Effect of drifting snow on the velocity field

Glaciers flow over irregular beds, and thus have undulating surface profiles. Furthermore, their transverse flow patterns may be influenced by nunataks or irregular valley walls. Patterns of both accumulation and ablation thus can be uneven owing to drifting and to shading from the

June snow depth (exaggerated)

June snow depth (exaggerated)

Figure 5.14. Effect of drifting snow on the surface profile of a glacier. Owing to the additional accumulation in the lee of the surface convexity at A, ws does not need to be as high at B as otherwise would be the case.

Figure 5.14. Effect of drifting snow on the surface profile of a glacier. Owing to the additional accumulation in the lee of the surface convexity at A, ws does not need to be as high at B as otherwise would be the case.

50 m sun during the melt season. We have just discussed one example of this from Siple Dome. Let us now consider some other examples.

To understand how drifting influences the flow field and surface profile, consider the hypothetical situation shown in Figure 5.14 in which a glacier flows over a convexity in the bed, resulting in a similar convexity in the surface. Owing to drifting in the lee of the surface convexity, the normal June snow depth at B is, say, 2 m, while that at A it is only 1 m. During a normal melt year, suppose that at A all of the snowand0.5mof the underlying ice melts, whereas at B, melting removes only the snow cover. Thus, the emergence velocity at A must be 0.5 m a-1, whereas at B it is 0. In the absence of the extra accumulation at B, the glacier would probably be thinner here as shown schematically by the dotted line in Figure 5.14. The greater surface slope between A and C would then provide the increased longitudinal compression needed to develop a positive emergence velocity at B.

The situation shown in Figure 5.14 occurs on a large scale on the surface of the Antarctic ice sheet above the western edge of Lake Vostok, a subglacial lake under 4 km of ice in central East Antarctica (see Figure 6.13). The increase and then decrease in surface slope reflects flow of ice over a steep slope down into the lake and then an abrupt decrease in basal drag as the ice moves out over the lake. As this is an accumulation area, the thicker accumulation (as at B) is advected downglacier and buried. Because flow rates are relatively low, ice moving over the lake experiences this excess accumulation for about 30 000 years. The excess shows up in radio echo profiles as an increase in the vertical distance between reflectors, and in an ice core from a borehole on the east side of the lake as a zone of high accumulation rate between—800 and—1100 m depth (Leonard et al, 2003).

Thule-Baffin moraines (Figure 5.15), first studied in detail by Goldthwait (1951), provide another geomorphologically significant

Figure 5.15. Thule-Baffin moraines (on skyline) in one of the type areas, Thule (Qanaq), Greenland. Note folding in foliation (sedimentary stratification) in superimposed ice in center of photograph. (From Hooke, 1970, Figure 8. Reproduced with permission of the International Glaciological Society.)

example of the importance of drifting snow. The casual observer will commonly be surprised to learn that although the crest of the moraine in Figure 5.15 is several tens of meters above the margin, the till is usually no more than a meter or so thick. Beneath the till is dirty ice with quite variable debris concentrations. The dirt in the ice is typically segregated into laminae or folia, millimeters in thickness, that dip steeply upglacier (Figure 5.16b). The sediment content of the dirt-bearing folia is normally only a few percent. Layers of clast-supported frozen till, sometimes exceeding 1 m in thickness, are also present. The wedge of ice downglacier from the moraine is clean. It is too thin to flow at an appreciable rate, and is frozen to its bed, preventing sliding. The low flow rate, in conjunction with the observed dip of the foliation, led to the mistaken impression that the foliation planes were shear planes, and that the dirty ice was actively shearing over the wedge of clean ice in such a way that debris, entrained at the bed, was carried to the surface on these planes.

(b) 100 Elevation, m

Strain rate, a-1

June snow cover, m

Snow and superimposed ice over debris

Snow and superimposed ice over debris

(e) Ablation, 0-8 emergence velocity,

-Ablation

-Ablation

(e) Ablation, 0-8 emergence velocity,

900 m

Velocity scale,

900 m

|~~1 Snow and undeformed J Estimated error in superimposed ice measurement

| | Blue ice I Measurement of same

Fvl riAk,™ „r „„ quantity in two successive hs,°° Debris in or on ice ^ '

years

Figure 5.16. Data from a Thule-Baffin moraine on Barnes Ice Cap, Baffin Island. (a) Map showing moraine and line of stakes used for velocity and mass balance measurements. Velocities are shown by arrows. (b) Surface profile along stake line, showing dip of foliation at surface and inferred dip beneath surface. (c) Strain rates. (d) June snow depth. (e) Net balance and emergence velocity along stake line. (Modified from Hooke, 1973a, Figure 3D. Reproduced with permission of the Geological Society of America.)

Velocity scale,

The geometry of this type of glacier margin can be understood in terms of the concepts we have been discussing. At "a" in Figure 5.16b, the mean June snow cover was about 1 m thick in the 1970s (Figure 5.16d). During an average 1970s summer, this snow and —0.55 m of the underlying ice melted. Longitudinal compression here (Figure 5.16c) resulted in a positive (upward) ws, and the emergence velocity was —0.50 m a-1 (Figure 5.16e). This was slightly less than the net balance, but high enough that a slight cooling could have brought about a steady state with Equation (5.26) satisfied.

At "b" in Figure 5.16b, the snow cover was 0, but so was the ablation rate as the till layer insulated the ice. On ridges like these, it is difficult to know what is meant by "emergence velocity", as a ridge has both up- and downglacier slopes. However, because there was still some longitudinal compression, the ridges were gradually increasing in height. This is a non-steady-state process: as the height of such a ridge increases, it becomes steeper until, eventually, till slumps off of it, exposing the underlying ice. This ice then melts rapidly owing to the lack of snow cover and to the thin covering of dirt that remains on it, decreasing its albedo. However, a new ridge begins to develop under the slumped till.

At "c" the June snow cover was nearly 2.5 m thick. During an average 1970s summer, this snow melted, but essentially no ice was lost: bn « 0. Thus, this sloping margin could exist despite the fact that the emergence velocity in it was negligible.

Thus, during a period of balanced mass budget along a glacier margin like this, till-covered ridges would go through cycles of growth and decay while the ice surface upglacier and the wedge of deformed superimposed ice downglacier remained unchanged. The primary change would be an increase in dirt cover as more debris melted out of dirty ice exposed by slumping from the ridges.

During a series of cool summers the sloping margin downglacier from the moraines becomes a local accumulation area at the edge of the glacier. If cool climatic conditions persist long enough, the glacier will advance, overriding and deforming this accumulation of superimposed ice, as shown in Figure 5.17. Recognition of this process provided an alternative to the shearing mechanism proposed by Goldthwait.

Three lines of evidence support the origin of Thule-Baffin moraines shown in Figure 5.17. Firstly, the less-deformed superimposed ice is fine grained (1-2 mm) and lacks any development of a deformation fabric such as would be present in highly deformed basal ice. Secondly, oxygen isotope ratios show that the deformed superimposed ice accumulated under conditions broadly similar to those prevailing today, yet it, in part (Figure 5.17), underlies ice with isotopic ratios characteristic of Pleistocene ice. Thirdly, folding of the downglacier dipping layers of superimposed ice occasionally can be observed in ice cliffs with the proper orientation (Figures 5.15 and 5.18). Further discussion of this process and the evidence for it is presented by Hooke (1970, 1973a, 1976) and Hooke and Clausen (1982).

Moraines are formed by the process illustrated in Figure 5.17 only under relatively cold climatic conditions. For example, in the Kanger-lussuaq area of Greenland, a few hundred kilometers south of Thule (now Qanaq), summer temperatures are warm enough to melt all of the snow at the margin, even though drifting can result in a rather thick June snow cover there. Thus, there is no marginal zone of superimposed ice. However, the process illustrated in Figure 5.17 was probably important in northern Wisconsin, Minnesota, North Dakota, Alberta, and

V

Glacier is frozen to bed so contact between glacier ice and superimposed ice - remains fixed through time

Warming climate results in thinning and separation of moraine from ice sheet

K'-v'i>l Undeformed and deformed M-V-.'l Debris-bearing glacier ice superimposed ice .-. , . .

|_| Clean glacier ice

~100 m Vertical exaggeration approximately 2x

Figure 5.17. Sequential cross sections showing schematically the processes by which superimposed ice is overridden during an advance of a glacier that is frozen to its bed at the margin. Light lines in superimposed ice show deformation of sedimentary layering. Last cross section shows how moraine becomes separated from glacier during subsequent retreat. (Modified from Hooke, 1973a, Figure 1. Reproduced with permission of the Geological Society of America.)

Saskatchewan where, over time spans of millennia as the ice advanced in the late Wisconsinan, geomorphic features indicate that huge quantities of till accumulated on the glacier surface (Attig et al., 1989; Moran et al., 1980) despite the lack of nunataks projecting above the surface. Ice wedge casts and other evidence indicate that the ice advanced over permafrost in these areas.

Disintegration ridges (Gravenor and Kupsch, 1959) are one of the primary geomorphic features suggestive of such a thick till cover.

Figure 5.18. Ice cliff near Thule, Greenland. Sedimentary bedding on the left is wrinkled and overturned by ice advancing from the right. The boundary between the active and the less rapidly deforming superimposed ice is marked by a dirt band (arrows) which can be traced into a Thule-Baffin moraine. (Person in white circle for scale.)

Figure 5.18. Ice cliff near Thule, Greenland. Sedimentary bedding on the left is wrinkled and overturned by ice advancing from the right. The boundary between the active and the less rapidly deforming superimposed ice is marked by a dirt band (arrows) which can be traced into a Thule-Baffin moraine. (Person in white circle for scale.)

Disintegration ridges are believed to form through a series of topographic reversals such as those just described as occurring on Thule-Baffin moraines (Clayton and Freers, 1967). As the ice sheet stagnated and began to melt down, it is suggested that ridges grew under areas of thicker debris and then melted when the debris slumped, only to be replaced by new ridges that developed under the slumped debris (Figure 5.19a). The only difference between this process and that which forms Thule-Baffin moraines is that in stagnating ice there is no significant longitudinal (or transverse) compression to generate upward flow, so the surface is continually lowered. Thus, the final slumping event deposits the debris directly on the bed. Relatively linear ridges are commonly formed by this process, although hills and circular ridges, informally called doughnuts (Gravenor, 1955), are also found. These features can be several meters high. Linear ridges can be hundreds of meters or kilometers in length, and may have most any orientation with respect to the

Ice-walled lake plain

Disintegration ridge

Disintegration ridge

Figure 5.19. Schematic sketches showing origin of (a) a disintegration ridge, and (b) an ice-walled lake plain. (Based on Clayton and Freers, 1967.)

Figure 5.19. Schematic sketches showing origin of (a) a disintegration ridge, and (b) an ice-walled lake plain. (Based on Clayton and Freers, 1967.)

ice margin, depending on whether they formed from morainal accumulations like those illustrated in Figures 5.16 and 5.17 or alternatively from accumulations of debris in crevasses or superglacial stream channels.

Ice-walled lake plains are another geomorphic feature formed during disintegration of debris-covered ice masses and commonly found in the upper midwest of the United States and adjacent areas of Canada (Clayton and Cherry, 1967). The lakes would have first formed when differential melting resulted in depressions in the ice surface. Because water is densest at +4 °C, surface water that warms to this temperature sinks. The resulting convection would have circulated +4 °C water to the bottoms of the lakes, where it could melt ice if the ice was not too cold and the sediment layer in the bottom of the lake not too thick. The lakes thus would have deepened. Field evidence shows that some eventually penetrated all the way to the bed. Superglacial streams then brought sediment to the lakes, forming typical lacustrine deposits with sand and gravel deltas near stream mouths and lacustrine silts and clays further from shore. When the surrounding ice melted, these deposits were left as flat-topped hills on the landscape (Figure 5.19b).

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