The Sedimentary Record of Large Proglacial Lakes

14.3.1 Introduction

As glaciers melted, they relinquished their load of sediment. The hydrology and sediment transport in glacial rivers and lakes have been studied by many (e.g. Church and Gilbert, 1975;

(A) Southern Outlet Control

(B) Middle Outlet Control

Level 3

Level 4

(C) Northern Outlet Control

Level 3

Level 4

(B) Middle Outlet Control

(C) Northern Outlet Control

Figure 14.2 Examples of water level changes in a basin (levels 1, 2, 3, and 4 in bathtubs) undergoing differential isostatic rebound, where the outlet is in the southern end (S), northern end (N), and in the middle of the basin. A) Outlet at southern end. Lake regresses throughout basin and blanket of sand may be deposited. B) Outlet in middle of basin. Lake regresses to north of outlet and transgresses to south. Blanket of sand deposited in north, but lake and beach transgress upslope in southern end; eventually beach stranded at maximum (level 4) when lake level drops. C) Outlet in northern end of basin; lake level transgresses everywhere, moving beach upslope to maximum. (After Larsen, 1987; Teller, 2001).

Figure 14.3 Example of more complex sequence of lake level development after southern outlet shown in Figure 14.2 A is abandoned (after Teller, 2003). A) Lake abruptly drops to level 4 after new outlet in middle of basin (Ml) becomes ice-free; no large beaches formed. B) Overflow through outlet Ml produces transgression south of outlet, regression to north; beach moves upslope in front of transgression. C) New outlet in middle of basin opens (M2) and lake abruptly drops from level 6 to 7, stranding beach at the transgressive maximum of level 6. D) Lake transgresses from level 7 to 8, where it begins to overflow again through outlet at southern end of basin, stranding beach at transgressive maximum of level 8.

Gustavson, 1975; Ostrem, 1976), as has subaqueous deposition at the glacier margin and in deeper parts of proglacial lakes (e.g. Jopling and McDonald, 1975). Studies indicate that sediment loads are high in modern meltwater rivers (e.g. Maizels, 1995) and, in fact, Friedman and Sanders (1978, p. 25) show that the highest rate of sediment discharge to the oceans from any continent is from Antarctica. In the glacimarine environment ofAlaskan fjords, Powell (1983) and Cowan and

Powell (1991b) measured sediment accumulation rates at an astonishing 9—13 m year-1, and high sedimentation rates occur in lakes near the mouths of glacial rivers where subaqueous density currents and slumps deliver large amounts of sediment (Smith and Ashley, 1985). Away from the point of sediment influx, accumulation rates are much less, and the varved sedimentary record of many large Pleistocene proglacial basins indicates that offshore (distal) accumulation was rarely more than a few centimetres per year (e.g. Antevs, 1925, 1951; Teller and Mahnic, 1988; Ringberg, 1991; Hang, 1997). Overall, glacilacustrine sediment thins away from the ice margin and the mouths of rivers.

The rate of sediment released from a glacier is a function of several factors, among them rate of melting, type of glacier and concentration of debris in the ice. The latter is partly a function of the availability and erodibility of bedrock and the overlying unconsolidated material, so lakes in areas of resistant lithologies such as granite, which are not as easily weathered and eroded as are sedimentary strata, tend to receive relatively small influxes of sediment that contain low percentages of clay-sized material.

Ice-marginal lakes have higher sedimentation rates than do their downstream proglacial counterparts, and the nature of the sedimentary record is different. Abrupt jumps in sedimentation rate or style, grain size, mineralogy, and biological composition will occur when the ice retreats from (or advances into) a lake basin. Progressive changes in these parameters will occur as an ice margin advances or retreats across the lake basin, and these changes can be used to interpret the location of the margin in the basin through time.

Morphologically, most proglacial lake deposits tend to form flat landscapes of clay and silt that drape over (and may ultimately obscure) previous features in the offshore region of the lake. The flat offshore areas of lake sediment typically are outlined by wave-cut cliffs or by coarser sediment deposited in shallow water, such as beaches and lags of wave-washed sediment.

The focus of the following sections will be on those sediments and landforms in large ice-marginal lakes dammed by continental ice sheets. In many ways, processes in these lakes are the same as those in non-glacial environments. However, there are some distinctive (even unique) aspects about the sediments that relate to the presence of the bounding (or nearby) ice margin and its highly dynamic nature, the abruptness of water inflow and outflow, iceberg calving, and isostatic rebound. Unlike many other lake settings, transgressions and regressions are common during the life of proglacial lakes because of the frequent and often large changes in ice margin, outlet depth and location, runoff into the lake, and differential isostatic rebound. As a result of erosion during transgression and subaerial exposure, as well as glacial readvance, unconformities commonly are present in the sequence, particularly in shallower parts of the basin.

14.3.2 Surface Appearance

Commonly, the surface of a large proglacial lake basin is flat, partly because continental ice has smoothed the preglacial surface and partly because wave action and sediment accumulation have further reduced the relief. Relatively impermeable sediments and low relief in many large palaeo-lake basins have resulted in poor drainage; up to 4 m of peat has accumulated in places on the floor of these ancient lakes (e.g. Dredge and Cowan, 1989). In Canada, many of these areas today are covered by muskeg, and the Hudson Bay Lowland and adjacent Canadian Shield support the largest area of organic terrain in the world (Dredge and Cowan, 1989).

The keels or edges of icebergs, and even winter ice, sculpted long linear and curved grooves in the mud on the floor of many lakes, with associated ridges immediately adjacent to them (Fig. 14.4) (see Woodworth-Lynas, 1996 for summary and discriminating characteristics), and these features are found over extensive areas of proglacial lake basins in North America as well as in modern lakes (e.g. Clayton et al., 1965; Dredge, 1982; Woodworth-Lynas and Guigné, 1990; Gilbert et al., 1992). These straight-to-curved features criss-cross each other. In oceans, iceberg scours have been found in water depths of 500-950 m (Barnes and Lien, 1988; Polyak et al., 2001). Sediment homogenization and faulting are commonly associated with these scour and plough marks. Although they are very subtle landforms, rarely having more than a metre or two of relief today, they are striking features from the air on an otherwise featureless offshore lake plain (Fig. 14.4). Tonal contrast results from the difference in organic accumulation in the soil between the better-drained ridges and adjacent lower areas. These ridges commonly lie adjacent to furrows and were interpreted by Clayton et al. (1965) as iceberg scour grooves and ploughed ridges, which are today more subdued than when originally formed. Woodworth-Lynas and Guigné (1990) argue that

Figure 14.4 Aerial photo of iceberg scour marks in the silty clay of the Lake Agassiz plain east of Winnipeg. Light linear tones are slightly elevated ridges. Main roads are spaced 1 mile (1.6 km) apart, and west to east distance across photo is about 5.5 miles; north is toward top.

these ridges were once the iceberg grooves which have been topographically inverted due to greater compaction of sediment on either side.

In addition to these extensive curvilinear scour features, Mollard (1983, 2000) describes and illustrates a variety of other irregular patterns that developed on lake floors in glacial regimes (Fig. 14.5), proposing various origins for them. The clayey and silty sediments in these features commonly exhibit convolutions, diapiric structures, and deformation (Mollard, 1983). All have very low relief. For example, as lake levels decline in an ice-marginal lake, icebergs may 'settle into the soft lake-bottom mud', resulting in folded and faulted silts and clays and a shallow depression that ranges from a few tens of metres to a few hundred metres in diameter (see Fig. 14.6). These may be rimmed by coarser drift that slipped off the ice mass, or by detritus that ablated into the depression (Thomas, 1984b; Mollard, 2000). Some of the low-relief 'doughnut' and brain-like patterns seen on aerial photographs in glacial terrains may be the result of this process (Fig. 14.5), although similar features are produced outside of lake basins in stagnant ice regions (Clayton and Freers, 1967; Mollard, 1983, 2000). Shallowing lake waters may also have resulted in the differential freezing (expansion) and thawing of lake muds, producing small pingo-like mounds that occasionally have central craters (Mollard, 2000). Even groundwater 'piping' from artesian

Jean Dubuffet
Figure 14.5 Brain-like pattern of deformed lacustrine silty clay, interpreted to have resulted from floating ice sinking into the soft mud (Mollard, 2000, fig. 6). The change to a less-distinct pattern with smaller rings in the lower right is interpreted to reflect deeper water.

flow from aquifers below the lacustrine muds may develop discharge pits and surrounding raised rims (Mollard, 1983).

Of course the morphology of a lake basin also reflects the underlying topography of the basin floor, which may be controlled by bedrock or an older erosional or depositional surface. Ice-marginal landforms deposited on the lake, such as end moraines, deltas and fans, may remain as distinct features on the lake floor, although they are vulnerable to wave erosion. A drape ofyounger lacustrine sediment may further diminish the relief, leaving relict (palimpsest) landforms.

14.3.3 Offshore Sediments

14.3.3.1 Iceberg Contribution

Icebergs may raft detritus of all sizes into ice-marginal lakes, releasing their load as they melt, tip or break up, and adding it in varying proportions to sediment on the floor of the lake (Fig. 14.6); this rain of detritus commonly deforms or even obliterates pre-existing laminae (Fig. 14.7), and consists of rock fragments, individual mineral grains, and fragments of unconsolidated sediment such as till (Ovenshine, 1970). The concentration of this ice-rafted detritus may be so great that the resulting sediment may resemble till (e.g. Dreimanis, 1979; see also Benn and Evans, 1998). In general, higher concentrations of this detritus are found close to the ice margin, beneath floating ice shelves, or where icebergs become grounded in shallow water and decay. In addition, sediment already deposited may be highly disturbed by icebergs ploughing into the soft sediment, producing an ice-keel turbate that may be similar in appearance to till (Vorren et al., 1983; Barnes and Lien, 1988; Woodworth-Lynas and Guigné, 1990).

14.3.3.2 Fine-Grained Offshore Accumulation

Most sediment entering a lake from rivers beyond the ice margin will initially be reworked by waves. Coarser sediment will be deposited in shallower waters around the lake margins (see section

Englacial Tunnel Image

Figure 14.6 Construction of a fan delta (A) into a lake at the mouth of an englacial tunnel (commonly referred to as a grounding-line fan), showing more distal lacustrine facies (B and C) and additions of unsorted detritus from glacier overhang and icebergs. Note grounded iceberg and resultant depression. (From Thomas, 1984b, fig. 5).

Figure 14.6 Construction of a fan delta (A) into a lake at the mouth of an englacial tunnel (commonly referred to as a grounding-line fan), showing more distal lacustrine facies (B and C) and additions of unsorted detritus from glacier overhang and icebergs. Note grounded iceberg and resultant depression. (From Thomas, 1984b, fig. 5).

Figure 14.7 Photo of laminated glacilacustrine sediment containing ice-rafted detritus. In the upper part a till-like sediment (diamicton) has been produced.

14.3.4 below), and accumulations may occur at the ice margin (see section 14.3.3.3, below). Where wave energy is greater, turbulently suspended finer sediment may be transported offshore in the surface layer of the lake (epilimnion), where it eventually is deposited during periods of low wave activity or after the lake surface has frozen and ends all wind stress. Of course, biogenic materials and even clays agglomerated by organisms into faecal pellets (Smith and Syvitski, 1982) may be added to these sediments. Where the load of a river entering a lake is fine grained, wave turbulence may keep some of the sediment suspended for long periods, resulting in some of these clays and silts overflowing from the basin through the outlet during the ice-free season (Smith et al., 1982).

In deeper waters, particularly away from the ice margin, the clays and fine silts that settle out of the epilimnion are commonly laminated, with fine-coarse couplets resulting from the seasonal variation in runoff to the lake and the seasonal influence of wind that keeps some sediment suspended until the lake freezes. These annual sediment couplets, called varves (Fig. 14.8), form extensive blankets in many proglacial lake basins. Although the thickness and nature of the couplets vary spatially and through time - being related to sediment supply, proximity to the source (ice margin, river mouth), and depth below wave base - they commonly remain monotonously similar through time (Fig. 14.8). In fact, this repetitive aspect argues for control by an annual lake rhythm. Where couplets can be identified with confidence as annual increments, they have been effectively used to establish a glacial and glacilacustrine chronology, (e.g. DeGeer, 1912; Antevs, 1922; O'Sullivan, 1983; Ringberg, 1991; Wohlfarth et al, 1993). Smith and Ashley (1985) discuss the difficulty of distinguishing true annual couplets from other rhythmic processes in lakes such as pulsing density underflow currents, repeated slumping, and

Figure 14.8 Typical light-dark (summer-winter) couplets (varves) of relatively coarse and fine sediment. (After Teller (1987, fig 3)

varying wave input. In fact, individual varves themselves commonly display seasonal grain-size variations that relate to wave energy and changes in sediment input due to meltwater runoff, precipitation, and slumping of unconsolidated sediment along the lake margin (Fig. 14.9). Ringberg (1984) has even identified a diurnal rhythm of accumulation within thick proximal varves. As described by Smith and Ashley (1985), true varves (versus surge rhythmites) may contain a lake-floor fauna, a seasonal progression of pollen and other organisms, internal interruptions in sedimentation, and grain sizes in the coarser part of the couplet that displays either an upward coarsening or no trend (versus an upward fining in gravitationally implaced sediments).

Interestingly, in some parts of the largest proglacial lakes, such as Lake Agassiz and the glacial Great Lakes, sediments are not varved, except in more protected areas (e.g. Antevs, 1951; Teller, 1976, 1987; Colman et al., 1994). In part this may be because wind mixing of the water column re-suspended several years of previously-deposited sediment, homogenizing the fine-grained winter accumulation with the coarser summer increment. Continuous density underflow throughout the year may have played a role in inhibiting varving. Additionally, iceberg and winter ice scouring of bottom sediments in tens, if not hundreds, of metres of water may have resulted in homogenization of materials to a depth of several metres (Woodworth-Lynas and Guigné, 1990). Low lake levels may be reflected in the deeper-water sedimentary sequence by erosional unconformities, alluvial units, coarser lacustrine sediment, or by accumulation of units enriched in either in situ or reworked organic matter.

14.3.3.3 Accumulation at the Mouths of the Rivers and Near the Ice Margin

High sediment loads are common in rivers entering proglacial lakes directly from the ice margin, as well as from newly deglaciated terrestrial areas; rapid sedimentation is common in oq0oqo°Sb°o°0 boo'

Fall and winter thaws

Each graded lamina related to increased runoff, wind energy, or nearshore instability

Spring melting

0 50 100

% of size fraction

Figure 14.9 Schematic diagram of grain size variation in a single varve, resulting from changes in runoff, wind, or gravitational transfer. Note that the typical overall coarse to fine (summerwinter) nature still is present. (After Teller (1987, fig. 7).)

proximal areas. Under some circumstances, traditional Gilbert-type deltas with steep foreset beds and associated topset (subaerial) and bottomset (deeper water) beds may be deposited (Fig. 14.10A) (Fyfe, 1990; Postma, 1990). These commonly form at the mouths of rivers where the velocity of the inflowing water is abruptly checked, and the positive imbalance between sediment supply and water depth leads to sediment storage on the delta slope that exceeds removal and downslope transfer (Nemec, 1990). In an ice-marginal environment, the delta surface is graded to the supplying river and to lake level (Fig. 14.10A). If the ice margin remains stable for a relatively long period of time, a subaerial distributary river system develops, depositing the alluvial sediments of the topset portion of the delta. Foreset thicknesses are controlled by the depth of water, and range from a few metres to over 100 m. These beds prograde into the lake in a complex way, mainly by avalanching, with gentler (24—29°) slopes in sandy sediment and steeper (30-35°) slopes in gravels (Jopling, 1965; Nemec, 1990); gentler concave-up slopes and more tangential transitions to bottomset beds are found in finer sediment because more sediment remains in suspension past the crest of the foreset slope (Jopling, 1965). Foreset beds commonly parallel the delta face and may extend over the entire downslope length of the delta face. However, variable avalanching and sediment supply may lead to discontinuous beds that wedge-out upslope or downslope, or that assume lobe or tongue shapes (Nemec, 1990; Fig. 14.10). Because water levels in proglacial lakes frequently change, deltas may become 'stacked' (Thomas, 1984a), drowned, or eroded during transgressive phases and be incised by rivers when lake level drops; new deltas or fans then form at the new lake level.

Figure 14.10 Three possible types of sedimentation in a lake at the margin of an ice sheet (Fyfe, 1990, fig. 7). A) Gilbert-type delta complex fed by two subglacial conduits, showing subaerial alluvial topsets and steeply-dipping foreset beds that merge with bottomset beds in deeper water. B) Subaqueous fan delta lobes fed by subglacial conduits, which may eventually aggrade enough to initiate a Gilbert delta; avalanching and mass movement dominate; sedimentation away from the ice margin is mainly from density underflow currents. C) Subaqueous fan apron in deep water fed by numerous conduits linked by a subglacial cavity system; the complex of rapidly-deposited fluvial sediment and diamicton from the ice commonly is called a grounding-line fan (see also Fig. 14.6).

Figure 14.10 Three possible types of sedimentation in a lake at the margin of an ice sheet (Fyfe, 1990, fig. 7). A) Gilbert-type delta complex fed by two subglacial conduits, showing subaerial alluvial topsets and steeply-dipping foreset beds that merge with bottomset beds in deeper water. B) Subaqueous fan delta lobes fed by subglacial conduits, which may eventually aggrade enough to initiate a Gilbert delta; avalanching and mass movement dominate; sedimentation away from the ice margin is mainly from density underflow currents. C) Subaqueous fan apron in deep water fed by numerous conduits linked by a subglacial cavity system; the complex of rapidly-deposited fluvial sediment and diamicton from the ice commonly is called a grounding-line fan (see also Fig. 14.6).

Sediment entering ice-marginal lakes through meltwater rivers at or near the base of a glacier may be deposited as subaqueous fans, the beds of which slope more gently into deeper water (e.g. Gustavson et al, 1975; Powell, 1990; Eyles and Eyles, 1992; Ashley, 1995; Figs. 14.10B, 14.10C, 14.11A). Benn and Evans (1998) discuss this environment in detail. High sediment concentrations in the inflowing river favour gentler slopes and fan (versus Gilbert delta) construction by density underflow currents (Jopling, 1965). As the ice margin retreats, a broad esker-like ridge may develop by overlapping fans if the tunnel remains active (Fig. 14.11B). These fan-shaped deposits may merge with other fans deposited at the mouths of subglacial conduits along the ice margin, forming a variably continuous ridge of subaqueous fluvial sediment (Fig. 14.10B and 14.10C), such as those of the extensive Salpausselka moraines across southern Finland that were deposited as ice retreated into shallower water along the northern side of the Baltic Ice Lake (e.g. Fyfe, 1990; Raino et al., 1995). These are commonly called grounding-line fans (Fig. 14.6).

Sediment in the resultant asymmetrical bedform has a steep ice-contact side, fines downslope into the basin and typically consists of gravels, sands and occasional diamicton units (Fig. 14.12); some have considered these ridges DeGeer moraines. Because sediment transport across ice-marginal subaqueous fans is mainly in highly turbulent flows and by gravitational transfer, sorting and bedding may be poor (Nemec, 1990; Benn and Evans, 1998). A drape or fringe of diamicton on these fans is common where there was an overhanging ice margin (Fig. 14.10), and Thomas (1984b) reports conical mounds of diamicton within the fan sequence that have

Figure 14.11 Schematic of meltwater deposition in proglacial lake at glacier margin. (After Sharpe et al., 1992). A) Subaqueous fan at mouth of subglacial tunnel. B) Series of overlapping subaqueous fans producing a broad esker-like ridge as ice retreats. C) Subglacial meltwater outburst that produces a broad moraine or extended ice-marginal fan below lake level.

Figure 14.11 Schematic of meltwater deposition in proglacial lake at glacier margin. (After Sharpe et al., 1992). A) Subaqueous fan at mouth of subglacial tunnel. B) Series of overlapping subaqueous fans producing a broad esker-like ridge as ice retreats. C) Subglacial meltwater outburst that produces a broad moraine or extended ice-marginal fan below lake level.

Transport Clayey Till
Figure 14.12 Moderately well-sorted and bedded ice-contact gravel in fan delta near Winnipeg that is overlain by 1 m thick diamicton (flow till) that thins toward the left and is capped by laminated glacial Lake Agassiz silts and clays.

up to 60 cm of relief (Fig. 14.6). Faulting and deformation in these units is common where deposition was on or against the ice (Fig. 14.13). Readvances of the margin may result in considerable deformation and thrusting of these ice-marginal sediments (e.g. Boulton, 1986); this may occur on an annual basis as a result of the decline in winter ablation of the ice margin

On the steep slopes of subaqueous fans deposited in deep water, debris avalanching is common across the surface, resulting in deposition of coarse lobes and anastomosing gravel- and sand-filled chutes like those found in subaerial alluvial fans (Fig. 14.10C) (e.g. Fyfe, 1990; Nemec, 1990; Prior and Bornhold, 1990). Debris-flow diamictons are commonly deposited, reflecting the proximity to the ice margin, and slumping of the proximal side of ice-contact deltas may occur as a result of the melting of buried ice and removal of ice-marginal support. These episodic, subaqueous mass movements may evolve into dense turbulent underflows or turbidity currents, which erode channels in pre-existing substrate and rapidly deposit finer sediment in deeper water (e.g. Prior and Bornhold, 1990; Benn and Evans, 1998).

Abrupt releases of subglacial meltwater may occur during glacial retreat (e.g. Shaw et al., 1989; Shoemaker, 1992; Shaw, 1996) and during rapid advances such as surges (Sharpe and Cowan, 1990). Short-distance surges (see Dredge and Cowan, 1989, Fig. 8.23), as well as extensive surges (Clayton et al., 1985; Clark, 1994a), appear to have been common in ice-marginal lakes and may have coincided with outbursts of subglacial water. Although the processes and magnitudes of such outbursts are controversial, they would have delivered large volumes of stored water and debris to the glacier margin in a short time. A study by Sharpe and Cowan (1990) in the Lake Agassiz basin led to the conclusion that some subaqueous moraines with a

Figure 14.13 Highly faulted bedded delta sands that were deposited over ice along retreating ice sheet in southern Sweden; note coarse ice-marginal gravels at right.

steep proximal and gentle distal form might have been deposited along the ice margin during such outbursts (Fig. 14.11C); these moraines consist of stratified-to-massive sediment that grades upward from gravel to sand and silt. They suggest that this outburst may have resulted from a rapid lowering of lake level, which produced a large hydrological differential (Fig. 14.14). This type of moraine may be similar in nature to subaqueous grounding-line fans that accumulated by 'normal' meltwater discharge from various subglacial tunnels. Beyond the ice margin, sedimentation rates, grain size and sedimentation style reflect such outbursts. In addition, when the confining margins of ice-marginal lakes fail, large slugs of water are normally released. This may have a domino effect on other lakes downstream, as well as on fluvial systems, and subaqueous fans may be very rapidly deposited in these downstream lakes (e.g. Kehew and Clayton, 1983; Kehew and Lord, 1987; Kehew and Teller, 1994). Turbulent high discharges of sediment from subglacial tunnels (or proglacial rivers), as well as those related to mass sediment transfer such as slumps and avalanches from delta faces and ice margins, commonly evolved into density underflows that moved along the lake floor until their turbulence and energy dissipated (Fig. 14.15). These underflows commonly were episodic, although some glacial rivers may have delivered a continuing, albeit varying, high concentration of sediment to proglacial lakes for long periods, which retained their identity over long distances and over gentle bottom slopes (e.g. Smith and Ashley, 1985). Much of this sediment was focused into relatively low topographic areas on the basin floor. In fact, most sediment in the deepest part of large proglacial lakes was probably transported there by these subaqueous gravity currents (see Smith and Ashley, 1985), and the so-called bottomset units of Gilbert-type deltas are partly composed of sediment from these flows and partly of material settling out from higher in the water column. Where there is a high influx of suspended fine sediment (usually fine sands and coarse silts) and where the sediment accumulation rate is high, a climbing-ripple

364 GLACIAL LANDSYSTEMS

(A) High lake level

Lake Agassiz y

(A) High lake level

Lake Agassiz y

(B) Rapid lowering H0

Possible

V-P0 supraglacial lake

(Unstable, rapidly? flowing ice)

(B) Rapid lowering H0

(Unstable, rapidly? flowing ice)

Subglacial drainage and erosion

Sedimentation starts

Subglacial drainage and erosion

Sedimentation starts

Surged margin? Extended roffle

Reduced potentiometric surface

Surged margin? Extended roffle

Reduced potentiometric surface

New moraine

Figure 14.14 Schematic of subaqueous moraine formation by subglacial outbursts into an ice-marginal lake. A) Initial high lake stage, showing stored subglacial water. B) Abrupt lowering of lake that increases pressure head and promotes outward flow of subglacial water and sediment; glacier surging may also occur at this stage. C) New moraine, expanded and thinner ice margin, and new stable condition following subglacial outburst. (Sharpe and Cowan, 1990, fig. 10).

New moraine

Figure 14.14 Schematic of subaqueous moraine formation by subglacial outbursts into an ice-marginal lake. A) Initial high lake stage, showing stored subglacial water. B) Abrupt lowering of lake that increases pressure head and promotes outward flow of subglacial water and sediment; glacier surging may also occur at this stage. C) New moraine, expanded and thinner ice margin, and new stable condition following subglacial outburst. (Sharpe and Cowan, 1990, fig. 10).

sequence may be deposited (Fig. 14.16), which consists of a long succession of ripples whose crests are incrementally offset from one another (in-drift climbing ripples) or lie directly above each other (in-phase climbing ripples) (e.g. McKee, 1965; Jopling and Walker, 1968; Gustavson et al., 1975).

Figure 14.15 Ice-marginal lake showing distribution of suspended sediment (light grey) arriving from glacier and river as underflow, interflow, and overflow in water. All but finest grains move across lake floor as relatively dense underflow suspensions into centre of basin, progressively depositing finer and finer sediment. Finer grains may be transported by wave energy as overflows and intermediate sizes may move at the epilimnion-hypolimnion density interface. (After Teller, 1987, fig. 4).

Figure 14.15 Ice-marginal lake showing distribution of suspended sediment (light grey) arriving from glacier and river as underflow, interflow, and overflow in water. All but finest grains move across lake floor as relatively dense underflow suspensions into centre of basin, progressively depositing finer and finer sediment. Finer grains may be transported by wave energy as overflows and intermediate sizes may move at the epilimnion-hypolimnion density interface. (After Teller, 1987, fig. 4).

14.3.4 Nearshore Sediments

Wave action and longshore drift commonly reworked sediment delivered to a lake, depositing grains around the margin that were too large to be transported into deeper water. In simplest situations, each beach is an indicator of lake level, and is a chrono-morphological and chrono-stratigraphic marker (Fig. 14.17). However, subaqueous bars, just offshore from the main beach, which commonly result from wave return flow, may form contemporaneously with the main beach (Figs. 14.18) (e.g. Walker and Plint, 1992). Around large lakes, where wind fetch is high, storm beaches may introduce further interpretive difficulties, because their elevation may extend well above the normal level of the lake (cf. Otvos, 2000). Thus, several levels of 'beaches' may form at the same time. Furthermore, wind activity during and after beach formation may add a cap of dunes on these lake-level landforms, adding to the difficulty of establishing exactly what the water level was in a lake at a given time.

As waves eroded their margins, helping to provide sediment for beaches, wave-trimmed scarps developed on headlands and exposed shorelines. Less distinct 'washing limits' of waves around the lake margin can also be correlated, and provide evidence for palaeo-lake level (Veillette, 1994).

When lake levels declined, nearshore deposits regressed downslope at the lake margin (e.g. Fig. 14.2A). Unless there were interruptions to this decline in lake level, or unless storm beaches were constructed, there was little morphology associated with deposition, and a blanket of sandy sediment was deposited, generally with only low-relief ridges on it (e.g. Posamentier et al., 1988; Walker and Plint, 1992; Thompson and Baedke, 1997).

As lake levels rose, beaches migrated upslope (e.g. Fig. 14.2C). This transgression resulted in waves reworking shoreline deposits so that, except when rapid drowning of the shoreline occurred, older

Figure 14.16 A) Schematic of climbing ripple lamination, showing evolution of in-phase ripples to in-drift ripples and the decrease in angle of climb as the ratio of fallout from suspension to bedload decreases (after Jopling and Walker, 1968). Flow left to right. Preservation of laminae on stoss side of ripple decreases upward. If suspended load/bedload decreases to where there is little sediment in suspension, ripples will only migrate and will not grow upward at the same time. B) In-drift climbing ripples.

Figure 14.16 A) Schematic of climbing ripple lamination, showing evolution of in-phase ripples to in-drift ripples and the decrease in angle of climb as the ratio of fallout from suspension to bedload decreases (after Jopling and Walker, 1968). Flow left to right. Preservation of laminae on stoss side of ripple decreases upward. If suspended load/bedload decreases to where there is little sediment in suspension, ripples will only migrate and will not grow upward at the same time. B) In-drift climbing ripples.

Glacial Lake Agassiz Beaches

Figure 14.17 Schematic of relative changes in level of Lake Agassiz from 1 1.7 to 7.7 ka, showing rising (transgressing) stages resulting from differential isostatic rebound after the opening of a new lower outlet caused an abrupt drop in lake level. The occurrence of transgressions indicates that these shorelines were south of the outlet. (After Teller, 2003).

Figure 14.17 Schematic of relative changes in level of Lake Agassiz from 1 1.7 to 7.7 ka, showing rising (transgressing) stages resulting from differential isostatic rebound after the opening of a new lower outlet caused an abrupt drop in lake level. The occurrence of transgressions indicates that these shorelines were south of the outlet. (After Teller, 2003).

beach morphology was destroyed. As discussed previously and illustrated in Fig. 14.2, in proglacial lakes around continental ice sheets, differential isostatic rebound resulted in lake level transgression south of the isobase through the overflow outlet, and regression north of the isobase. This means that large beaches did not form north of the isobase that extends through the outlet, and that beaches to the south reflect only the maximum level of transgression for any given phase of a large lake (Teller, 2001).

Figure 14.18 Cross-section sketch of barrier beach between lagoon and lake, showing gently dipping (planar) forebeach laminae (FB), nearly flat top-beach laminae (TB), and angle of repose backbeach laminae (BB) that prograde into lagoon during periods when waves overtop the barrier. Subaqueous bar may be tip of growing spit or ridge accumulated by wave return flow.

Figure 14.18 Cross-section sketch of barrier beach between lagoon and lake, showing gently dipping (planar) forebeach laminae (FB), nearly flat top-beach laminae (TB), and angle of repose backbeach laminae (BB) that prograde into lagoon during periods when waves overtop the barrier. Subaqueous bar may be tip of growing spit or ridge accumulated by wave return flow.

As with beaches around non-glacial lakes, bedding gently slopes lakeward on the foreshore (Fig. 14.18). Straight, narrowly spaced, symmetrical ripples, typically with rounded crests, form in the nearshore zone where there is an overall balance between wave uprush and return flow, whereas asymmetrical ripples form where unidirectional flow dominates; close to shore, megaripples may form and these become planar laminae on the beach face (e.g. Clifton et al., 1971). Although grain size is related to wave energy, which is partly a function of wind fetch, it is limited by the supply of grains from shoreline erosion and longshore drift. Headlands commonly have no beaches, or only coarse-grained lags because wave energy is focused on them, in contrast to the finer-grained beaches in protected bays and in shallow water. Although beach sediments are typically well sorted and bedded, this is not always the case in ice-marginal lakes because of the shorter ice-free season for wave action and the potential for disturbance by floating ice. Ice-rafted coarse material may be added to these beaches.

Morphologically, beaches may form berms attached to the mainland or offshore barrier islands, separated from the mainland by a lagoon. Barrier beaches commonly form by the growth of a spit and may have an additional steeply dipping set of cross beds (backset beds) that form when waves overtop the barrier and flow into the lagoon, burying these organic-rich sediments (Fig. 14.18); beds on the lakeward-facing side may be eroded if the barrier is transgressing into the lagoon. In large proglacial lakes, such as Lake Agassiz, these barrier beaches can be hundreds of metres wide and stand tens of metres above the adjacent floor of the lake. In contrast to the generally sparse organic content in sediment of large ice-marginal lakes, these lagoons commonly provided an important palaeo-ecological record because of their warmer and less-turbid environment and may lie buried beneath barrier beaches of proglacial lakes (e.g. Teller et al., 2000; Fig. 14.18).

In general, most proglacial lakes developed an irregular stair-step of beaches (Fig. 14.17) around the basin south of the isobase that extended through the outlet. Most beaches are discontinuous and some have associated wave-trimmed shorelines. In differentially rebounding areas these shorelines can be used to define the outline of the lake at a given transgressive maximum. Today, these once-horizontal beaches have been deformed by differential isostatic rebound, with elevations rising toward the centre of maximum rebound. In the Lake Agassiz basin, most beaches in northern regions are now over 150 m above their contemporaneously formed equivalents 500-700 km to the south (Fig. 14.19) (Teller and Thorleifson, 1983).

Rebounded ancient beaches can be used, in concert with ancient marine beaches, to reconstruct the rates of isostatic rebound as well as total postglacial rebound of a continent (e.g. Andrews and Peltier, 1989; Dyke, 1996). In addition, by comparing the curvature (slope) of upwarped shorelines in the same region, their relative age can be determined. Specifically, older beaches in a sequence will have undergone more differential rebound than younger beaches, so will have steeper gradients which diverge from those of younger beaches (Fig. 14.1).

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