commonly show a thin (<3 m) discontinuous cover of pale-coloured, silt-rich ‘rubble till’ (Fig. 12D) which was likely a highly effective abrasional agent when being driven across underlying rock surfaces under fast flowing ice. Coarse-grained bouldery basal debris produced by subglacial quarrying and plucking from up-ice facing escarpments, was deposited as areas of boulder-rich hummocky end moraine resting on drumlinized and megascale glacially lineated rock; a conspicuous example is the large Withrow Moraine that marks the margin of the Okanogan Lobe on the Omak Plateau (Kovanen and Slaymaker, 2004, Fig. 12). These deposits are comparable to ‘rubble moraine’ on the glacially-streamlined limestone hard beds of central Canada, that formed by subglacial erosion of up-ice facing escarpments under fast flowing ice (e.g., Eyles and Doughty, 2016). Examples of rock cut MSGLs from the bed of the CIS add to a growing body of evidence showing their locally common presence on bedrock (Eyles, 2012; Eyles and Putkinen, 2014; Krabbendam et al., 2016). Directly analogous grooved hard beds are inferred from modern Antarctic ice streams (Livingstone et al., 2012, 2016) further supporting a genetic link with fast flow. Enhanced glacial abrasion and erosion is also expressed along the floors of Cordilleran valleys by narrow, glacially-overdeepened so-called ‘fiord lake’ basins (e.g., Adams, Okanagan, Shuswap, Pend Oreille, etc.) where water depths are considerable (in excess of several hundred metres) and sub-bottom geophysical data suggest underlying glaciolacustrine sediment fill thickness that can locally exceed 1 km (Eyles et al., 1990, 1991a, b; Mullins et al., 1990, 1991; Vanderburgh and Roberts, 1996). These inland basins in the core zone of the CIS are strikingly similar in their overall morphology to overdeepened fiords along the Pacific coastal margin (e.g., Shaw and Lintern, 2016) and they trapped very large volumes of subglacial and proglacial glaciolacustrine sediment during ice expansion (to be reworked into deformation tills that may have sustained fast ice flow as discussed below). By analogy with the modern Greenland and Antarctic ice sheets these large deep water bodies

may have survived as subglacial lakes at the time of maximum ice volume; indeed Livingstone et al. (2013) modelled and predicted a high potential for subglacial lakes below CIS, a reflection of the mountainous topography and relief of the ice sheet's bed. These water bodies may also have aided disintegration of the ice sheet during deglaciation when they were greatly enlarged and deepened still further by ice damming (e.g., Eyles and Clague, 1991; Lesemann and Brennand, 2009) (see Discussion below). 4.1.2. Mixed bed subglacial terrain As related above, the core zone of CIS is defined by upland plateau of the Intermontane Belt in what is now central British Columbia lying at elevations greater than 1200e1500 masl. This area shows large swaths of drumlinized till with protruding and variably streamlined bedrock giving rise to a geomorphologically distinct ‘mixed bed’ subglacial terrain. Drumlins are variably composed of rock (‘rock drumlins’), sediment preserved in the leeside or stoss-side of rock knobs (‘composite drumlins’) or are composed entirely of sediment (‘drift drumlins’). This mixed subglacial terrain type is particularly well developed around the community of Kamloops where it was mapped in detail by Fulton (1976a, b, c, d) (Figs. 13 and 14), east of Prince George (Fig. 15) in British Columbia, across the Omak Plateau of Washington State (Fig. 6) and parts of Idaho and Montana (Fig. 7). Bedrock drumlins across the Thompson Plateau were described by Lesemann and Brennand (2009) and McClenagan (2013). While recognizing that they are clearly erosional in origin they downplayed the effects of direct subglacial abrasion and emphasized instead, the erosional action of massive subglacial sheet floods (the ‘meltwater hypothesis’ of Shaw, 2002). This is not the place for a full discussion of that hypothesis except to note that the subglacial fluvial interpretation of McClenagan (2013) was rejected by Stumpf et al. (2014). While much water was undoubtedly stored as subglacial lakes under the Cordilleran Ice Sheet (e.g., Livingstone et al., 2013) a megaflood origin for drumlins is not considered further. ‘Ribbed moraine’ is another much less common type of subglacial landform identified in the present study and these are typically associated with mixed beds across the core zone of CIS. This landform type consists of closely-spaced, gently-rounded

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Table 1 Sources used in assessment of Quaternary geology and subglacial bed types across the Cordilleran Ice Sheet. Publication

Geographical Area

Alden, 1932 Alley and Chatwin, 1979 Arnold et al., 2016 Bednarski, 2015 Carrara, 1989 Carrara et al., 2007 Clague, 1976 Clague, 1983 Clague, 1986 Clague, 1987 Clague, 1988 Clague, 1991 Clague et al., 1982 Cui et al., 2015 Eyles, 1995 Eyles and Clague, 1991 Eyles and Kocsis, 1989 Fullerton et al., 2004 Fulton, 1965 Fulton, 1967 Fulton, 1971 Fulton, 1976a Fulton, 1976b Fulton, 1976c Fulton, 1976d Fulton et al., 2004 Geological Survey of Canada (2014) Herzer and Bornhold, 1982 Hughes et al., 1969 Johnsen and Brennand, 2004 Levson and Giles, 1995 Levson et al., 1998 Lipovsky and Bond, 2014 Locke and Smith, 2004 Margold et al., 2013a Margold et al., 2013b Muller, 1967 Muller and Christie, 1966 Perkins and Brennand, 2015 Plouffe, 1991 Plouffe, 1992 Plouffe, 1994a Plouffe, 1994b Plouffe, 1995 Plouffe, 1996a Plouffe, 1996b Plouffe, 1996c Plouffe, 1996d Plouffe, 1997 Prest et al., 1968 Richmond, 1986 Spooner and Osborn, 2000 Thorson, 1980 Tipper, 1971a Tipper, 1971b Tipper, 1994

Eastern Montana Southwestern Vancouver Island British Columbia and Yukon Northeastern British Columbia Glacier National Park Region, Montana Southeastern Alaska Strait of Georgia, British Columbia Skeena River, British Columbia British Columbia Williams Lake, British Columbia Quesnel, British Columbia Cariboo Region, British Columbia Vancouver Island, British Columbia British Columbia Cariboo Region, British Columbia Central British Columbia Cariboo Region, British Columbia Northern Montana Southern British Columbia Kamloops Region, British Columbia Southern British Columbia Shuswap Lake, British Columbia Vernon, British Columbia Kamloops Lake, British Columbia Merrit, British Columbia British Columbia Canada Southwestern Vancouver Island Yukon Thompson Basin, British Columbia Nechako Plateau, British Columbia Smithers and Hazelton, British Columbia Yukon Montana Northern British Columbia and southern Yukon Central British Columbia Kluane Lake, Yukon Kluane Lake, Yukon Fraser Plateau, British Columbia Northern interior British Columbia Central British Columbia Chuchi Lake, British Columbia Tezzeron Lake, British Columbia Manson River and Fort Fraser, British Columbia Cunningham Lake, British Columbia Burns Lake, British Columbia Fraser Lake, British Columbia Tsayta Lake, British Columbia Central British Columbia Canada Northern Rocky Mountains Stikine River Valley, British Columbia Puget lowland, Washington Central British Columbia Central British Columbia Central British Columbia

morainal ridges oriented transverse to ice flow direction (Fig. 21). These are identified as Rogen moraines and available outcrops indicate that they are composed in the main, not of till, but coarsegrained, very poorly-sorted ice proximal outwash sands and gravels that are commonly extensively glaciotectonized. Rogen moraines are commonly transitional downflow to drumlins recording modification of the original transverse bedform as ice velocity in€ ller, 2006 for creases and begins to streamline its bed (e.g., Mo Scandinavian examples). The distribution of Rogen moraines across CIS was not systematically assessed by the present study but available data indicate in general, that they are typically associated with former ice divides where ice flow was sluggish and the ice € ller, sheet was possibly cold-based (e.g., Dunlop and Clark, 2006; Mo

2006; Trommelen et al., 2014). 4.1.3. Soft bed subglacial terrain Soft bed subglacial terrain encompasses extensive areas of wellstreamlined subglacial surfaces at elevations generally below 15001200 masl that are underlain by thick (up to 20 m) and extensive deposits of matrix-rich Late Wisconsin till. In central and southern British Columbia this till has been mapped as ‘Fraser Till’ (Fulton et al., 2004 and refs therein) (Fig. 9) with locally-named stratigraphic equivalents in the US (e.g., Vashon Till; see Easterbrook, 2003). These areas of soft beds are extensively drumlinized (Armstrong and Tipper, 1948, Fig. 9B) but very significantly also show contain well defined flow sets of megascale glacial lineations

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Fig. 9. Simplified regional extent of ‘soft bed’ terrain across the core zone of the Cordilleran Ice Sheet divided into drumlinized areas (A) and those areas of much more elongated MSGLs (B) that identify trunks of paleo ice streams (Fig. 20); note presence of much smaller areas of MSGLs identifying complex local streaming as part of a possible onset zone (see Fig. 17 and text).

(Fig. 5, Fig. 9C and Fig. 16e20) that are the footprint of fast flowing paleo ice streams (e.g., Barchyn et al., 2016; Briner, 2007; Clark, 1993; Clark and Stokes, 2013; Ely et al., 2016a; Evans et al., 2008, 2014; Eyles and Doughty, 2016; Eyles et al., 2016; Jamieson et al., 2016; King et al., 2009; Livingstone et al., 2012; MacLean et al., 2017; Ross et al., 2009; Spagnolo et al., 2016, 2017; Stokes and Clark, 2002; Stokes et al., 2013) (Fig. 20). Eyles et al. (2016) and Ely et al. (2016a) recently argued that drumlins and MSGLs are part of a geomorphic continuum but as related above, they are differentiated in the present study for the specific purpose of separating the spatial footprint of the trunks of paleo ice streams from the drumlinized beds of their catchment areas upglacier. Those areas downglacier that show transitions from drumlins to MSGLs are interpreted as paleo ‘onset zones’ where ice was beginning to accelerate (e.g., Barchyn et al., 2016). 5. Paleo ice streams within the Cordilleran Ice Sheet The result of ice sheet-wide mapping of MSGLs from digital data sets identifies a number of paleo ice streams within CIS (Fig. 20). As emphasized above, this paper does not attempt to map the beds of paleo ice streams in detail nor identify the full suite of landforms associated with each. A detailed case-by-case assessment using the full range of geomorphic criteria identified by Stokes (2011) and

Clark and Stokes (2013) based on a detailed comparison with modern ice stream types found today in Greenland and Antarctica, is now the subject of ongoing work. Most paleo ice streams so far identified within CIS can be described as ‘terrestrially terminating’ and some are bounded by large arcuate end moraine complexes (e.g., Okanogan Ice Stream # 3; Two Medicine Ice Stream # 14; Fig. 20); most others lack welldefined downglacier limits such as the case of the Georgia-Puget Ice Stream which instead, terminated as extensive outwash deposits (Easterbrook, 2003). In this case, any end moraines appear to have been destroyed, or were prevented from forming, by topographic confinement of the former ice stream and associated reworking by meltwater or marine waters during rapid deglaciation. As will be discussed below, many ice streams likely were present within coastal troughs along the Pacific margin; and their megalineated beds are obscured by younger marine sediments. Consequently, a large degree of preservational bias is clearly evident from consideration of the spatial distribution of paleo ice streams on Fig. 20. The location of paleo ice streams in CIS (Fig. 20) appears to reflect severe topographic forcing of ice flowing out from the ice sheet's inner core zone and being forced to escape through lowland gateways and narrow outlet valleys; very high ice velocities enhanced by topographic constriction likely explains the presence

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Fig. 10. Megascale glacially-lineated hard bed cut by the Nass Ice Stream (A) (#7; Fig. 20). Along the Nass River Valley (Fig. 3B). Note sharp lateral boundary of streamlined rock against fluvially-dissected valley sides. B: Area of glacially-streamlined hard bed on rock cut below Babine Ice Stream (#9; Fig. 20) near Babine Lake; ice flow to bottom right (see Fig. 3B). Note prominent rock drumlin and MSGLs (outlined).

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Fig. 11. Flow set of rock-cut megascale glacial lineations cut by Skeena Ice Stream (A and B) (#8; Fig. 20) on south flank of Hazelton Mountain, Skeena River Valley (Fig. 3B). Flow set is approximately 1.4 km wide. MSGLs likely record accelerated ice flow over the high standing Hazelton Mountain (55 160 2000 N, 127 500 2700 W) which formed an obstacle on the bed of the ice stream.

of flow sets of MSGLs cut into rock along several outlet valleys (e.g., Nass #7, Skeena #8) (Fig. 20). Again, this is subject to the proviso that such bedforms may reflect erosion during multiple glacial cycles. The largest topographically-controlled ice stream in CIS flowed south from numerous tributary valleys in the Coast Ranges and on Vancouver Island into Strait of Georgia-Puget Sound (#1) (Fig. 20); the longest ice stream flowed along the Rocky Mountain Trench where it is possible that the Nechako Ice Stream (# 13: Fig. 20) may have been the uppermost part of an extended and narrow ice stream (Rocky Mountain Trench Ice Stream # 4) (Fig. 20) that terminated as the large Flathead Lobe (Fig. 4). The smallest ice stream so far identified in CIS occurs immediately east of the continental divide in northern Montana, where extensive flow sets of MSGLs are exceptionally well preserved on the bed of the former Two Medicine piedmont lobe (#14: Figs. 20 and 24) which was fed

by Cordilleran ice flowing out of the Rocky Mountains through the relatively low lying Marias Pass in Glacier National Park (Alden, 1932; Carrara, 1989; Locke and Smith, 2004; Fullerton et al., 2004, Fig. 25). This composite ice stream (it was made up of several confluent ice flow units) was approximately 50 km in width and at least 55 km long (similar to today's Malaspina Glacier) and unusually, left a series of well-preserved hummocky till-cored moraines during its decay. Because of the semi-arid climate and lack of any vegetative cover, this is the best preserved bed of a former Cordilleran ice stream so far identified in the present study. Initial investigation, indicates the importance of glacier flow down mountain front pediment surfaces cut across soft shales rich in swelling clays. Along the eastern margin of CIS in southern Alberta near Calgary, there is clear geomorphic evidence on the floor and shoulders of the Bow River Valley for fast flowing ice moving east out of the

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Fig. 12. Dissected drift drumlin (A) north of Withrow in Washington State on the bed of the Okanogan Ice Stream (#3; Fig. 20) (47 420 1200 N, 119 320 6000 W). B: Ground view of streamlined landforms shown in A (arrowed) with MSGLs in foreground. C: Grooved glacially-abraded basalt surface exposed in cross-section in road cut. D: Typical pale-colored, silt-rich rubble till resting on bedrock surface shown in C. E: Withrow Moraine with very large glacially-quarried basalt boulders. See Fig. 4 for locations. (1.5 column size).

Rocky Mountains (Fig. 25). The valley contains prominent drumlins (which Fisher and Spooner, 1994 attributed to subglacial meltwater floods) together with newly identified transitions from ‘cloned’ drumlins to MSGLs (this study). This relatively small Cordilleran paleo ice stream is tentatively referred to as the Bow River Ice Stream (# 15; Fig. 20). The regional trend of MSGLs shows that it was deflected south by convergence with the Laurentide Ice Sheet margin located on the plains immediately east of Calgary. Ongoing work suggests it was part of a much larger composite ‘Foothills Ice Stream complex’ composed of other fast flowing ice units moving east out of the Rocky Mountains that were forced south by the Laurentide Ice Sheet, carrying large quartzite erratic blocks of the Foothills Erratics Train (e.g., Jackson et al., 1997). Mapping and demarcating the confluence zone of the Cordilleran and Laurentide ice sheets in Alberta is traditionally reliant largely on stratigraphic assessment, dating and mapping of the geographic limits of montane and Laurentide till sheets dominated by Rocky Mountainsourced or Shield-derived clasts respectively (e.g., Holme et al., 2000). Going forward, the availability of high resolution digital imagery now allows direct geomorphic mapping of the margins of the two ice sheets along the eastern margin of the Rocky Mountains and western plains. It can be observed that MSGLs (and thus paleo ice streams) are uncommon in areas of mixed bed (with the notable exception of the Rocky Mountain Trench Ice Stream; #4; Fig. 20) underscoring the possible limiting role on ice velocity created by emergent, highstanding bedrock knobs protruding through the till bed (e.g.,

Fig. 14A) and suggesting the importance of bed roughness as a brake on ice velocity. Drumlinized mixed beds are more typical of the extensive catchment areas upglacier that fed individual ice streams at lower elevations where ice was forced through valleys and lowlands. Ice stream ‘onset zones’ are identified by downglacier transitions from drumlins to MSGLs. On this basis, complex spatial transitions from sluggish to streaming flow appear to be the norm (e.g., Fig. 5B, Fig. 16 and 17) possibly reflecting changes in basal ice temperature, bed rheology and other factors (e.g., Stokes et al., 2007) that await detailed assessment. It is intriguing to note that the prominent flow set of MSGLs identifying the bed of the Kluane Ice Stream near Kluane Lake (Fig. 19) occurs in the same area where many modern valley outlet glaciers in the nearby Kluane Icefield are surge-prone (see Bevington and Coupland, 2014 and refs therein). Detailed examination of the geology and geomorphology of the bed of the Kluane Ice Stream could shed light on conditions that may have triggered fast ice flow, and thus possible the nature of geological controls that contribute to modern surge behaviour in the area. Elsewhere across the core zone of CIS, small (and unnamed) patches of MSGLs lying outside the much longer trunk areas of the major ice streams (Fig. 20) are not as yet well understood and could reflect local streaming of otherwise sluggish moving ice forced to flow around, over or between emergent bedrock highs; again, more detailed field assessment is required. The Cordilleran Ice Sheet has been recently described as the least understood of all northern hemispheric Pleistocene ice sheets; the results of the present study show that the drainage of the ice

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Fig. 13. Large tract of mixed bed on plateau basalts of the Interior Plateau north of Kamloops in central British Columbia. A: Drumlinized Fraser Till. B: Rock-cored drumlins with streamlined rock (C), and non-streamlined bedrock highs (D) protruding through the till cover. Ice flow to south (Fig. 3B for location).

sheet was fundamentally no different than that of modern (e.g., Greenland and Antarctica) and ancient counterparts (LIS) in having numerous fast flowing ice streams. The total area of ice stream trunks identified by this study (Fig. 20) undoubtedly represents a small fraction of the total ice sheet footprint (Fig. 9) (see also Stokes et al., 2016 in reference to the LIS) but the present inventory likely only identifies just a small fraction of the total number of paleo ice streams primarily because of a lack of appropriate offshore data from deep troughs along the 1000 km-long fiord-indented Pacific coastal margin of British Columbia and coterminous parts of southeast Alaska. Submarine troughs such as Dixon Entrance and the adjacent Skeena Valley, Moresby Trough, Mitchell's Trough and Goose Island Trough (Fig. 20) were likely occupied by ice streams from the Coast Ranges and interior that were funneled through narrow outlet valleys and fiords (e.g., Herzer and Bornhold, 1982). Seafloor imagery of subglacial bedforms along the deep axis of these troughs are unavailable. The work of Swartz et al. (2015), Ely et al. (2016b) and Montelli et al. (2017) on the offshore geomorphic

footprint of marine ice streams along the far northwestern part of the CIS on the Gulf of Alaska and Alaskan Peninsula is a guide and predictor of what may be present offshore to the south. In retrospect, the very considerable thickness (5 km) of the glaciomarine infill of the forearc basin along the Yakataga section of the Gulf of Alaska east of the Copper River (e.g., Eyles, 1988; Eyles and Eyles, 1989; Eyles et al., 1991a) may reflect not only rapid tectonic uplift and rapid glacial erosion resulting from extreme precipitation values along the mountainous coast but also the enhanced delivery of debris offshore by fast-flowing paleo ice streams flowing through the submarine troughs that cross the continental shelf (e.g., Montelli et al., 2017). It is speculated that most of the Pacific fiords and offshore troughs in British Columbia and Southern Alaska hosted marinebased paleo ice streams (e.g., Hicock and Fuller, 1995). Indeed, repeated episodes of fast ice flow and enhanced subglacial erosion during successive glacial cycles may have contributed to long term overdeepening of fiords (e.g., Kessler et al., 2007). The same model

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Fig. 14. Representative example of typical mixed bed (A) occurring south of Kamloops consisting of rounded streamlined basalt highs protruding through drumlinized glacial cover Reproduced with permission of Geological Survey of Canada from Fulton (1976a, b, c, d). B and C: Drumlins composed of Fraser Till, and high standing streamlined rock knobs (Fig. 3B for location).

could apply inland to the glacially-overdeepened interior bedrock valleys of central British Columbia whose floors extend well below sea level and which are now occupied by deep fiord lakes (e.g., Okanagan, Shuswap) and the deep lake basins in the US most notably the Lake Pend Oreille basin in Idaho. 6. Geologic controls on ice streams in the Cordilleran Ice Sheet This section briefly outlines the principal geological factors that appear to have controlled the location of paleo ice streams within

the Cordilleran Ice Sheet emphasizing the importance of a soft sediment bed and topographic confinement as ice escaped from the core zone through the surrounding mountain rim. It can be noted that MSGLs are largely confined to soft beds where paleo ice streams flowed over matrix-rich till (e.g., Georgia-Puget Ice Stream #1, Fraser #5, Nechako # 13, Two Medicine #14) (Fig. 20). In these areas till commonly rests unconformably on underlying glaciotectonized sediment and was clearly ‘manufactured’ (Hershey, 1897) by subglacial cannibalisation and mixing of older sediments which were homogenized to varying degrees and re-transported subglacially. Very thick ‘sub-till sediments’ are widely distributed

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Fig. 14. (continued).

and exposed across the Intermontane Belt lowlands and plateau that comprise the core zone of CIS where deeply-cut Cenozoic paleovalleys and depressions are filled with great thicknesses (<1 km) of fine glaciolacustrine sediments (e.g., Andrews et al., 2011, 2012; Clague, 1986; Eyles and Clague, 1991; Eyles et al., 1987, 1990; Hickin et al., 2016; Lian and Hicock, 2001; Vanderburgh and Roberts, 1996; Nichol et al., 2015; Plouffe et al., 2011; Spooner and Osborn, 2000; Stumpf et al., 2004). In these areas, drumlinized and megalineated till commonly rests unconformably on fine-grained sediments deposited in proglacial lakes during the earlier growth phase(s) of the ice sheet (Eyles and Clague, 1991; Eyles et al., 1987; e.g., Fig. 22) and in addition perhaps, in subglacial waterbodies at times of maximum ice sheet volume (see Livingstone et al., 2013). Soft bed till is commonly massive or crudely bedded with lithic clasts floating in a fine-grained matrix; these facies typically occur as superposed, sub-horizontal, undulatory beds of variable thickness (<1 m to several metres) composed of reworked and partiallyto-completely homogenized glaciolacustrine sediment (Fig. 22)

(see also Hickin et al., 2016). Clasts and rafts of intact partiallyreworked pre-existing sediment are very common and diagnostic. In cases where drumlin and MSGL cores are exposed it is apparent that the bedform truncates underlying sediments (Fig. 22). The same observations apply to drumlins and MSGLs of the Puget Lowlands; their cores are very well-exposed in lengthy coastal bluffs and show Vashon Till that rests erosively on thick proglacial subaqueous sediments deposited during ice expansion (see Easterbrook, 2003; Easterbrook et al., 2003). The paleoglaciological significance of these older sediment fills is that they are finegrained and easily reworked subglacially; it seems very reasonable to argue that they allowed the generation of an extensive soft till bed that would eventually promote and sustain ice streaming. The role of ice marginal water bodies and their antecedent sediments in lubricating fast ice flow was also stressed by Stokes and Clark (2004) in their study of the former Dubawnt Lake Ice Stream which flowed on hard crystalline rocks of the Canadian Shield; it is suggested that the same model likely has wide application across the bed of the Cordilleran Ice Sheet.

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Fig. 15. Mixed bed consisting of rock drumlins cut into volcaniclastic strata of the Nicola Group (outlined) and drift drumlins cut into Fraser Till east of Prince George (Fig. 3B).

6.1. Drumlins and megascale glacial lineations as a bedform continuum Drumlins are an iconic subglacial landform and many tens of thousands of sediment-cored (‘drift’) drumlins' and rock-cored (‘composite’) drumlins occur across the mixed and soft beds that formed under the core zone of CIS (Figs. 9, 15e17, 23). This is not the place for a lengthy discussion of drumlin forming processes but some brief comments are warranted. Field evidence from the Cordillera suggests that they are predominantly erosional in origin given 1) they are associated with rock-cut counterparts (rock drumlins), and 2) observations of available outcrops (admittedly small in number compared to the total population of bedforms) shows their surface form commonly truncates underlying deposits such as till and/or associated glaciolacustrine sediment. As related above, these bedforms are likely associated with relatively sluggish ice flow velocities however precisely defined, and form a spatial continuum with MSGLs (Fig. 9C and Fig. 20). While some models see such bedforms as the product of the upwards growth of periodic Rayleigh-Taylor ‘instabilities’ at the interface of ice resting on a soft bed (see Spagnolo et al., 2017) we have previously suggested that this spatial relationship is more consistent with a common

erosional origin for drumlins and MSGLs involving the progressive dissection of relatively large ‘parent’ drumlins as ice velocity increased to streaming flow. In this model drumlins are incised longitudinally (i.e., dissected) by flow-parallel bands of deforming subglacial debris (the ‘erodent layer’ of Eyles et al., 2016) to produce longer, narrower daughter ridges (MSGLs). This process when repeated, lowers the bed to ultimately create successively narrower and longer bedforms; a process specifically termed ‘cloning’ by Clark (2010, p. 1022) and ultimately producing a bed of markedly lower relief allowing ice to flow faster. The production of MSGLs by some form of erosional cloning process from drumlins is evident across large areas of the core zone and periphery of CIS (Figs. 5, 16 and 17) and is amply supported by prior work in British Columbia. Armstrong and Tipper (1948, p. 289e90) reported the common presence of what they called ‘grooved drumlins’ with prominent longitudinal scours (Fig. 23) which they attributed to scoring by large boulders held in basal ice moving over the drumlin bedform. They pointed to the common presence of large erratic boulders (up to 5 m in diameter) lying on drumlins. Essentially the large longitudinal grooves are ‘megastriations’ or ‘megagrooves’ no different from those formed on rock surfaces and indicative of selective linear subglacial erosion by a

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Fig. 16. Complex areal transitions from drumlins to MSGLs on the bed of the Fraser Ice Stream (#5; Fig. 20) underlain by Fraser Till and older glaciolacustrine sediments. Ice flow bottom left to top right (Fig. 3B for location). Note large numbers of dissected drumlins (X) (see Fig. 24) suggesting the area lay close to the onset zone of the ice stream.

relatively hard ‘indenter’ whether a large single boulder, clast clusters or patches of stiffer deforming debris being swept downglacier across the drumlin (see Eyles and Doughty, 2016; Krabbendam et al., 2016). Downglacier movement of deforming

subglacial debris is inferred to be an active process at the base of modern ice streams in Antarctica (e.g., King et al., 2009; Smith et al., 2012) and was defined as ‘excavational deformation’ by Hart (1995). It can be suggested that the ‘grooved drumlins’ of

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Fig. 17. Area of MSGLs surrounded by drumlins northeast of Prince George (Fig. 3B) recording northeast ice flow. Note pervasive dissection of drumlins into MSGLs at X named as ‘grooved drumlins’ by Armstrong and Tipper (Armstrong and Tipper, 1948; Fig. 23) indicating the area was the onset zone of a Nechako Ice Stream (#13, Fig. 20) flowing north into the Rocky Mountain Trench.

Armstrong and Tipper (1948) record the initial dissection of large parent drumlins (Fig. 23) by subglacial till containing very large clasts. The same authors noted numerous drumlins transitional downglacier to what they termed ‘poorly-defined parallel groovings’ (the MSGLs of this study). Their statements (italicized herein) that ‘drumlins are most common where ice was flowing upslope’ (i.e., is interpreted as moving more sluggishly) and ‘groovings where ice was flowing downslope’ (i.e., where ice was accelerating) may be simplistic in a strictly glaciological sense but are prescient in the light of modern work suggesting a fundamental relationship

between bedform elongation and ice velocity. Further, they also considered that ‘parallel groovings’ developed where ‘the force of the ice was sufficient to destroy the projecting till’ (p.298); this is consistent with progressive lowering and cloning of a relatively high relief initially drumlinized till surface by a thin (<1 m) ‘erodent layer’ of deforming subglacial debris. In light of the above observations, drumlins on the bed of the Cordilleran Ice Sheet are a record of relatively sluggish ice flow velocities; ‘grooved’ and ‘cloned’ drumlins appear to be typical of ice stream ‘onset zones’ where ice was accelerating to streaming velocities and are transitional

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Fig. 18. Megascale glacial lineations on the bed of the Anahim Ice Stream masked by thin glaciolacustrine sediment on Fraser Till (see Fig. 3B for location).

bedforms to highly elongated MSGLs defining the trunks of fully developed, fast-flowing paleo ice streams.

7. Timing of ice streaming and possible relationship to abrupt deglaciation of the Cordilleran Ice Sheet Recent glaciological modelling stresses that the timing of the maximum extent and subsequent decay of CIS was markedly outof-phase with regional hemispheric paleoclimate records suggesting the influence of additional, as yet unknown topographic and glaciological controls (Seguinot et al., 2016). In short, the ice sheet appears to have reached its maximum volume relatively late and disappeared very quickly immediately thereafter. Given this history, it seems reasonable to propose that the paleo ice streams identified on Fig. 20 likely developed only during the very latest

phase(s) of ice sheet growth when it reached its maximum volume. Indeed, this notion is supported by earlier work elsewhere. Dowdeswell and Elverhøi (2002) argued that Late Weichselian (Wisconsin) ice streaming of the Barents Sea Ice Sheet along the Svalbard continental margin in northern Europe, was triggered just when the ice sheet reached its maximum volume and outflowing ice drainage began to be strongly topographically controlled by deep coastal troughs resulting in accelerating flow velocities. In British Columbia, ice accumulating across the broad core zone of the Intermontane Belt was subject to the same type of topographic forcing as suggested for the Barents Sea Ice Sheet when Cordilleran ice was funneled through narrow outlet valleys in the mountains that surround the core zone to the west and east. Ice flowed westward through narrow fiords into offshore troughs and marine straits (e.g., Puget Sound); that ice flowing eastward was also forced

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Fig. 19. Present day ice cover on the north east flanks of the St. Elias Ranges adjacent to Kluane Lake in the Yukon Territory (A). Many glaciers have a surge-prone history. Inset (B) megascale glacial lineations on sediment left by the Kluane Ice Stream (#12; Fig. 20) near Burwash Landing on Kluane Lake (see Fig. 3B for location).

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Fig. 20. Paleo ice streams within CIS identified by presence of megascale glacial lineations on hard and soft beds and named after principal parent valley, basin or community. Ice streams originate within catchments across large areas of drumlinized terrain in the Intermontane Belt and their location primarily reflects topographic funneling of outflows through narrow outlet valleys and lowlands. The Fraser Ice Stream (#5) may have been the uppermost part of an extended Rocky Mountain Trench Ice Stream (#4) as much as 1000 km in length terminating as the Flathead Lobe in the US (Fig. 4). Possible locations of additional inferred paleo ice streams draining to Pacific Ocean through coastal fiords and offshore troughs are shown in red. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

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Fig. 21. Area of rib-like Rogen moraine ridges (A) formed transverse to ice flow under sluggish or stagnant ice below the former Williams Lake ice divide (R on Fig. 3A). B: Glaciotectonized ice-proximal outwash sands and gravels within core of Rogen moraine.

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Fig. 22. Fraser Till (base arrowed) (A) resting unconformably across thick glaciolacustrine fill along Nicola Valley south of Spences Bridge. Till is derived from glacio-tectonism and subglacial reworking of underlying sediment. B: Exposure through drumlin showing truncation of glaciotectonized glaciolacustrine sediment by drumlin surface which is clear evidence of an erosional origin. Note absence of surface till. C: Massive Fraser Till facies composed of subglacially-mixed glaciolacustrine silt and clay matrix and generally subrounded clasts. D: Heterogeneous Fraser Till facies produced by incomplete subglacial mixing of glaciolacustrine silty clays and outwash gravels, overrun during ice sheet expansion.

through deep elongate rock basins some of which are occupied by ‘fiord lakes’ that today are hundreds of metres deep but whose floors are underlain by thick lateglacial sediment fills indicating much greater water depths during maximum ice volume and deglaciation. Consequently, the model of Dowdeswell and Elverhøi (2002) involving the inset of ice streaming as a consequence of topographic forcing at time of maximum ice volume is very attractive in inferring the possible onset of fast ice flow within CIS. If this hypothesis is correct then the entire population of paleo ice streams is of the same age since they must have been very shortlived; deglaciation occurred rapidly after maximum ice sheet volume was attained. This suggests further, that rapid deglaciation was causally related to the onset of ice streaming. Fulton et al. (2004) suggested that CIS essentially stagnated in situ; the absence of large recessional end moraines other than those of the lower Fraser Valley of southern British Columbia; the socalled ‘Sumas’ advances (see Clague et al., 1997; Easterbrook, 2003), together with fields of possibly annual moraine ridges on part of the Fraser Plateau (e.g., Eyles and Clague, 1991) and also the end moraines associated with the Two Medicine Ice Stream in Montana, support the absence of major still stands during ice sheet deglaciation. Rapid ice loss has been traditionally attributed to regional climate warming and the associated abrupt rise in equilibrium line elevation resulting in a strongly negative mass balance cutting off ice flow and leaving ice to stagnate over large areas (e.g., Booth, 1987; Fulton, 1991; Fulton et al., 2004; Guilbault et al., 2003; Margold et al., 2013a, b; Seguinot et al., 2016; Robert, 1991). This simple downwasting model is inconsistent with both the timing of deglaciation which is unrelated to climate (Seguinot et al., 2016) and also its rapidity where a very large volume of ice was lost in a few thousand years. Menounos et al. (2017) have more recently suggested a much more complex pattern of deglaciation possibly beginning as early as 14.5 ka when CIS may have lost some 50% of

its mass in as little as 500 years in response to hemispheric climate warming. Recognition of fast flowing ice streams within CIS now permits consideration of other glaciological processes that may have initiated rapid deglaciation such as ice stream collapse where they terminated in water. This model is amply supported by geologic data. Drainage of the southernmost Pacific margin of the ice sheet was dominated by the Georgia-Puget Ice Stream (Fig. 20) which is the largest ice stream artery so far recognized in the ice sheet. This ice mass is known to have retreated well north of Seattle by 13.6 ka as a consequence of marine waters flooding into Juan de Fuca Strait and Puget Sound (Booth, 1987) resulting in what has been called ‘rapid disintegration’ of the Puget Lobe (Easterbrook et al., 2003, p. 22). The importance of glacio-isostatic crustal depression and high relative sea levels along the coastal margin of the CIS can also be stressed as a possible factor in rapid ice retreat (e.g., Bassis et al., 2017). By 13.0 ka, ice had completely evacuated the Strait of Georgia to the north between Vancouver Island and the mainland (Clague, 1980; Clague and James, 2002; Dyke et al., 2003; Guilbault et al., 2003) pointing to collapse of the entire paleo ice stream in the Strait of Georgia. The same story of rapid ice retreat and possible ice stream collapse is apparent along the northern Pacific coastal margin of CIS where ice disappeared very early from Dixon Entrance (the site of a very large postulated ice stream) (Fig. 20) by 13,500 to 13,000 14C yr BP (Barrie and Conway, 1999). Clague (1985) inferred rapid deglaciation of Hecate Strait (another confluence area of several inferred paleo ice streams) by c. 12.7 ka and suggested ‘wholesale destabilization of the western periphery of the Cordilleran ice sheet by eustatically rising seas’ (p.256). The susceptibility of CIS to massive ice loss by calving along its Pacific sector has been emphasized by results of numerical modelling (Hendy and Cosma, 2008). Continuing rapid ice loss inland between 11 and 10 ka (Blaise

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Fig. 23. Schematic representation (A) of a ‘grooved drumlin’ (redrawn after Armstrong and Tipper, 1948). B: Grooved drumlin south of Kamloops representing the subdivision (cloning) of a large parent bedform into more elongate forms accompanying an increase in ice flow velocity. The process if continued may result in the formation of MSGLs.

et al., 1990; Carrara, 1995; Fulton et al., 2004) as outlet glaciers evacuated the Pacific coast is recorded by an abrupt spike in the volume of meltwater and sediments being moved out from the core zone of CIS resulting from the abrupt drainage of deep ice-dammed lakes (Burke et al., 2012; Eyles et al., 1991a, b; Johnsen and Brennand, 2004, 2006; Lesemann and Brennand, 2009) and captured in the sediment record of offshore basins along the Pacific margin (Blais-Stevens et al., 2003). Large interior lakes may have influenced rapid collapse of narrow outlet glaciers in the same way as did the glacio-isostatically depressed coastal margin troughs on the Pacific coast. Complex changes in ice flow direction (and thus ice sheet dynamics) that occurred as CIS waned (e.g., Arnold et al., 2016) may be the geomorphic record of abrupt ‘flow switching’ where highly dynamic unstable ice streams that are retreating, abruptly surge forward in different directions (e.g., Greenwood
Cofaigh et al., 2010; Winsborrow et al., 2012). et al., 2012; O This paper offers no data that might throw light on why CIS reached its maximum volume much later (14.5 ka) than the much larger Laurentide Ice Sheet at c. 18e20.0 ka. To some degree this diachroneity, if real (see Menounos et al., 2017), is the expected outcome of their vastly different sizes and volumes and especially their very different locations relative to sources of moisture. This

aspect was discussed by Dyke et al. (2002) who suggested the relatively late development of the Cordilleran Ice Sheet allowed the growing Laurentide Ice Sheet access to Pacific moisture. Hebbeln et al. (1994) showed that the Barents Sea Ice Sheet formed much faster and was out of phase with other ice sheets because of its marine setting and it is highly likely that the same factor was important in the life cycle of the Cordilleran Ice Sheet given its mountainous setting and proximity to Pacific Ocean moisture sources. A simple working hypothesis arising out of the above discussion, is that the onset of ice streaming in the Cordilleran Ice Sheet occurred as a result of topographic forcing brought about when the ice sheet attained its maximum volume (e.g., Dowdeswell and Elverhøi, 2002). Fast ice flow was sustained by the widespread presence of a soft till bed reworked from antecedent glaciolacustrine sediment trapped in topographic depressions and paleovalleys under its core zone inland. Rapid iceberg calving from marine terminating ice streams along its glacio-isostatically depressed Pacific margin could have triggered the abrupt collapse of the entire ice sheet in response to unsustainable ice losses from collapsing ice streams (e.g., Bassis et al., 2017). Additional significant catastrophic ice loss may have occurred across the interior of

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Fig. 24. Megascale glacial lineations along the bed of the former Two Medicine Ice Stream which flowed some 55 km east of the Rocky Mountain Front onto the interior plains (Fig. 4).

the core zone in response to rapid calving of outlet glaciers terminating in, or underlain by, deep (1 km) ‘fiord’ lakes and other subglacial water bodies (see Pattyn, 2003; Burke et al., 2012; Livingstone et al., 2013).

sediment (e.g., Fig. 9) could provide a well-exposed analog for assessing subglacial conditions under the Greenland Ice Sheet and their relationship to fast ice flow. 9. Conclusions

8. Comparison of the Cordilleran Ice Sheet with the modern day Greenland ice sheet There are sufficient glaciological similarities to suggest that the now exposed bed of CIS provides an appropriate analog for the subglacial geology and bed types currently obtaining below the Greenland Ice Sheet (Fig. 26). CIS is 50% larger but was of similar overall basal topography to that of the Greenland Ice Sheet which also has a ‘core zone’ of much lower subglacial relief amplitude and elevation surrounded by high mountains through which ice streams flow to the coast (Bamber et al., 2013a; Herzfeld et al., 2012; Rippin, 2013). The reduced subglacial relief of the core zone of CIS is in part, the product of magmatic loading and isostatic depression below extensive voluminous Neogene plateau basalts of the Intermontane Belt whereas the geology of the core zone in Greenland is composed of Precambrian basement and its low elevation (below sea level) is primarily the result of glacio-isostatic depression but both include deeply-incised subglacial fiord-like and canyon-like fluvial valleys in their interiors (Bamber et al., 2013b; Eyles et al., 1991a, b). The existence of thick subglacial sediments across the depressed core zone in Greenland has also been inferred from geophysical data (e.g., Rippin, 2013) and argued to be a key factor in sustaining fast ice flow within parts of the ice sheet (Walter et al., 2014). In conclusion, the extensively streamlined soft-bedded core zone of the Cordilleran Ice Sheet, underlain by a soft bed of till formed by subglacial reworking of older

After more than a century of detailed study of its complex geologic and stratigraphic record, the paleoglaciology and drainage of the Late Wisconsin Cordilleran Ice Sheet has remained obscure; the ice sheet has been said to be the least understood of all northern hemispheric Pleistocene ice sheets. High resolution mapping of the geomorphology of its bed using digital data imagery from western Canada and the northern United States, shows that its glaciological structure and drainage was no different than modern ice sheets such as in Greenland and Antarctica, and also its Late Wisconsin North American counterpart (Laurentide Ice Sheet). The Cordilleran Ice Sheet contained paleo ice streams now recorded by megascale glacial lineations (MSGLs) cut into a soft till bed, and locally bedrock. The onset zones of most Cordilleran ice streams were located across the ice sheet's core zone which was largely restricted to extensive plateau and lowland surfaces of the northsouth elongated Intermontane Belt of British Columbia. This zone shows transitions from drumlins to megascale glacial lineations cut into a thick soft bed of till, and which were also cut into rock along the floors of confined outlet valleys supporting the hypothesis of an erosional origin for both drumlins and megascale glacial lineations regardless of substrate, and confirming that these bedforms are end members of a subglacial ‘bedform continuum’ that reflects increasing ice flow velocities and the reduction of bed friction arising from progressive dissection of large drumlins into MSGLs. It is speculated that the onset of fast ice flow within the

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Fig. 25. MSGLs of the Bow River Ice Stream west of Calgary, Alberta (Fig. 20).

Cordilleran Ice Sheet was triggered by topographic funneling along outlet valleys draining the core zone of the ice sheet across the Intermontane Belt when the ice sheet expanded to its maximum volume. Fast ice flow was likely sustained by an extensive soft bed of matrix-rich deformation till reworked from antecedent sediment; the onset of fast ice flow may have triggered the immediate

collapse of the ice sheet in response to unsustainable losses of ice from ice streams calving along its Pacific margin, combined with that lost by calving and disintegration in deep ‘fiord lakes’ under the ice sheet central core zone. The bed of the former Cordilleran Ice Sheet may provide a suitable (and accessible) paleoglaciological analog for that of the modern Greenland Ice Sheet.

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Fig. 26. Size and basal topography of the present day Greenland Ice Sheet (after Rippin, 2013; Morlighem et al., 2014) compared to that of the Late Wisconsin Cordilleran Ice Sheet (Fulton et al., 2004; this study). Note presence below both ice sheets of a comparatively low relief central core zone flanked by a rim of high mountainous relief. The presence of thick subglacial sediments under the Greenland Ice Sheet has been inferred from geophysical data; the extensive soft-bed that underlay the core zone of the Cordilleran Ice Sheet that was inherited from proglacial glaciolacustrine basins, is an exposed analog (Fig. 9).

Acknowledgements Eyles thanks the Natural Sciences and Engineering Council of Canada for generous funding. We are particularly grateful to Alpine Lakes Air in Telkwa, British Columbia for excellent logistical assistance during the course of field work and we thank Kirsten Kennedy and Amalia Kolovos for very helpful discussions in the field and laboratory. Chris Stokes, Neil Glasser, Jeremy Ely, Alain Plouffe and two anonymous reviewers are thanked for their very helpful comments and thoughtful reviews of earlier drafts of the manuscript.

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