Chapter 5. Subglacial Processes and Sediments

CHAPTER SUBGLACIAL PROCESSES AND SEDIMENTS 5 J. Menzies1, Jaap J.M. van der Meer2 and W.W. Shilts3 1Brock University, St. Catharines, ON, Canada, 2Queen Mary, University of London, London, United Kingdom, 3Illinois State Geological Survey, University of Illinois, Champaign-Urbana, IL, United States

5.1 INTRODUCTION Subglacial processes and sediments, in terms of thickness and areal extent, dominate glaciated continents and ocean basins. Subglacial erosional and depositional processes are complex. Incomplete knowledge of the subglacial environment and its processes and landforms stems from the largely inaccessible environment beneath ice masses and the limited modern analogues to Pleistocene and pre-Pleistocene subglacial environments. This chapter is not exhaustive in the use of older literature. For extended reviews of previous literature, the reader is directed, for example, to Menzies and Shilts (2002), Evans et al. (2006), Benn and Evans (2010) and Schomacker et al. (2010a). Subglacial sediments are of paramount importance to society. The topographic and volumetric abundance of subglacial sediments in glaciated countries cannot be understated. Large areas of all continents where glaciation has occurred, since the earliest of geological times, are covered by these sediments (see chapters: Precambrian Glacial Deposits: Their Origin, Tectonic Setting, and Key Role in Earth Evolution; The Early Palaeozoic Glacial Deposits of Gondwana: Overview, Chronology, and Controversies). The foundations of roads, railroads, aircraft runways, the waste disposal sites for sew- age, toxic waste sites, the utilization of agricultural land (see chapter: Glacioaeolian Processes, Sediments, and Landforms), the distribution and movement of groundwater in aquifers, the utilization of glacial sediments in drift prospecting (see chapter: Drift Prospecting in Glaciated Terrain) and use of glacially derived construction materials in those areas last glaciated in the Pleistocene, are all dependent on subglacial sediments, their structural architecture, grain size and stratigraphic complexity. As our increasing awareness of environmental hazards and rational land use develops in the 21st century, the demand for expert and applied knowledge of glacial sediments and environments will rise, with a prime place given to knowledge of the subglacial environment (De Mulder and Hageman, 1989; Coates, 1991; Knight, 2008; Paulen and McMartin, 2009). Historically, major population cen- tres developed in areas that were glaciated because of the rich soils that promoted agriculture provid- ing sustenance for large populations. Perhaps no other glacial environment impacts on the daily lives of millions of people to such a high degree as does the subglacial (Box 5.1). The subglacial environment can be viewed as containing a series of sediment packages that each have boundary interfaces that are sandwiched between an upper ice mass and an underlying

BOX 5.1 DEFINITION Where the subglacial environment ceases and proglacial, submarginal and/or subaqueous glacial environments commence remains, at times, debatable. The subglacial environment is that glacial subsystem directly underlying an ice mass in close contact with the overlying ice, including those cavities and channels beneath the ice that are not directly influenced by subaerial processes. On land, the subglacial environment may continue for some distance beyond the apparent surface ice margin since buried active ice may underlie debris that has been deposited either from supraglacial sources or from various meltwater streams exiting the ice (Fig. 5.1A, B)(Bernard et al., 2014; Banwell et al., 2016; Hiester et al., 2016). Likewise, on entering a body of water an ice mass may be sufficiently thick or the water sufficiently shallow to permit ice to remain in contact with the bed, the basal environment thereby remaining subglacial. Where ice begins to rise off its bed due to buoyancy from lake or oceanic waters, the subglacial environment can be said to cease down-ice of this dynamic hinge point, the grounding-line (Fig. 5.1C). Down-ice from a grounding-line an extensive proglacial subaqueous proximal environment may exist. Such a proximal subaquatic zone may become transiently subglacial for short periods of time and at different places on the bed (see chapters: Modern Glaciomarine Environments and Sediments; Glaciovolcanism: A 21st Century Proxy for Palaeo-Ice).

bed (composed of bedrock or underlying sediments) where a complicated set of processes interact altering the morphological, thermal and rheological state of the sandwiched sediments, evolving into several and, at times, a single interface junction (cf. Smith et al., 2015) (see Fig. 5.8). The nature of this interface controls glacial erosion, transportation and deposition processes. The key to subglacial bed types and associated forms lies in understanding the mechanics of this complex set of interfaces. These boundary zones migrate across the landscape with every advance and retreat of the ice masses. Therefore, all terrains covered by glaciers have been affected and altered by the passage of this boundary zone. Subglacial depositional and erosional phenomena are often destroyed, radically altered or obscured by later processes associated with proglacial environments following gradual ice retreat or by more complex subglacial overriding and subsequent ice retreat, or by subaerial diagenesis. The transport of sediments within the glacial system requires an understanding of the sediment delivery systems (SDS) within supraglacial, proglacial and subglacial subenvironments. Glacial deposits and landforms are end products of these sediment systems (Evans, 2003). Sediment delivery occurs as a function of ice dynamics, the type of ice mass, basement and subjacent geology and sedimentology, the temporal and spatial variability of sediment discharge (flux), associated hydro- logical regimens and topography (Menzies et al., 2016; Spagnolo et al., 2016). Fig. 5.1 illustrates the many and complex pathways of these transport systems. Individual local sediment delivery systems depend on ice internal and basal thermal conditions, whether polar or temperate (Menzies and Shilts, 2002; Frederick et al., 2016; MacGregor et al., 2016). Under polar frozen bed conditions, it was once thought that little or no transport of sediment occurred other than in the supraglacial system. However, evidence now shows that even under polar or subpolar bed conditions, some limited transportation occurs as part of a slow deforming bed (Shreve, 1984; Echelmeyer and Zhongxiang, 1987; Hallet et al., 1996; Alley et al., 1997; Lloyd Davies et al., 2009)(Table 5.1). In contrast to frozen bed conditions, large quantities of sediment are transported in all the sediment subsystems under temperate wet bed conditions (Vorren and Laberg, 1997; Kirkbride, 2002; Ottesen et al., 2005, 2016; O’Brien et al. 2007; Benn and Evans, 2010; Schomacker et al., 2010b)(Fig. 5.1). FIGURE 5.1 Sediment delivery systems (SDS) within glacial environments: (A) SDS within valley glacial systems, (B) SDS within marginal ice sheet systems, and (C) SDS within subaqueous glacial environments. (A) Modified from Boulton, G.S., Eyles, N., 1979. Sedimentation by valley glaciers; a model and genetic classification. In: Schlu¨chter, C. (Ed.), Moraines and Varves: Origin/Genesis/Classification. A.A. Balkema, Rotterdam, pp. 11À23. 108 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

Table 5.1 Subglacial Interface Types

As ice conditions fluctuate within an ice mass over diurnal, seasonal and annual cycles, there changes occur in the volume and rate of sediment transfer. Daily, as discharge of meltwater increases as a function of solar radiation, transport of both bedload and suspended sediment increases within the supraglacial, englacial and proglacial subenvironments. There may be a slight daily increase in subglacial and englacial transport rates as meltwater penetrates a thin ice mass, such as a cirque or valley glacier or the marginal edge of an ice sheet. In general, however, sediment transport in subglacial systems is likely to remain relatively steady, even over an annual cycle (Truffer et al., 2000; Habermann et al., 2013). Seasonal changes in sediment transport are influenced by local weather conditions. When freezing temperatures prevail, supraglacial and subsequently englacial meltwater transport ceases. At the same time, subglacial transport is likely to be restricted, there being little or no additional surface meltwater but frictional heating and geothermal effects maintain some level of subglacial meltwater transport under temperate basal conditions (Ottesen et al., 2005; Nyga˚rd et al., 2007; Quincey and Luckman, 2009; Dowdeswell and Ottesen, 2016). It is evident that sediment transported by ice masses reflects the underlying and surrounding lateral (in the case of valley glaciers) and up-ice bedrock geology (see chapter: Application of Till Mineralogy and Geochemistry to Mineral Exploration). Although ice masses may transport sediment extremely long distances of hundreds of kilometres (Shilts, 1977, 1984, 1993; Meer and Wicander 1992; McClenaghan and Peter, 2016). Much of the sediment load in the subglacial system, is derived from little more than a few kilometres up-ice, from as little as 10À15 km (Dreimanis and Vagners, 1971; Hallet et al., 1996; Menzies and Shilts, 2002; McMartin and Paulen, 2009). Much of this work on glacial transport was first introduced by Dreimanis and Vagners (1965, 1971, 1972) using the concept of ‘terminal grade’ that indicates that a particular bedrock type or mineral is ground down until a grain size is reached where further size reduction 5.2 EROSION IN THE SUBGLACIAL 109

does not occur because of the mineral’s physical characteristics (the terminal grade) (Haldorsen, 1978, 1981; Anderson, 2005). It might be said that the harder a rock type, the longer it is likely to survive in transport (cf. Shilts, 1995; Levson 2001; Larson and Mooers, 2004; Hooke et al., 2013). In areas of previous glaciation, sediment is commonly sourced directly from relict deposits. Sediment from confined valley systems provides overwhelming supplies of transportable debris from the surrounding steep terrains, where debris flows, landslides, rock avalanches, and surface runoff in the summer supply considerable volumes of sediment (Hewitt, 2009)(Fig. 5.1A). It has been demonstrated that as basement bedrock types vary in terms of erosivity, the sediment yields for transport change drastically (Shilts, 1973; Hallet et al., 1996; Alley et al., 1997; Anderson, 2005; Sugden et al., 2005; MacGregor et al., 2009; Norton et al., 2010; Herman et al., 2011; Cook and Swift, 2012). Slow-moving, basal frozen, polar ice masses sliding over hard crys- talline bedrock yield many magnitudes less eroded material than fast-moving temperate ice crossing soft sedimentary rock types (Colgan et al., 2002; Bonow et al., 2007; Du¨hnforth et al., 2010; Corbett et al., 2013; Frederick et al., 2016; MacGregor et al., 2016). Sediment transport pathways within glacier/ice sheet systems are intricate and much remains to be understood about them (Gurnell and Clark, 1987; Evans, 2003; Jamieson et al., 2010, 2014; Cook and Swift, 2012; Jaeger and Koppes, 2016). Large temporal and spatial variations in sediment transport processes and flux rates occur in subglacial subenvironments (cf. Schomacker et al., 2010a; Menzies et al., 2016). Although flux rates have been studied extensively in valley glacier systems (Kirkbride, 2002; Benn et al., 2003) there has been limited discussion of sediment flux rates, up until recently, below present-day ice sheets, ice streams and ice shelves (cf. Alley et al., 1989; Dowdeswell and Siegert, 1999; Ottesen et al., 2005; Dowdeswell et al., 2006, 2010; Bradwell and Stoker, 2015; Spagnolo et al., 2016). Ice streams, for example, transported sediments to the margin of the northern portion of the Barents-Kara Sea ice sheet at a rate of B4cma21 (0.13 cm a21 averaged over the fan) (Dowdeswell and Siegert, 1999). Fast-flowing ice streams are responsible for considerable sediment transfer by ice sheets (Dowdeswell and O´ Cofaigh, 2002; Ottesen et al., 2005; Dowdeswell et al., 2006, 2010; Bingham et al., 2010; Andreassen et al., 2014; Kyrke-Smith et al., 2014; Perol et al., 2015; Frederick et al., 2016; Hiester et al., 2016). However, these are relatively narrow zones of fast-moving ice and it should not be construed that basal ice elsewhere does not transport large quantities of sediment. Care needs to be exercised in attempting to estimate such sediment flux rates as only limited ‘snapshots’ of subglacial conditions exist at present. It is interesting to speculate that a relationship would appear to exist between drumlins, ribbed moraine, and the rate of sediment discharge at the iceÀbed interface (Meer et al., 2003; Dunlop et al., 2008; Smith et al., 2015; Menzies et al., 2016; Spagnolo et al., 2016).

5.2 EROSION IN THE SUBGLACIAL As the active subglacial environment does not allow the direct observation of processes, we must deduce these processes from their legacy. The legacy of erosion can be seen on both rock surfaces and as hiatuses, in some cases, sharp erosional contacts within sediment packages. Following glacial retreat rock surfaces show many erosional features on a range of scales. The smallest of these are striae (Fig. 5.2AÀD), long scratches that follow the direction of ice movement. They are produced when harder rocks carried by the ice are dragged over the surface or when soft sediment is deformed or squeezed under hydrostatic pressure between the ice and the 110 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

FIGURE 5.2 (A) Cross-cutting striae, Pyhajarvi, Finland. (B) Striae across bedrock surface, Grimsel Pass, Switzerland, note nail head striae in foreground with coin and oblique chattermarks and fractures. (C) Striations across a Silurian bedrock surface, Port Colborne, Ontario, approximately 3 metres from side to side. (D) Striae on flat Palaeozoic dolomite beds south of the Shield boundary in Manitoba. (E) Striations at Northern College Campus, Kirkland Lake, Ontario. Older steep side wall striations trending 205 (black) are the oldest, followed by younger flat lying surface striations at 175 (yellow) on the protected lee (south) side of the outcrop. The youngest ice flow has striated the uppermost surface towards 145 (red). (D) Courtesy of I. McMartin, GSC. (E) Courtesy of Beth McClenaghan GSC (McClenaghan et al., 1995). bed (Dahl, 1965; Shaw, 1988; Krabbendam et al., 2016). Observing striae, at times, can permit the direction the ice moved to be defined (Fig. 5.2E). Isolated glacial striae seldom occur, but affect whole rock surfaces, demonstrating that striation leads to aerial abrasion. When observed, in detail, it is clear in some cases, that the rocktool is not dragged over the surface in a smooth movement, but that it is actually jerky; it is repetitive localized high-stress percussion (Lister, 1981; Riley, 1982; Sharp and Gomez, 1986; Tulaczyk et al., 1998)(Fig. 5.3). Typically, striated rock surfaces and clasts carry multiple striations often of varying directions, at times it is not uncommon to find clasts with striations on their underside at 180 degrees opposite on their upper surfaces, indicative of clast rotation within the basal ice, at the iceÀbed interface or within soft deforming sediment layers. Striated surfaces often also show crescentic gouges where lunate fragments have been removed from the surface (Fig. 5.4A, B). The horns of the lunate fracture, typically, point down-ice and are thus a good indicator of ice direction. 5.2 EROSION IN THE SUBGLACIAL 111

FIGURE 5.3 Ribbed striae as a result of skipping movement of rocktool—white strokes indicating contact points, Sandflugtdalen, W. Greenland.

FIGURE 5.4 (A) Crescentic gouge at Slettjokull,¨ Iceland. Tests showed that a force of 156 MPa was needed on this basalt rock to produce the gouge. (B) Chattermarks on flat Palaeozoic dolomite beds south of the Shield boundary in Manitoba. Courtesy of I. McMartin, GSC. 112 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

Many striated surfaces also show smoothed P-forms (glacially eroded, plastically moulded smooth-walled linear and/or convoluted depressions; often straight or curved, extending a few cen- timetres to many tens of metres). These forms indicate that subglacial meltwater played a part in their formation, suggesting an alternation beneath water flowing at the contact between ice and rock and ice being in direct contact with its bed (Shaw, 1988; Lowe and Anderson, 2003; Rea, 2007; Eyles and Doughty, 2016)(Fig. 5.5AÀD). To explain how ice can remove bedrock, refreezing (regelation) was thought to be the mechanism. When basal ice meets a bedrock bump, the pressure on one side would lead to pressure melting, the water penetrating fractures in the rock on the other end of the bump and refreezing. Consequently, a block or some part of the bedrock would freeze to the bottom of the ice and be quarried (cf. Ugelvig et al., 2016). However, under temperate ice, refreezing is much less likely to occur on a large scale or underneath the whole glacier. As many glaciers are underlain by moving basal debris, at times attached to the basal ice while in other locations separated from the basal ice by a sharp discontinuity, adfreezing must be patchy at best and occur at the top of this sediment (Lawson et al., 1998). In recent years, the role of pressurized subglacial water has been recognized as playing an impor- tant role in glacial erosion. It is now known that this water pressure can be high enough to pry apart

FIGURE 5.5 (A) Roche moutonnees,´ Turtmanngletscher, Switzerland. (B) P-forms across bedrock hillslope, near Grimsel Pass, Switzerland. (C) P-forms of Quaternary age across a Precambrian diamictite boulder near Whitefish Falls, Ontario, Canada. (D) P-forms on Kelleys Island offshore in Lake Erie, Ohio, USA. The scale is approximately 5À6 m wide. 5.2 EROSION IN THE SUBGLACIAL 113

rocks once the water penetrates fractures and joints (Rijsdijk et al., 1999; Phillips et al., 2013a; Blignault and Theron, 2015; Evans et al., 2015; Lelandais et al., 2016), a process known as hydrofracturing. Other processes involving freeze on and abrasion can lead to similar forms being developed. These various processes may lead to the up-ice end of a bump (protuberance) being smoothed and rounded, while the down-ice end is rough and stepped (cf. Rastas and Seppa¨la¨,1981; Roberts and Long, 2005; Krabbendam and Glasser, 2011)(Fig. 5.5A, B). The resulting feature is known as a roche moutonnee.´ Roche moutonnees´ come in many sizes, the largest (hill- to mountain-sized) being known as whalebacks (Roberts and Long, 2005; Bradwell et al., 2008)(Fig. 5.6A). Pressurized water also explains erosion of bedrock where there are no discernible protuberances.

FIGURE 5.6 (A) Bedrock eroded by cold-based ice, Alan Hills, Antarctica. (B) Roche moutonnees´ in the proglacial zone, Nygardsbreen, Norway. 114 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

The largest forms resulting from glacial erosion are basins and fjords. Large basins can be formed in both sediment and rock, while fjords are only found in rocks. Many basins are associated with glaciotectonics (see chapter: Glacitectonics; de Gans et al., 1987; Aber and Ber, 2007). Fjords are characterized by their great length compared to their width, they are overdeepened and many show one or more sills (Syvitski et al., 1987; Margreth et al., 2016; Lelandais et al., 2016). Furthermore, they only occur in mountainous areas that formerly carried ice sheets. Basins can also be overdeepened as demonstrated by, for instance, the circumalpine basins in Europe (Schlu¨chter, 1979a,b), the Finger Lakes in New York State, United States (Mullins et al., 1996) or the lakes to the east of the Andes in Patagonia, and most lakes on the Canadian Shield. Glacial erosion is not restricted to bedrock; sediment packages are eroded as well. This erosion can occur at the ice bed contact, in the sediment pile (Truffer et al., 2000), as well as at the base of a deforming bed (Meer et al., 2003; Eyles and Doughty, 2016). Overall it can be said that eroding glaciers are nondiscrimi- natory when it comes to size, they can literally move mountains. For example, in Alberta large bedrock masses that appear to have been removed (‘pop-up hills’) from depressions up-ice from them were well recognized (see Slater, 1927; Kupsch, 1962; Christiansen and Whitaker, 1976). Previously, it was thought that cold-based ice did not have any erosional effect as the ice was thought to be frozen to the bed with all forward movement occurring within the ice. Recent work in the Antarctic has shown that this is not the case, cold-based ice does erode its bed, but not aeri- ally. Erosion, both abrasion and quarrying, is concentrated on bedrock steps (Fig. 5.6B)(Davis et al., 2006; Kleman et al., 2008; Lloyd Davies et al., 2009). Given the limited extent of these features their preservation potential after deglaciation and subsequent climate change is very small.

5.3 TRANSPORT: MECHANICS Glaciers transport debris at different levels: supraglacial (see chapter: Supraglacial Environments), englacial, basal (Fig. 5.7A) and subglacial (Boulton, 1978; Alley et al., 1997). Although this is largely passive transport, these levels are not separated but interconnected (Schlu¨chter, 1979a,b; Lukas et al., 2013). Supraglacial debris can enter the glacier by falling into crevasses or by washing in by supraglacial meltwater; it will thus become englacial or even basal debris (Fig. 5.7B). Englacial debris can become supraglacial by movement along rising shear planes or by surface melting. Similarly, it becomes basal debris by basal melting or by transport in meltwater. Basal debris can become supraglacial in the same ways as englacial debris, but this will occur mainly near the snout. Of course, when a glacier tumbles in an icefall all the debris will be mixed up. Although debris at the supraglacial level can typically be easily distinguished from basal debris (angular vs rounded and striated) the mixing process means that each level can still contain material that is suggestive of the other. It is instructive to indicate that not all supraglacial debris is necessar- ily angular. It is not uncommon in boulder trains and as ablation deposits in many glaciated areas to find well-rounded clasts and boulders transported and deposited from supraglacial environments depending on bedrock sources, and ice basal scavenging (debris entrainment) by removal of rock and debris causing, through debris entrainment, clasts to rise within the ice to the supraglacial (Shilts, 1973; DiLabio and Shilts, 1979; Lawson, 1995; Darvill et al., 2015; Lovell et al., 2015). 5.3 TRANSPORT: MECHANICS 115

FIGURE 5.7 (A) Folding of basal debris leading to thickening of the basal debris layer; rock in centre is ca. 40 cm long, Glacier de Tsijiore Nouve, Arolla, Switzerland. (B) Debris within supraglacial ice on the Stein Glacier, Switzerland. (C) Striated boulder within proglacial debris at margin of Stein Glacier, Switzerland. 116 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

Because of the typical velocity profile of a glacier or ice sheet (Paterson, 1994, Fig. 11.12) debris that has become part of the glacier at various locations either on the surface or from the sides (as in the case of valley glaciers) and from beneath, will not reach the frontal margin at the same time. Due to dif- ferential velocity, debris particles—especially at the basal level—will be forced into contact with each other. Depending on difference in hardness and whether the ice holding the particles will melt under the pressure, the particles may erode each other. Due to this, larger particles become polished and striated (Fig. 5.7C), while sometimes they are broken up. These comminution processes are also supposed to result in the sandy nature of many tills (Haldorsen, 1981; Paulen and McMartin, 2009). Since forward movement in the bed can account for a large proportion of the total glacier velocity, mass transport of sediment at the bed must be substantial, only depending on the thickness of the deforming layer. The latter is one of the biggest unknowns in glacial geology, having been observed to be very thin (Fuller and Murray, 2000)tomanymetresthickintheory(Boulton, 1996), while all tills show extensive evidence of sediment deformation at both macro- and microscales (Meer et al., 2003). Transport in the glacier bed happens at different scales. On the one hand, it is microscopic by rotation and shearing (see chapter: Glacitectonics) as well as by liquefaction (Phillips et al., 2013b). On the other hand, macroscopic by squeeze and even by movement of the whole sediment pile (Kjær et al., 2006) after detachment at its base. As large erratics and bedrock rafts in all glaciated areas dem- onstrate, ice masses can also transport all sizes of rocks and sediment packages. However, after being released by the ice, the largest erratics cannot be transported by water and come to rest. Despite all this, most of the debris eroded by glaciers is moved out of the system by running water (Marren, 2005). One only must look at meltwater rivers when they emerge from the ice to see how charged they are with debris. Furthermore, when comparing sediment wedges on the continental shelves, as compared to glacial debris on land, the sediment loads on the shelves are multiples of what is on land.

5.4 DEPOSITION: MECHANICS The mechanics of subglacial deposition of debris from ice masses are based on the bond between debris within the ice being released through passive melt-out, active pressure melt-out, by the effect of meltwater melting the ice or by the immobilization of subglacial debris at the basal interface between the bed and active ice, this later process is possibly volumetrically the greatest. A major further source of sediment deposition is due to gravitational debris flows released subglacially or owing to the immobilization of existing subglacial sediment under soft deforming bed conditions (Fig. 5.8). Primarily sediment release can be achieved when debris is pressed against the bed or against other debris and its internal frictional strength is greater than that applied by the forward movement of the active ice or mobile sediment. The debris may become lodged (Fig. 5.9) (lodgement/emplacement) or immobilized due to a reduction in external stress levels or increases in internal shear strength of the material, both resulting in immobilization of the debris (traction/melange).´ Other release mechanisms occur when the ice ceases forward movement, becomes passive (dead) and the ice slowly melts away through basal melting (melt-out) [cf. Bering Glacier in Alaska (Fleisher et al., 2006, 2010)]. Lodged or emplaced particles can be re-activated or they can be enclosed in further debris accumulation, after which they can become part of the deforming bed. Deformation will then obliterate many or all previous structures (Menzies, 2012), except for the largest boulders at the surface (Fig. 5.10A). Melt-out should preserve the structure and distribution 5.4 DEPOSITION: MECHANICS 117

FIGURE 5.8 A schematic diagram of subglacial conditions beneath an active soft deforming bed. Note position of upper (If1) and lower (If2) interfaces, and ‘H’ hard undeforming, ‘QM’ quasi mobile, ‘QH’ quasi-hard bed conditions. Blue lines indicate broken (porewater) and unbroken (open meltwater) hydraulic activity.

FIGURE 5.9 Lodged boulders at base of till and overlying rhythmites; note dragged water escape structure, San Martin de los Andes, Argentina. 118 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

FIGURE 5.10 (A) Group of lodged boulders remaining in place while the finer sediment deformed around the group, Slettjokull,¨ Iceland. (B) Meltout structures in Lennoxville Till near the confluence of Rivieres Fraser and des Plante. (C) Basal debris bands melting out of basal ice and exposed on the side of Aktineq Glacier on Bylot Island, Nunavut, Canada. 5.5 SUBGLACIAL SEDIMENTS 119

of the debris in the ice (Larson et al., 2003, 2015). Should the deposit become re-activated in the deforming bed by renewed glacier activity, melt-out structures will likely be obliterated or badly disrupted (Paul and Eyles, 1990; Moller,¨ 2010; Menzies, 2012, Fig. 6b; Cook et al., 2011). In areas of low glacier bed relief, melt-out deposits may have a low survival potential because they, typically, are subject to subsequent reworking within active glacial environments (cf. Evans et al. 2006, Table 6; Larson et al., 2015). However, some melt-out tills deposited marginally and directly from supraglacial sources and reworked by meltwater streams do in some instances survive as large sediment packages. Examples from the Riviere des Plantes in southern Quebec (Fig. 5.10B) and other places in the northern Appalachians, and the margin of the Aktineq Glacier on Bylot Island, Nunavut, Canada illustrate such melt-out tills (Fig. 5.10C).

5.5 SUBGLACIAL SEDIMENTS Subglacial sediments can be subdivided into sorted (glaciofluvial and glaciolacustrine) and unsorted (tills or, with no genetic connotation, diamicts or diamictons) (Table 5.2). The latter are, at times, difficult to identify in terms of origin, location and type. Some controversy surrounds the usage and definition of the term, till (Hambrey and Harland, 1981; Dreimanis, 1989; Menzies and Shilts, 2002). Tills are polygenetic and typically regarded as of continental (on land) origin and, as such, are very easily confused with glaciomarine/glaciolacustrine diamictic sediments in certain locations, and in other locations may have intercalated debris flows, flow tills both land-based and marine, and

Table 5.2 Deposition and/or Emplacement Processes in Subglacial Environments 120 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

volcanic lahars (Shilts, 1978; Dreimanis, 1989; Meer et al., 2003; Evans et al., 2006; Meer and Menzies 2011; Evans and Benn, 2014). As diamicts, the subglacial till matrix is often dominated by locally sourced materials (, 15 km up-ice), however, where fine-grained matrix is produced these clay- and silt-sized sediments may persist within down-ice sediments for 10s and even 100s of kilometres. The sediments exhibit a wide range of grain sizes, from coarse sands to clays, clasts (pebbles to small boulders) and are typically subangular to subrounded, and have striae on their surface facets. Stratigraphically, subglacial tills contain intraclastic units ranging from a few centimetre- sized stringers to occasionally large block units of sorted sands and gravels that typically have sharp contacts with the more clay-rich till and, like many of the larger clasts and boulders, ‘float’ within the sediment matrix. Many of the intraclasts occur at a range of angles, some exhibiting rotation while others appear boudinaged. In a few cases these tills contain little or no large clasts and are matrix-dominated. In many places till can be seen to vary in matrix colour, the result of differently sourced sediments (cf. Perrin et al., 1979; Hamblin et al., 2005). Geotechnically, tills are often clas- sified at the macroscale as massive and are often overconsolidated but on closer examination tills exhibit structure, are stratigraphically complex and at the microscale exhibit ‘bedding’, shear lines and a multitude of complex microstructures (McGown and Radwan, 1975; Brockman and Szabo, 2000; Phillips et al., 2002; Meer et al., 2003; Menzies et al., 2006; Klint et al., 2010; Meer and Menzies 2011, Waller et al., 2011; Kessler et al., 2012; Narloch et al., 2012; Menzies et al., 2016; Gehrmann et al., 2017). Tills may form within several distinct glacial subenvironments: subglacial, englacial, proglacial (terrestrial/subaquatic) and supraglacial; and may be deposited from active temperate wet-based and polar cold-based ice masses, under stagnant ice conditions, and subaquatic environments of both grounded and floating ice masses (Fig. 5.1A, B). Table 5.1 illustrates the complexity and interrelationships that exist within till-forming glacial environments. Within subglacial wet-based, active temperate conditions considerable input from meltwater can be expected, as well as thermal fluctuations, and transient fluctuations in ice velocity and stress conditions (Fig. 5.8; see Box 5.2). Subglacial tills formed under active temperate conditions used to be recognized as lodgement, flow

BOX 5.2 SUBGLACIAL INTERFACES Subglacial interface types formed under polar, temperate and polythermal bed conditions are best illustrated in the figure (Fig. 5.8) and considered as a mosaic of subglacial interface states (Meer et al., 2003; Piotrowski et al., 2004). In the figure the conditions at the base of an ice mass are highlighted such that conditions may be frozen (polar), thawed (temperate) or a mix of both (polythermal). Under these varying states sediments may be frozen, or saturated and thawed. Under the overlying strain applied by the ice, sediment may, under certain conditions and levels of saturation (effective stress levels), be susceptible to being deformed (mobilized) or, if saturation is too low and effective stress levels too high, the sediment remains undeformed (immobile). Where conditions permit, meltwater, under temperate and polythermal conditions, may permeate the underlying sediment and bedrock, or may not. If the sediment and/or bedrock is of very low permeability then meltwater does not permeate and most, if not all, the meltwater remains at the iceÀbed interface (If1) and moves across the bed as discrete sheets, or more likely, channelized water. Where permeation occurs, the sediment may become saturated and subject to deformation in other instances meltwater may permeate a highly permeable sediment/bedrock and flush out of the ice/bed system. Where sediment deforms, a lower interface (If2) or plane of decollement´ will develop that may fluctuate across the subglacial bed to form a mobile layer of deforming sediment of widely fluctuating thickness (a deformation ‘front’) (Fig. 5.8). It is likely at the interface between underlying immobile and mobile sediment that various bedforms may develop. 5.5 SUBGLACIAL SEDIMENTS 121

and melange´ (traction), with a variety of local till subtypes (cf. Menzies, 2003, 2012; Evans et al., 2006; O´ Cofaigh et al., 2007; Menzies et al., 2006; Leeman et al., 2016). Subtypes of tills are often identified owing to some ‘specific’ characteristic of grain size, provenance and/or structural attribute or even distinctive colour and clast content. Where ephemeral passive ice conditions occur within active ice masses melt-out tills can be deposited. Although comparatively rare, under dry polar cli- mates sublimation tills can form (Shaw, 1989; Hindmarsh et al., 1998; Marchant et al., 2002). Both lodgement and melt-out tills have been extensively described in the literature (Dreimanis, 1989; Ruszczynska-Szenajch,´ 2001), whilst subglacial melanges´ (traction tills) have only recently been recognized as separate, distinct subglacial sediments (Curry et al., 1994; Hoffmann and Piotrowski, 2001; Menzies and Shilts, 2002; Evans et al., 2006)(Tables 5.2, 5.3). These subglacial tills are formed as a consequence of: (1) high shear stress resulting in debris being released from basal ice in a ‘smearing on’ process where high pervasive shear stress results in ice pressure melting and debris release (lodgement till); or (2) by slow undermelt releasing debris gradually (melt-out till) (Shaw, 1987; Larson et al., 2015); or (3) as a result of portions of a soft deforming subglacial bed either reaching shear strengths higher than the mobilizing basal ice stresses due to onset of freezing

Table 5.3 A Schematic Till Classification (see Menzies, 2003) 122 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

temperatures, or reduced porewater content, or porewater pressure reduction due to reduced overbur- den stress (Menzies, 1989a; Fuller and Murray, 2000; Leeman et al., 2016; Sobota et al., 2016) producing melange;´ and finally; (4) by the glacial crushing of basal debris and bedrock (H-bed conditions; Menzies, 1989a,b) resulting in a form of pulverized bedrock being produced (Haldorsen, 1981; Aylsworth and Shilts, 1989a,b, pp. 413, 414) (Fig. 5.8, see Box 5.2). There are several divergent ideas concerning till origins and classification, for example, that tills typically exhibit a bimodal grain size distribution, are overconsolidated, have very low permeabil- ities, are unstratified, homogeneous, massive, and unstructured, and, finally, that tills of differing colours are of different depositional phases and therefore indicative of more than one glaciation (Fig. 5.11). Although many tills do exhibit many, if not all, of the above characteristics many other tills do not. Where tills can be seen forming at present the usage of the term till seems appropriate. Where, as is so often the case, either the actual means of deposition is obscured or the sediments are ancient, the till-forming mechanisms and environments within which they were originally deposited are unknown; in such cases the nongenetic terms diamict or diamicton for unlithified sediments and diamictite for lithified sediments may be preferred. A persistent, and perhaps unsolv- able problem, remains in identifying as tills those sediments that are marine as opposed to glaciomarine or lacustrine compared with glaciolacustrine (Meer and Menzies 2011). With the advent of glacial micromorphology, a clearer understanding of the internal in situ stratigraphy and microstructures furthers understanding of this complex, and often enigmatic, glacial sediment (Meer et al., 2003; Menzies et al., 2016) (see chapter: Glaciohydrogeology). The classification of tills has undergone considerable modification in recent years as the style and method of deposition and/or emplacement has been discussed (Evans et al., 2006; Menzies, 2012)(Table 5.4). Where the terms, lodgement till, melt-out till, flow till, waterlain and ablation till were often used rather indiscriminately it has become apparent through detailed studies of subglacial environments that such tills may often be peripheral to the dominant form of subglacial tills, namely deformation tills, traction tills and melange-style´ tills, and perhaps tectomicts that all lead to a different and more complex view of subglacial till formation (Evans et al., 2006; Menzies et al., 2006; Menzies, 2012: Phillips et al., 2013a; Larson et al., 2015; Gehrmann et al., 2017). It can be generally assumed that most subglacial tills are deposited or emplaced under temperate wet-based subglacial conditions where meltwater may be active at the iceÀbase interface and that sediments may be saturated and subject to overlying basal ice strain leading to macro- and microdeformation. The so- called soft-sediment deformation hypothesis appears to explain much of till sedimentology under these conditions (Murray, 1992; Fuller and Murray, 2000; Thomason and Iverson, 2006; Brisbourne et al., 2017). However, there are other instances where cold-based ice masses, such as those on Bylot and Baffin Islands, do produce tills (DiLabio and Shilts, 1979; Echelmeyer and Zhongxiang, 1987; Zdanowicz et al., 1996; Cuffey et al., 1999, 2000; Atkins et al., 2002). As Menzies and Shilts (2002) suggested there are likely several soft subglacial interface conditions from fully melted to dry bed states under near-frozen bed interface states (Box 5.2). The concepts of quasi mobile (QM) to quasi-hard (QH) conditions may lead to melange´ tills (traction tills) under fully saturated soft bed deformation conditions. Where meltwater cannot infiltrate the bed under Q or H conditions, lodgement tills may be deposited under high strain rates (Menzies, 2012)(Table 5.1). A new till classification table is shown in Table 5.4 that follows this set of bed interface subglacial conditions. Subglacial stratified sediments have been traditionally thought of as indicative of sediments of proglacial origin that have been incorporated by active ice into the subglacial environment. In 5.5 SUBGLACIAL SEDIMENTS 123

FIGURE 5.11 (A) Till section, Estonia, showing multiple colour contrasting sublithofacies units within a single lithofacies package. (B) Till exhibiting planar colour changes within a single lithofacies unit, Oakville, Ontario, Canada. (C) Till showing at least two sublithofacies units in a single till lithofacies unit, Pine Point Mine, Pit K-62, NWT, Canada. (D) Laminae of silt deformed around rock clasts (centre) part of meltout till unit, Matanuska Glacier, Alaska, USA. Courtesy of Larson, G.J., Menzies, J., Lawson, D.E., Evenson, E.B., Hopkins, N.R. 2015. Macro- and micro-sedimentology of a modern melt-out till—Matanuska Glacier, Alaska, USA. Boreas 45, 235À251. many cases these sediments, typically sands and gravels, are deposited within channels and tunnels or as thin sheet-like layers at the iceÀbed interface, especially under ‘hard’ bed conditions (Eyles et al., 1982; Ravier et al., 2014; Brisbourne et al., 2017). Many tills contain stratified sediment intraclasts indicative, in numerous instances, of subglacial meltwater deposition. It is now 124 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

Table 5.4 Release Mechanism and Resultant Till Types

commonly recognized that within subglacial cavities, glacier bed depressions, channels and tunnels, considerable volumes of stratified sediments may be deposited, ranging from delicately laminated lacustrine clays to coarse bouldery gravels to small and large till balls (e.g., Menzies and Woodward, 1993; Goldschmidt, 1994; Roberts and Hart, 2005)(Fig. 5.12). These sediments exhibit evidence of large short-term fluctuations in subglacial meltwater discharge, velocity and depth. This variability in meltwater stream competency is manifest in a wide range of sedimentary facies and internal structures, and grain size distributions. Sedimentation is usually rapid (within hours), causing these sediments to exhibit rapid porewater expulsion and migration features, as well as load and shear structures. Subglacial glaciofluvial sediments tend to be areally restricted; occurring as thin patches or linear ribbons (eskers) usually parallel or subparallel to the ice direction (Fig 5.13). Subglacial and proglacial stratified sediment can be segregated on the basis that the latter are tabular, widespread, thick ( . 15 m) sediment packages that may have pitted surfaces pock-marked by kettles. Subglacial stratified sediments commonly reach thicknesses ,15 m, and are areally limited in distribution. Owing to overriding by ice or mobile sediment, stratified sediments may exhibit evidence of liquefaction, fluidization, faulting and shearing (Fig. 5.14). The lack of sediment cohesion usually pre- cludes folding but postdepositional cryostatic conditions may support folding and buckling. In 5.5 SUBGLACIAL SEDIMENTS 125

FIGURE 5.12 (A) Fine-grained till units with delicate sand stringers, Lichtenegg, southwest Germany. (B) Heavily brecciated till with fine sand intraclast below with a small clast at the contact between the two sublithofacies units, Mohawk Bay, Ontario, Canada. (C) Banded till units close to base of a drumlin, Chimney Bluffs, New York State, USA. (D) Brecciated till with distinct ‘bedding’ planes lined with fine to coarse sands, Pine Point Mine, pit O-28, NWT, Canada.

some cases, these sediments may be ‘rafted’, becoming part of subglacial melanges.´ Within subglacial stratified sediments, inclusions are common in the form of till balls, injections from overlying sediments and underlying diapirs, clay shards, and other ‘exotic’ sediments that may have fallen into subglacial meltwater streams and cavities. Where dense clay sediment overlying saturated stratified sediment has suffered rapid loading, dish and pillar structures and, in some cases, complete ball and pillow structures may develop (Fig. 5.15A, B). FIGURE 5.13 An intradiamicton lens of trough cross-stratified coarse sand to granules, overlain by dropstone-rich laminated clays. White arrowheads indicate dropstones. Courtesy of Ravier, E., Buoncristiani, J.-F., Clerc, S., Guiraud, M., Menzies, J., Portier, E. 2014. Sedimentological and deformational criteria for discriminating subglaciofluvial deposits from subaqueous ice-contact fan deposits: a Pleistocene example (Ireland). Sedimentology 61, 1382À1410.

FIGURE 5.14 An SSE-verging recumbent fold overlying two NNW-verging overturned folds. Note that the recumbent fold is truncated by the overlying diamicton. The dotted lines correspond to the limit among three different horizons (H1/H2/ H3) of deformation. The interpretation of this section is based upon a deforming bed model when ice is coupled to the bed, explaining the reverse directions of tectonic transport between the Horizons 1 and 2. Shearing occurs at the ice/ H1 interface and at the H1/H2 interface along a slip surface. H3 represents the stable horizon with no deformation. Courtesy of Ravier, E., Buoncristiani, J.-F., Clerc, S., Guiraud, M., Menzies, J., Portier, E. 2014. Sedimentological and deformational criteria for discriminating subglaciofluvial deposits from subaqueous ice-contact fan deposits: a Pleistocene example (Ireland). Sedimentology 61, 1382À1410. 5.6 SUBGLACIAL LANDFORMS/BEDFORMS DIRECTLY ATTRIBUTABLE 127

FIGURE 5.15 Load structures within tills. (A) A load structure at Mohawk Bay, southern Ontario. Structure is approximately 1 m across at base of structure. (B) Pink diamict, derived from the Athabasca Basin of northeastern Alberta within a dark, locally derived Cretaceous shale-rich overlying till. Courtesy of Paulen, R.C., Paradis, S., Plouffe, A., Smith, R.I., 2011. Pb and S isotopic composition of indicator minerals in glacial sediments from NW Alberta, Canada: implications for ZnÀPb base metal exploration. Geochemistry: Explor. Environ. Anal. 11, 309À320.

Subglacial glaciolacustrine sediments are often found in the leeside of large topographic obstructions, boulders or on the floor of cavities. In the case of the Hudson’s Bay Lowlands, however, subglacial glaciolacustrine sediments extend for kilometres (Andrews et al., 1983; Thorleifson et al., 1992). The presence of small units of glaciolacustrine sediment, within other subglacial sediments, is common (Ravier et al., 2014, and references therein). However, stratified inclusions in tills should not automatically be judged to have formed subglacially, as glaciers are well capable of plucking unlithified, preglacial or proglacial sediment clasts and rafts and transport them long distances.

5.6 SUBGLACIAL LANDFORMS/BEDFORMS DIRECTLY ATTRIBUTABLE TO ACTIVE/PASSIVE ICE ACTIVITY At the subglacial interface, fluctuations or perturbations can occur that translate into landforms/bedforms (Aylsworth and Shilts, 1989a,b; Clark et al., 2009; Johnson et al., 2010; Hooke et al., 2013; O´ Cofaigh et al., 2013; Stokes et al., 2013a,b; Jo´nsson et al., 2014; Trommelen and Ross, 2014; Dowling et al., 2015; Moller¨ and Murray, 2015; Banwell et al., 2016; Menzies et al., 2016; Peters 128 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

et al., 2016)(Table 5.5). Such perturbations appear to be the result of interactions across this set of basal interfaces due to changes, for example, in basal ice strain, basal ice velocities, thermal vari- ables, sediment rheology, bedrock lithology and bulk strength, sediment flux rates, pre-existing topographies and interface ‘roughness’. Such perturbations that can and do occur translate in some

Table 5.5 Directly and Indirectly Iinduced Land-/Bedforms 5.6 SUBGLACIAL LANDFORMS/BEDFORMS DIRECTLY ATTRIBUTABLE 129

instances into bedforms that align parallel or transverse to the main flow direction of the overlying ice mass or are roughly topographically planar forms (Lamsters et al., 2016). A fundamental question that still requires an adequate solution is why do such interface perturbations occur where and when they do? Is there an inherent instability in basal ice interface conditions (Dunlop et al., 2008; Clark, 2010; Fowler, 2010a,b; Stokes et al., 2013a; Hillier et al., 2016; McCracken et al., 2016) that results in forms being developed or evolved at various scales and orientations and why? To answer these critical questions, an improved understanding of basal ice dynamics and interface thermal conditions and sediment rheologies of the basal load is necessary (Hart, 1997; Hindmarsh, 1999; Schoof, 2007; Dunlop et al., 2008; Menzies and Hess, 2013; Vreeland et al., 2015; McCracken et al., 2016; Menzies et al., 2016). In trying to understand the origins of drumlins, fluted moraine, megascale glacial lineations (MSGLs), Rogen moraine or other basal ice interface forms, there are several crucial, yet enigmatic, aspects of their origins and development to be considered. (1) Are the patterns formed distinctive and would they appear related to the position and location beneath the ice? (2) Are the forms developed as a group over a short period or over longer and repeated episodes of formation? (3) Most forms are composed of a variety of sediments rather than just till (Hoffmann and Piotrowski, 2001; Menzies and Brand, 2007; Menzies et al., 2016). (4) Most forms developed parallel to ice motion would seem to be elongated because of either higher ice velocities or the subglacial sediment flux rate or a combination of both factors, also perhaps in relation to the bedforms’ positions close to or beneath ice streams (cf. King et al., 2016). (5) It seems likely that all forms developed at the iceÀbed interface are not developed by the same set or a combination of processes. (6) In many instances, overprinting or reorientation of some or all forms likely occurs. (7) There may well be interrelationships between these forms and other appar- ently unrelated forms, such as drumlins and end moraines, or between the interface forms and topographic slope or proximity to lakes or other large bodies of water where basal ice dynamics suddenly change. (8) Finally, is it possible that there is a relationship between the size, morphology and shape of, at least, the parallel forms (drumlins, fluted moraines, etc.) and the sediment flux rates at the iceÀbed interface (Spagnolo et al., 2014; McCracken et al., 2016; Vreeland et al., 2015; Menzies et al., 2016)?

5.6.1 DRUMLINS Drumlin shape can deviate considerably from the classical tear-shaped form so commonly portrayed in textbooks (Fig. 5.16A, B). Drumlins can vary in size from a few metres to over 200 m in height and can stretch a few metres long to over a kilometre. Typically, they form in large ‘swarms’ or fields many thousands in number (cf. Clark et al., 2009; Hess and Briner, 2009)butmayalsooccur singly or in small groups. Drumlins typically contain a wide range of sediments from single tills to multiple tills, sometimes in ‘onion skin’ carapaces, to occasional bedrock cores, or boulder piles, sand cores or, at times, even more puzzling no apparent core (reviewed by Menzies, 1979, 1989b). Recent satellite imagery clearly demonstrates the apparent close relationship between drumlin fields and ice stream locations (Stokes and Clark, 2001; Clark and Stokes, 2003; Clark et al., 2003, 2009; Greenwood and Clark, 2010; Hooke and Medford, 2013; Margold et al., 2015; Hillier et al., 2016). Several current hypotheses exist as to the formation of drumlins and drumlin fields. In all cases, an ‘event’ or a ‘trigger’ appears to be necessary for their formation and development. Once an initial nucleation occurs (stage 1) possibly due to a ‘sticky spot’ developing at the lower interface 130 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

FIGURE 5.16 (A) Drumlins, Clyde County, New York State, USA. Note these are elongated drumlins either side of a broad interdrumlin area. (B) Drumlins in northern Finland. Courtesy of Risto Aario.

(Alley, 1993; Sergienko and Hulbe, 2011; Luthra et al., 2016; Roeoesli et al., 2016; Walker et al., 2016) (sticky spots being small areas of higher basal shear stress or areas of localized higher shear strength sediments or bedrock protrusions), it can be demonstrated that in some places the form will persist, grow and possibly migrate (stage 2) while elsewhere the form after stage 1 is demolished (Menzies et al., 2016). The problem with all drumlin formative hypotheses is the question of what is the ‘trigger’? Following the ‘trigger’, a relationship between sediment added (basal sediment flux) to the original nuclei, erosional processes on top of and between drumlin needs to be considered (McCracken et al., 2015; Vreeland et al., 2015). Currently, three broadly acceptable hypotheses exist that attempt to account for the formation and subsequent development of drumlins: 1. Deforming sediment bed accumulating a proto-nuclei around a bed ‘sticky spot’ (stage 1) that, in some cases, conditions permit survival (stage 2) (Savage, 1968; Boulton, 1987; Menzies, 1989a,b; Smith et al., 2007a,b; Luthra et al., 2016; Menzies et al., 2016; Walker et al., 2016). 2. Groove ‘ploughing’ in which an uneven topographic expression (stage 1) is further ‘sharpened’ into stage 2 forms (Tulaczyk et al., 2001; Clark et al., 2003). 5.6 SUBGLACIAL LANDFORMS/BEDFORMS DIRECTLY ATTRIBUTABLE 131

3. Interface instability (stage 1) in which bed topography is locally formed into positives and negatives; the former developing at stage 2 into drumlins (Hindmarsh et al., 1989; Hindmarsh, 1999; King et al., 2007; Schoof, 2007; Fowler, 2010a,b).

5.6.2 FLUTED MORAINES AND MEGASCALE GLACIAL LINEATIONS (MSGLS) There would appear to be an association of like-bedforms developed at subglacial interfaces between fluted moraines, MSGLs, drumlins and Rogen moraines (Rose, 1989; Clark, 1993, 1994; Roberson et al., 2011; O´ Cofaigh et al., 2013; Stokes et al., 2013a,b; Trommelen and Ross, 2014; Spagnolo et al., 2016). Hypotheses like those used to explain drumlin origin have been suggested for fluted moraine, except that, unlike drumlins, flutes can extend for many kilometres, are much narrower, in general, and typically lower in height. Flutes of fluted moraines, typically, are between 1 and 10 m in height and can extend from a few decimetres to several kilometres in length (Fig. 5.17). MSGLs, on the other hand, can be several metres to over 10 m in height and extend for many kilometres. Both forms are generally composed of a melange´ of sediments that have appar- ently been collectively basally ‘scavenged’ by forward basal ice motion (either directly at the iceÀbed interface or between the interface of mobile and immobile sediments within a deforming bed) in the process of formation (Smith et al., 2007b). The old hypothesis that many flutes grow, or at least begin developing in the lee of boulders (Fig. 5.18) but in many cases, and especially so with MSGLs, the cause of nucleation streamlined elongation is often less apparent. It seems likely that the formative processes involved in both drumlins and MSGLs are very similar and any differentiation may be, in the case of flutes and MSGLs, more the result of a relatively narrowly confined, high-sediment flux rate at the iceÀbed interface and the relatively high basal ice velocities as is typical for ice stream locations (Dunlop and Clark, 2006; Dunlop et al., 2008; Stokes et al., 2008; Greenwood and Clark, 2010; Winsborrow et al., 2010; Menzies et al., 2016).

5.6.3 ROGEN MORAINE These morainal forms, also termed ribbed, are formed transverse to the dominant ice flow directions. It was probably Lundqvist (1989) who first suggested that rather than considering Rogen moraine in isolation or as unique landforms, the moraines were part of a continuum of forms either emanating from or passing into parallel fluted moraine and drumlins (Fig. 5.16). In some instances, Rogen moraine and drumlin exist almost side by side (Aylsworth and Shilts, 1989). Rogen moraine occur as ridges 6 10 m in height stretching transverse to ice flow for hundreds of metres to several kilometres and may pass down- or up-ice into drumlinized terrains (Lundqvist, 1970; Menzies, 1987) and vice versa (Wagner, 2014, and references therein). Like fluted moraine, MSGLs and drumlins, the sediment content of these ridges is varied and essentially a melange´ of basal available sediment. The surfaces of these moraines are often covered in boulders of varying sizes (Aylsworth and Shilts, 1989, p. 421; Shilts, 1976, 1982). Hypotheses of Rogen moraine origin are like drumlins in that both are a subglacial interface ‘problem’. However, these moraines can be viewed as wave-like forms that may result from rapid transverse glaciodynamic responses, such as proximal grounding-line lift-off events or instabilities inherent in the sediment flux, basal ice stress patterns or subglacial sediment flux rate variabilities or possibly basal ‘crevasse-squeeze’ forms (Bouchard, 1989; Lundqvist, 1989; Fisher and Shaw, 1992; Ha¨ttestrand, 1997; Knight and McCabe, 1997; Schoof, 2002; Dunlop and Clark, 2006; 132 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

FIGURE 5.17 (A) Fluted end moraine at Slettjokull, Iceland. (B) Fluted moraine within the ‘Cabot Ice Stream’ in an area east of the Labrador Ice Centre, north of the Smallwood Reservoir, Newfoundland and Labrador, Canada. Note the esker cutting through the flutes in the centre of the photo beyond the algae-covered lake. (C) Fluted moraine south of Pine Point, NWT, Canada. Older, larger northwest trending landforms were cannibalized to form these westÀsouthwest trending flutes during subsequent ice streaming. These flutes remarkably maintained their morphology, despite being winnowed by Glacial Lake McConnell. Great Slave Lake is to the north. Courtesy of Roger Paulen. 5.6 SUBGLACIAL LANDFORMS/BEDFORMS DIRECTLY ATTRIBUTABLE 133

FIGURE 5.18 Flute developed in the lee of a boulder on the forefield of Storbreen Glacier, Jostedal, Norway. The boulder is approximately 1.5 m in height.

Angelis and Kleman, 2008; Dunlop et al., 2008; Stokes et al., 2008, 2013a,b; Ankerstjerne et al., 2015; Menzies et al., 2016). Many questions remain as to the origin of Rogen moraine, for example, their tendency to be found close to ice divides in Canada where basal ice motion must have been limited and where the moraines take on an almost ‘thrust plate’ appearance (cf. Aylsworth and Shilts, 1989b; Shilts, 1976, 1982; Stokes and Clark, 2003)(Fig. 5.19aÀd).

5.6.4 MARGINAL MORAINES Other moraines that form transverse to the ice flow direction, but not formed subglacially, may be the result of push from advancing ice or the upward squeezing of subglacial sediments at ice margins or accumulate at the ice frontal margin as end, recessional or terminal moraines (Bennett and Boulton, 1993; Kru¨ger, 1995, 1996; Bennett, 2001; Evans and Hiemstra, 2005; Benediktsson et al., 2008, 2016). Such moraines can vary in height from a few metres to several tens of metres and commonly have an asymmetric transverse profile (Fig. 5.20). These moraines often contain a range of sediment types from various subenvironments picked up by advancing ice. For instance, the marginal moraines in Illinois and most of the mid-United States are largely, but not always, tills (cf. Webb et al., 2012), whereas those in Arctic Canada may contain stratified and unstratified sediments, deltaic and lacustrine sediments. Where the clay content of the moraine is sufficiently high, the moraines may attain an almost vertical slope profile. The volume of any marginal moraine is very much a function of the residency time the ice margin is at or close to a specific location. In many instances, the ice margin may return to the same location within the terrain, thus continuing to build up the moraine repeatedly over time (Kru¨ger, 1995; Bennett, 2001; Kru¨ger et al., 2002; Stiff and Hansel, 2004; Vacco et al., 2009; Curry et al., 2011; Eyles et al., 2011). The sediment content of most marginal moraines reflects a wide diversity of sediment: supraglacial, englacial and subglacial. FIGURE 5.19 (A) Rogen (ribbed) moraines south of Ennadai Lake depicted using various multispectral band combinations. All Landsat ternary images have been pan-sharpened using ETM 1 panchromatic (15 m resolution) band 8. (a) Landsat 7 ETM 1 4,3,2 (RGB) FCC of visible and near-infrared bands. (b) Landsat 7 ETM 1 5,3,2 (RGB) FCC of visible and mid-infrared bands. (c) Landsat 7 ETM 1 7,4,2 (RGB) FCC of infrared, near-infrared and visible bands. (d) SPOT 5 panchromatic image (10 m resolution). (B) Examples of Rogen confined with an area of low topography near South Henik Lake, Nunavut, Canada. (C) Rogen moraines near Whitbourne, the Avalon Peninsula, Newfoundland, Canada. (D) Rogen moraines near Lough Oughter, Ireland. (A, B) Courtesy of Wagner, K. 2014. Ribbed Moraines and Subglacial Geomorphological Signatures of Interior-Sector Palaeo-Ice Sheet Dynamics (M.Sc. thesis). Brock University, St. Catharines, Ontario, 274 pp. (C) Courtesy of Tom Fisher. FIGURE 5.19 (Continued)

FIGURE 5.20 Marginal moraines along the northern edge of the Findelen Glacier, Switzerland. 136 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

Hummocky moraine is a distinctive form of moraine that occurs over large areas of submarginal and proximal proglacial areas. Past depositional models argue that hummocky moraines were deposited supraglacially from stagnant debris-rich ice (‘disintegration moraine’) (Bennett and Boulton, 1993; Eyles et al., 1999; Ham and Attig, 1996; Munro-Stasiuk, 1999; Moller¨ and Dowling, 2015)(Fig. 5.21AÀC). Boone and Eyles (2001), in contrast, suggest that hummocky moraine may be a product of subglacial erosion rather than supraglacial letdown during ice

FIGURE 5.21 (A) Hummocky moraine Kap James Hill, Greenland. (B) Landsat ETMþ satellite image of upper Glen Turret (red dot and arrow marks the location of photograph in (C)). (C) Hummocky moraine in Glen Turret, Scotland, of Loch Lomond Readvance age. Moraine in foreground is approximately 8À9 m in height. 5.7 SUBGLACIAL LANDFORMS/BEDFORMS INDIRECTLY ATTRIBUTABLE 137

disintegration. Eyles et al. (1999) note that across southern Alberta hummocky moraine is composed of fine-grained till as much as 25 m thick, containing rafts of soft, glacitectonized bedrock and sediment. Much hummocky moraine is chaotic and nonoriented that appears, in places, to pass downslope into weakly oriented hummocks (Rogen moraine) that are transitional to drumlins in topographic lows. Ice disintegration, subglacial ice ‘pressing’, and possibly subsequent chaotic overprinting on drumlins, fluted moraines and Rogen moraine may all help explain hummocky moraines in different parts of the world. In some places, behind these hummocks can be found ice- walled lake sediments, thermokarst features and, at times, associations with surge events resulting in ice/sediment passive disintegration (cf. Ham and Attig, 1996; Curry and Yansa, 2004).

5.7 SUBGLACIAL LANDFORMS/BEDFORMS INDIRECTLY ATTRIBUTABLE TO ACTIVE/PASSIVE ICE ACTIVITY Landforms/bedforms within this category of subglacial landform are largely attributable to the influence or impact of meltwater activity (Table 5.4). A subdivision of these forms can be made based on those formed subglacially, and others developed either submarginally or marginally to an ice mass. Although these forms may grade into other topographic features, they tend not to be part of a continuum of bedforms but typically occur together within the same area of terrain. For example, where one finds eskers, it is not unusual to find kames, kettle holes and kame terraces (Fig. 5.22D).

5.7.1 ESKER SYSTEMS Eskers are straight to sinuous ridges of glaciofluvial gravel and/or sand that vary from a few tens of metres to over 100 km in unbroken length (Fig. 5.22A, B). The ridges locally range from a few metres to over 50 m in height and from less than 50 m to hundreds of metres in width at their base. Eskers may be transitional to large ice-contact stratified drift complexes deposited by interlobate meltwater streams, and this is reflected by ambiguous references to the origin of large, linear glaciofluvial complexes in the literature (Brennand and Sharpe, 1993; Clark and Walder, 1994; Warren and Ashley, 1994; Ashley and Warren, 1997). Esker ridges generally trend parallel to the final direction of ice flow in the area where they are preserved, suggesting that although they may be formed by meltwater drainage throughout a glaciation, only those formed during the latest stages of glaciation are preserved (Fig. 5.22C). Notwithstanding the above observation, eskers can be seen to cross ice flow features, such as drumlins, at sharp angles, suggesting that regional ice flow was, at best, sluggish during their formation (Fig. 5.22D). Eskers tend to occupy the lower portions of a landscape. Where esker ridges pass from one depression or valley to another, they are draped, with no apparent deformation, over divides or cols. Observations of modern glaciers show them to be honeycombed by tunnels, passages, crevasses, etc., conduits that pass through the ice mass in various orientations from vertical to subhorizontal (Banwell et al., 2016; Dowdeswell and Ottesen, 2016). During the melt season, in the ablation zone, these conduits are commonly full of water, and those at or near the base of the glacier are under a hydrostatic head (Greenwood et al., 2016). Meltwater in conduits that reach or pass near the base of the glacier become charged with mud FIGURE 5.22 (A)This image, taken from an altitude of about 3000 m, shows an esker ridge superimposed at almost right angles on a drumlin. (a) Esker at De Pas, Labrador, Canada, north of the Smallwood Reservoir with the George River in the background. (b) Esker bifurcating into several linked ridges in the distance, De Pas, Labrador, Canada, north of the Smallwood Reservoir with the George River in the middle ground. (c) Esker ridges in the Hand Hills in southern Alberta, Canada. (d) Kame terrace at Bardebreen, a tributary glacier to Sefstrombreen in Spitsbergen. (B) This image, taken from an altitude of about 3000 m, shows an esker ridge superimposed at almost right angles on a drumlin. The bouldery ridge in the background is the inner face of the substantial lateral moraine surrounding the rapidly retreating Stagnation Glacier (B28). The landform is composed mostly of glacial ice preserved by the insulating cover of a metre, more or less, of bouldery esker gravel. (Photograph taken by Christian Zdanovich). (C) Details of the Rennie lake esker system showing the dendritic pattern of the esker tributaries; (D) Esker ridge crossing a drumlin obliquely between Turquetil Lake, Keewatin (Nunavut) and Lynn Lake, Manitoba. (A, B) Courtesy of PRI, WW Shilts Image Gallery. (a, b) Courtesy of Roger Paulen; (c) courtesy of Tracy Brennand. (C) Adapted from Shilts, W.W., Aylsworth, J.M., Kaszycki, C.A., Klassen, R.A. 1987. Canadian Shield. Geomorphic Systems of North America, vol. 2. Geological Society of America, Boulder, Colorado, pp. 119À161. (D) Courtesy of PRI, WW Shilts Image Gallery. 5.7 SUBGLACIAL LANDFORMS/BEDFORMS INDIRECTLY ATTRIBUTABLE 139

FIGURE 5.22 (Continued)

melted out of the glacier’s debris-rich basal layers. The finer portions of the mud are carried through and out the mouth of the conduit into proglacial streams or directly into lake or marine water, depending on whether the glacier is retreating in the sea, up gradient (allowing unimpeded proglacial drainage), or down gradient (blocking drainage and forming proglacial lakes). The coarser components are largely trapped in the conduit, partially or completely filling it, and are left as sandy or gravelly esker ridges when the glacier melts away (cf. Ma¨kinen and Palmu, 2008; Rampton and Sharpe, 2014). Since the till adjacent to eskers represents the total basal load of the glacier that deposited the esker, and since individual till sheet thicknesses rarely average more than a few metres, it is evident 140 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

that some process(es) enhance sediment accumulation in esker conduits. Since meltwater flowing through glacial conduits is generally at a temperature above the pressure melting point of the ice in which the conduits formed, and since frictional heat is generated by the fast-flowing water-sediment suspension, tunnel walls are subjected to melting, thus enlarging the meltwater cavity. As ice flows toward the conduit to maintain a stable cross-section, it brings basal debris with it so that new material from lateral sources is constantly being dumped into the conduit, sorted, and deposited in or carried out of the conduit to be deposited ultimately in proglacial outwash streams, lakes or marine waters. By this process, the esker grows to be orders of magnitude thicker than it would be if it were just sweeping out the debris contained in basal ice into which the tunnel was excavated. Esker systems form as a function of the location and form of a meltwater channel or drainage system either subglacially, englacially or at least submarginally along the outer edge of an ice mass, yet our understanding of subglacial and englacial meltwater systems remains limited (Veillette, 1986; Shilts et al., 1987; Aylsworth and Shilts, 1989a; Clark and Walder, 1994; Clarke et al., 1984; Warren and Ashley, 1994; Kleman and Ha¨ttestrand, 1999; Brennand, 2000; Boulton et al., 2001, 2009; Bennett et al., 2009; Cummings et al., 2011; Burke et al., 2012; Greenwood et al., 2016; Perkins et al., 2016). Essentially, an esker is the result of a sediment- choked meltwater channel (Boulton et al., 2007a,b, 2009; Hooke and Fastook, 2007). Some eskers would appear to form within part or completely within an englacial conduit that, on ice melting, overlies subglacial sediments. In some instances, eskers drape drumlins often at an oblique angle to the drumlin long-axis orientation (Aylsworth and Shilts, 1989b; Salt and Evans, 2004)(Fig. 5.22D). Since eskers mirror the pathways taken by meltwater channels and drainage systems, they generally do not follow directly the same azimuth as the ice flow lines within an ice mass. Esker ridges (Fig. 5.22A) can be a few metres to several tens of metres in height and may run across terrain for a few tens of metres to many hundreds of kilometres (Margold and Jansson, 2012). Examples from the Northwest Territories and Nunavut, Canada, illustrate the location and distribution of eskers on the Canadian Shield and below the central area of the Laurentide Ice Sheet during the Late Wisconsinan (Menzies and Shilts, 2002, Fig. 8.44; Margold et al., 2013; Storrar et al., 2013) (Fig. 5.22C). It seems that the location and size of eskers and entire esker patterns, in some instances, may be related to groundwater sources below ice sheets in which the transmissivity of the bed exerts control (Boulton et al., 2007a,b, 2009). In other cases, eskers may trace out the surface drainage supraglacial patterns that subsequently descend to the ice base. In many cases a dendritic pattern emerges (cf. Shilts et al., 1987)(Fig. 5.22C). In other instances, esker patterns can be deduced from basal meltwater recharge rates coupled with patterns of palaeo-groundwater flow and the seasonally varying magnitude of discharge from stream tunnels at the retreating ice sheet margin (Boulton et al., 2009; Hewitt, 2011). Major channel/esker systems appear to form under quasistable conditions close to the ice margin, at least over several centuries, during the retreat of an ice sheet (Hooke and Fastook, 2007; Shilts et al., 2010). The development of esker systems would appear to be interlinked with hydraulic systems supraglacially, englacially, subglacially and possibly, within the coupled underlying groundwater systems (Rothlisberger¨ and Lang, 1987; Boulton and Caban, 1995; Boulton and Zatsepin, 2006). Esker system patterns indicate, likely, ice margin configurations of the former ice sheet retreat phases during deglaciation. Large trunk eskers characteristically form perpendicular to ice margins. Esker distributary systems give clues to understanding palaeo-glaciology and how large subglacial meltwater drainage systems operate under present ice sheets. 5.8 REPETITIVE SEDIMENTOLOGIC HISTORIES 141

5.7.2 TILL DELTAS/TONGUES AND GROUNDING-LINES Where ice masses enter large bodies of water and begin to float at a grounding-line, a large tongue of till may develop that has been variously termed a till delta or a till tongue (King et al., 1991; King, 1996; Shaw et al., 2006; Anandakrishnan et al., 2007; Graham et al., 2009; Larter et al., 2009; Smith and Anderson, 2010)(Fig. 5.1C). Under conditions of subglacial soft sediment deformation, it seems likely, based on research in Antarctica and on Pleistocene ice sheets, that the till and subglacial melange´ sediments may emerge as deforming units into the water body, thus developing a wedge of till out into the bed of the water body. In other cases, where subglacial melt tunnels emerge into water bodies, substantial subaqueous fans and deltas may form in these ice-proximal environments with esker systems, in some cases, extending back up ice from the fans and deltas (Ma¨kinen and Palmu, 2008; Ahokangas and Ma¨kinen, 2014; Greenwood et al., 2016). Likewise, under these soft bed conditions, as an ice mass retreats, a till tongue may slowly begin to form beneath the subglacial margin at or very close to the grounding-line. At a grounding-line, there are significant changes in subglacial stress and hydraulic conditions that will lead to rapid changes in sediment rheology at that point. In many instances, major meltwater portals emerge at the grounding-line producing large marginal glaciofluvial deposits as subaqueous fans and deltas. These major changes at the grounding-line in terms of stress and hydraulics are transmitted back up-ice and affect subglacial conditions for some considerable distance back under the ice. The impact of such grounding-lines is only now being investigated but already ideas regarding subglacial soft sediment deformation, and how this bears upon drumlins and MSGL development, require investigation in the field (Le Meur and Hindmarsh, 2001; Schoof, 2007). A further aspect of grounding-line dynamics is the potential effect on ice sheet stability that is triggered or otherwise by the stability and location of these grounding-lines (Conway et al., 1999; Rignot et al., 2004; Vieli and Payne, 2005; Pollard and DeConto, 2009; Joughin et al., 2014).

5.8 REPETITIVE SEDIMENTOLOGIC HISTORIES IN SUBGLACIAL ENVIRONMENTS It is commonplace in glacial environments that recurring events involving erosion, transport and deposition occur repeatedly at the same site (Menzies and Shilts 2002). A result of reactivation, readvance following retreat partially or fully from an area is that it is impossible to predict the subglacial surfaces in place following final retreat of the ice. In subglacial environments, there are two levels of repetitive event histories that can be distinguished: (1) those occurring during a single phase of glacial activity and (2) those during each successive phase of glaciation. Too often subglacial environments are discussed with a somewhat static view of sedimentological activity, when in fact the subglacial environment is one of repeated dynamic activity in which multiple events occur with such large variations in time and space that differentiation of each event history is an intricate and, at times, almost impossible task (Boulton and Deynoux, 1981; Larsen et al., 2004; Evans and Benn, 2014). During single glacial events, sediments can pass through many successive sequences of erosion, transport and deposition. The ambiguous term ‘resedimentation’ has been extensively used in attempts to convey a sense of the active nonstatic sedimentary systems within subglacial 142 CHAPTER 5 SUBGLACIAL PROCESSES AND SEDIMENTS

environments. Landforms and bedforms may be formed and overprint earlier forms and sediment packages, or they may be partially or wholly destroyed or unaffected, depending upon time of formation, sediment shear strength, effective stress levels, etc. In some cases, the change of one bedform to another may be only partially complete, while in other locations a new form emerges replacing the previous form with no vestige of the earlier form visible. However, sediments, their characteristics, structures and properties, may still retain the imprint of past sedimentological histories that might indicate frozen conditions, rafting, thrusting, etc. Overprinting, reorientation, reworking, realignment and redirection are all terms indicative of episodes of repetitive event histories. Where deglaciation has occurred and subsequent ice re-advance has occupied an area once again, a quite different set of circumstances must be considered. First, the length of time involved between separate advances. Second, the impact of climatic changes on soil development, vegetation colonization and the effect of subaerial processes in altering topographic features and slopes. Third, the presence of postglacial lakes or sea-level rise in flooding areas of terrain. Finally, however long land surfaces have been exposed, pre-existing subglacial sediments and landforms may be drastically or only slightly altered. In many instances, all records of previous subglacial sediments may be totally removed and reincorporated into a new subglacial environment. It is possible that no sediment units may survive even as rafted inclusions. However, geochemical signatures may be detectable in the form of altered minerals and neoformed clay minerals. In some cases, the presence of organics in the form of woody materials, pollen or bone fragments may be indicative of past warmer climate conditions. However, such evidence is commonly lost through reworking by overriding glacial erosion. Some artefacts cannot be so easily destroyed (striae, chattermarks, conchoi- dal fractures) but instead cross-cutting and overprinting relationships may be detected. It is critical that field researchers be aware of the influence and extent of overprinting and multiple event histories. The overlaying effect of multiple sedimentary events within one or multiple glacial phases must be considered whenever the origin of subglacial sediments and forms are debated and when glacial stratigraphy is examined (cf. Shilts et al., 2007; Shilts and Caron, 2017).

5.9 FUTURE PERSPECTIVES: CHALLENGES AND OPPORTUNITIES Subglacial sediments in terrigenous glaciated areas of the world are crucial in understanding the diverse use of these sediments, their geotechnical and geochemical aspects, amongst so many attributes of these sediments. It is critical that a clear understanding and grasp of subglacial lithofacies, bedforms and lithofacies associations be fully appreciated. Volumetrically, other than glaciomarine sediments, these sediments cover large areas of the Earth’s glaciated surface. A full appreciation of their influence on human activities whether, for example, in agriculture, construction as foundations and aggregate resources, drift prospecting, in utilizing these sediments as sites for waste, toxic and nuclear disposal, or as aquifers of major groundwater sources for large urban areas is needed. There remains a considerable wealth of knowledge to be gleaned from these sediments and bedforms. As our understanding of subglacial processes from past glacial environments increases, so new avenues of research emerge. Likewise, paradigms shift and mutate such that what has been accepted understanding is now and remains in some cases in a state of flux. For example, the analyses of microstructures within diamictons has opened a large area of research that demands new perspectives on how all diamictons are deposited and what conditions prevail in subglacial environments and how these environments fluctuate in terms of temperature, porewater and meltwater content and under varying strain conditions. Similarly, there is REFERENCES 143

an increasing need to grasp the intricate interplay between thermal, geological, glaciological and rheological conditions that coexist at subglacial interfaces of which subglacial sediments are the end product. Interpretation of subglacial facies associations, subfacies units and the diagnostic criteria that might be used to discriminate environments, processes and forms remain productive research areas. The utilization of satellite and Lidar, among many forms of remote imagery, provide challenges in terms of interpretations but also an array of new and exciting associations among subglacial bedforms, landform associations and the glaciodynamics of ice sheets, especially ice stream activity. It is essential that relationships be understood beneath past ice sheets in terms of depositional environments and ice stream activity. Although a great deal is known about subglacial processes, there remains limited understanding concerning rates of processes, subglacial sediment flux rates and the changing rheological states within underlying deforming and undeforming sediments. It is likely, that in many cases sediments and attendant macro- and microstructures may be deposited and generated over very short periods of a few hours or days, while others such as melt-out may take centuries. Since past subglacial sediments are uti- lized today for innumerable purposes, it is vital that the effects of postdepositional diagenesis leading to pedogenesis be much better understood (cf. Shilts and Kettles, 1990; McClenaghan et al., 2000; Ayuso et al., 2005; Dempster et al., 2013). Diagenesis appears to begin almost immediately after sediments are deposited even before ice retreat and then it accelerates, especially in the first years of subaerial exposure (Stankowski, 1976; Sna¨ll, 1985; Byrne, 1994; Krklec et al., 2015). Since a large percentage of glacial sediments have not gone through chemical winnowing processes, these sediments contain active mineral surfaces subject to postdepositional chemical weathering, where labile components are easily altered. However, the extent, depth and degree of alteration caused by diagenetic and pedogenic processes remains an area of much needed new research. It would be fair to state that our understanding of past ice sheets, the dynamics of their growth and decay and the interrelationship of ice sheets with oceans and lakes is revealed by our somewhat restricted grasp of many of the processes in subglacial environments. Over the past 20 years there has been a tendency to view research findings from the specific to the general, partly due to the costs involved in extensive field exploration in remote areas; however, it is worth pointing out that the con- verse view from the general to the particular must not be lost in future studies of these environments. However, major headway in this aspect of glacial environments is being made in the North Sea, northern coastal Norway and in the Barents Sea and off Svalbard in Europe and in the offshore around Antarctica. The ‘drumlin enigma’ stands as a conspicuous instance of how much there is still to understand about ice sheet dynamics and the interplay of glacier motion, sediments, topography, source rock composition and properties, and the environmental ‘conditions’ at the base of ice sheets and valley glaciers. Until we fully understand the complexities of environments beneath ice sheets and other ice masses, and spatial distributions of glacial sediments and sediment thicknesses, it is unlikely we can appropriately postulate the explanation(s) necessary for subglacial landform development, sediment deposition and predict other geomorphic processes at the iceÀbed interface.

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