Arctic Glaciers

Net surface mass balance is the sum of winter accumulation and summer losses of mass from glaciers and ice sheets. Where this is positive, glaciers grow and where it is negative, they recede. Net surface mass balance is determined primarily by changes in climate. Dowdeswell et al. (1997) examined records of surface mass balance for more than 40 Arctic ice caps and glaciers going back in some cases as far as the 1940s. They deliberately excluded the large Greenland Ice Sheet because its surface mass balance is so poorly constrained. The remaining Arctic glaciers and ice caps cover about 275 000 km2 of the archipelagos of the Canadian, Norwegian, and Russian High Arctic and the area north of about 60°N in Alaska, Iceland and Scandinavia. The records show that most Arctic glaciers have experienced predominantly negative net surface mass balance over the past few decades. Dowdeswell et al. (1997) found no uniform trend in mass balance for the entire Arctic, although some regional trends occur. Examples are the increasingly negative mass balances for northern Alaska, due to higher summer temperatures, and increasingly positive mass balances for maritime parts of Scandinavia and Iceland, due to increased winter precipitation. The negative mass balance of most Arctic glaciers may be a response to a step-like warming in the early twentieth century at the termination of the cold 'Little Ice Age'. Dowdeswell et al. (1997) calculated that the Arctic ice masses outside Greenland are at present contributing about 0.13 mm yr-1 to global sea-level rise.

Source: Dowdeswell, J.A., Hagen, J.O., Bjornsson, H., et al. (1997) The mass balance of circum-Arctic glaciers and recent climate change. Quaternary Research, 48,1-14.

To understand how the glacier system responds to climate change we need to consider what happens when a glacier thins or thickens due to changes in its mass balance. If climate deteriorates - temperatures fall and/or snowfall increases -every part of the glacier is likely to thicken. This thickening will result from either an increase in accumulation or a reduction in ablation and may cause the ice margin

Positive A

mass balance

Negative mass balance y

10 20 30 Time (years)

Figure 3.18 Short-term and long-term trends in mass balance to advance. Conversely, a climatic warming will lead to overall thinning of the glacier and retreat of its margin. Glacier response to such changes in mass balance may be either stable or unstable. A stable response is one in which the glacier thickens or thins in proportion to the size of the change in mass balance. An unstable response is one in which the glacier thickens or thins out of proportion with the size of the change in mass balance that triggers it. The type of response depends on whether the flow is extending or compressive (Figure 3.8). Unstable behaviour occurs more often with compressive flow. This can be illustrated in relation to Figure 3.19 where A-A and B-B are two cross-sections in a glacier subject to compressive flow. The discharge (Q volume per unit time) of ice through cross-section B-B is less than that through cross section A-A by the amount of ablation (ab) which occurs on the glacier surface between points A and B (QAA = QBB + ab). If we now place a uniform layer of snow over the glacier surface it can be shown mathematically that the increase in discharge caused by the layer of snow is proportional to the original flow through the section. Consequently, because the flow through cross-section A-A is larger than that through B-B, the increase in discharge due to the new layer of snow will be greater at A-A than at B-B (QAA > QBB + ab). Therefore the ice must thicken between point A and B in order to accommodate this increased volume. With continued flow this effect will be accentuated down-ice and the glacier will progressively thicken towards its snout, eventually causing it to advance.

Changes in mass balance are propagated down-glacier by means of kinematic waves. A kinematic wave can be viewed as a bulge in the glacier surface. The greater thickness of ice in the bulge will locally increase the basal shear stress and therefore the rate of ice deformation, and consequently the glacier velocity. The bulge will therefore move down ice faster than the thinner ice on the up-ice and down-ice side

Positive A

mass balance

Negative mass balance y

Loss by ablation

Ice surface

Figure 3.19 The response of a glacier to an increase in accumulation in an area of compressive flow. See text for details. [Modified from: Sugden and John (1976) Glaciers and Landscape,

Edward Arnold, figure 3.12, p. 48]

Bedrock

Figure 3.19 The response of a glacier to an increase in accumulation in an area of compressive flow. See text for details. [Modified from: Sugden and John (1976) Glaciers and Landscape,

Edward Arnold, figure 3.12, p. 48]

of the bulge. It is important to stress that it is the bulge that moves and not the ice: a boulder on the surface would temporally move at a higher velocity when it was on top of the bulge and then return back to its normal velocity as the bulge passed by. Kinematic waves originate in the vicinity of the equilibrium line. In general the equilibrium line represents the junction between predominantly extending flow in the accumulation zone and the compressive flow typical of the ablation zone. An increase in ice thickness, caused by a deterioration in climate, will therefore cause a stable thickening of the glacier above the equilibrium line and an unstable response below it. It is this difference in thickening response that initiates a kinematic wave. The wave, once formed, travels down the glacier at a rate faster than the ice velocity. When the wave reaches the snout, often years later, it may initiate an advance of the ice margin. The passage of a kinematic wave is similar to the passage of a flood wave within a river. It is important to emphasise that kinematic waves are rarely visible at the glacier surface. They are simply the mechanism by which mass balance changes are propagated throughout the glacier system: the means by which it adjusts to changes in climate by extension or contraction of the ice margin.

The rate at which a kinematic wave passes through a glacier is highly variable. Some glaciers respond quickly to changes in mass balance whereas others do not. The length of time for glacier adjustment to a change in mass balance, the response time, depends on the sensitivity of the particular glacier to change and upon the nature of the change. For example, an excess of ablation at the glacier snout one year may cause a glacier snout to retreat almost immediately, by contrast an advance in the snout due to an excess of accumulation must first feed through the whole glacier before it has an effect. The response time is also controlled by the sensitivity of the glacial system, which depends on a variety of parameters, including glacier morphology and activity

(i.e. the mass balance gradient). At a simple level the larger the ice body the slower the response time. Large ice sheets will respond only to large and sustained changes in mass balance, whereas small cirque or valley glaciers will often respond quickly to minor fluctuations. The rate of response of these smaller ice bodies is influenced by their morphology, as illustrated in Figure 3.20. Here two valley glaciers with different slopes are affected differently when the equilibrium line is raised by an identical amount. The ratio of accumulation to ablation area of Valley Glacier B is much more sensitive than that of Valley Glacier A and therefore will be more responsive to climatic change. In a similar manner glaciers which are very active, that is, they have large mass balance gradients, will respond much more quickly to changes in mass balance than those that are less active.

B: Valley glacier B

Figure 3.20 The impact of a rise in equilibrium line altitude (ELA) on two valley glaciers with different gradients. The impact of a change in the ELA is greater on valley glacier B because it has a shallower gradient and the rise in ELA therefore affects a larger surface area. [Modified from: Kerr (1993) Terra Nova, 5, figure 4, p. 335]

Figure 3.20 The impact of a rise in equilibrium line altitude (ELA) on two valley glaciers with different gradients. The impact of a change in the ELA is greater on valley glacier B because it has a shallower gradient and the rise in ELA therefore affects a larger surface area. [Modified from: Kerr (1993) Terra Nova, 5, figure 4, p. 335]

The link between climate, mass balance and glacier response is complex and may not always be apparent. This is particularly the case where glaciers terminate or calve into water. On land a glacier can react to changes in input by extending or withdrawing its snout. An extension of the snout means that an increased surface area is exposed to ablation because the glacier advances into warmer areas at lower altitude. A glacier in a deep fjord - a glacial valley that has been drowned by the sea - cannot do this so easily. This can be illustrated by considering a tidewater glacier in an ideal fjord of constant width and depth (Figure 3.21). The ablation area of the glacier is limited to the lower reaches of the glacier and, more importantly, to the amount of ice that can be melted or discharged as icebergs from the cross-sectional area of the snout. If there is a shift to a positive net balance the glacier will begin to advance. As it cannot extend to lower altitudes to enhance ablation, it will continue to advance until it can spread out and increase the cross-sectional area exposed to melting and calving. This will occur only at the fjord mouth or at a point at which the fjord widens or deepens. Consequently, fjord glaciers are particularly sensitive to changes in mass balance, and relatively minor climatic changes can cause spectacular variations in the position of snouts. The most stable positions for calving glaciers within fjords include: (i) fjord mouths; (ii) fjord bifurcations; (iii) points where fjords widen, narrow or are bordered by low ground; and (iv) where they deepen or shallow. These locations are known as pinning points, points where the ablation geometry of the ice margin may be changed and control the location of calving ice margins (Figure 3.21).

Figure 3.21 The role of fjord geometry in controlling the location of ice margins. (A) In afjord of constant width and depth a glacier can advance to the fjord mouth, but will not advance beyond the fjord mouth because of the increase in calving rate at this point. (B) Examples of topographic pinning points (dashed lines) for a calving glacier. Calving glaciers will tend to be stable at these pinning points

Figure 3.21 The role of fjord geometry in controlling the location of ice margins. (A) In afjord of constant width and depth a glacier can advance to the fjord mouth, but will not advance beyond the fjord mouth because of the increase in calving rate at this point. (B) Examples of topographic pinning points (dashed lines) for a calving glacier. Calving glaciers will tend to be stable at these pinning points

Understanding the interaction of climate, mass balance and glacier response is important not only in examining the response of glaciers to relatively minor climatic fluctuations but also to the more dramatic and long-term fluctuations associated with the onset of a glacial cycle. The growth and decay of large ice sheets is a complex problem. Traditionally, ice sheets are considered to grow through a sequence of larger and larger ice bodies - snow patches > cirque gla-ciers>valley glaciers>icefields>ice caps>ice sheets - developing first in upland areas and then expanding into lowland regions as the ice bodies merge and grow. More recently computer models have suggested that the sequence of growth may follow a slightly different pattern - snow patches ! cirque glaciers ! valley glaciers ! piedmont lobes ! small ice sheet. In this scenario valley glaciers first develop in mountainous areas and flow out into low-lying areas where the ice spreads out as large lobes, known as piedmont lobes. These lobes merge and thicken rapidly in an unstable fashion to produce small ice sheets, which then merge to produce larger ones. The piedmont lobes thicken and grow dramatically because they have low gradients and consequently low ice velocities, and therefore cannot discharge the ice pouring into them from the fast flowing valley glaciers that drain the steep mountainous areas behind. Topography may also play a very important part in the rate at which an ice sheet develops. This has been illustrated by a computer model of a former ice cap that existed in the Scottish Highlands at the close of the last glacial cycle, during a cold period known as the Younger Dryas (10 000 years ago). This computer model illustrated the sensitivity of this ice cap to the mountain topography of the highlands: certain types of topography accelerated its growth. As illustrated in Figure 3.22 for a given deterioration in climate, parts of the ice cap that advanced into large basins grew more dramatically than

Time (years)

Figure 3.22 The effects of topography on rates of ice-sheet growth in a numerical ice-sheet model. The top panel shows ice-sheet volume through time, and situations (A) and (B) represent ice advancing into a topographic basin (A) and ice advancing down a mountain slope (B). Situation (A) is represents rapid non-linear growth in which a small change in climate can have a dramatic effect on ice volume, while situation (B) represents stable growth. [Modified from: Payne and Sugden (1990) Earth Surface Processes and Landforms, 15, figure 6, p. 632]

those centred on topographic ridges. Consequently the location of large topographic basins have a dramatic effect on the rate at which ice sheets grow. In these locations small deteriorations in climate may have a dramatic effect on the size of the ice body.

Once established an ice sheet will continue to grow provided that there is an excess of precipitation over that which the ice sheet can discharge and ablate. Growth will be driven by a variety of positive feedback systems. For example, the Earth's atmosphere is characterised by strong altitudinal gradients in temperature and precipitation. As an ice sheet grows an increasing proportion of its area will lie at more favourable altitudes for accumulation, facilitating further growth. Ice sheets will grow until: (i) their size is limited by the available space, such as the edge of the continental shelf and the presence of deep water; (ii) precipitation starvation sets in when the interior of an ice sheet becomes so removed from sources of precipitation around its margins that its rate of accumulation falls; and (iii) climate changes, reducing accumulation or increasing ablation. Theoretically ice sheets should reach an optimum or equilibrium size for the prevailing climate and topographic location. The ice sheet will exist until some change in this climate occurs to cause it to decay (deglaciation).

Deglaciation may be driven by either a decrease in precipitation and therefore accumulation or an increase in ablation or a combination of both. An increase in ablation may be achieved not only through a rise in air temperature but also through a rise in sea level. Rising sea level may increase the area subject to calving and therefore induce rapid ablation. Rapid calving will increase the discharge of ice towards these margins, a process that may initiate the development of large fast-flowing ice streams draining the ice-sheet centre. It has been suggested that this process may cause the rapid 'draw-down' and deglaciation of an ice sheet. Traditionally there are two models used to explain how deglaciation occurs.

1. Catastrophic down-wasting by areal stagnation. In this model the equilibrium line rises quickly above the ice sheet, depriving it of accumulation. As a consequence down-wasting is rapid and traditionally believed to be associated with little forward flow. This model has been used to explain the rapid decay of mid-latitude ice sheets during the Cenozoic 'Ice Age' and the presence of extensive areas of stagnation-type landforms within the limits of these former ice sheets.

2. Active deglaciation by regular ice-marginal retreat. In this model ice sheets decay in a regular fashion through the contraction and retreat of the ice margin. Forward flow is maintained even as the margin of the glacier retreats. This mechanism can occur both slowly and quickly.

Although widely postulated as a mechanism of glacier decay, areal stagnation is now considered by many geologists to be of only regional or local importance. It is unlikely that ice sheets will be completely deprived of accumulation during decay. More significantly, ablation will always be maximum at the margin and minimum at the ice divide, maintaining the surface ice-sheet gradient, basal shear stress and therefore forward flow. Stagnation on a local or regional scale may occur, however, where very low-angled lobes of ice exist, a situation that may result, for example, after a glacier has surged. For the most part, therefore, glaciers decay actively through the regular retreat of their ice margin. This process may involve a fast, continuous and therefore rapid retreat, or it may be punctuated by a series of still stands, seasonal readvances, surges or larger prolonged readvances. The pattern of deglaciation is therefore rarely a simple and uniform one.

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