Glacier fluctuations in western Scotland

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Quaternarylnternational, Vols 38/39, pp. 137-147, 1997.

j Pergamon PH: S1040-6182(96)00018-3

Copyright© 1996INQUA/ElsevierScienceLtd Printed in Great Britain. All rights reserved. 104(I---6182/97$32.00

GLACIER FLUCTUATIONS IN WESTERN SCOTLAND D o u g l a s I. B e n n

Department of Geography, University of Aberdeen, Aberdeen AB9 2UF, U.K.

Glacier fluctuations between the Last Glacial Maximum (LGM) and the early Holocene are reviewed. At the LGM (< 25 ka BP) most of Scotland was covered by an ice sheet that terminated on the continental shelf, mostly beyond the present coastline. Deglaciation was interrupted by stillstands or readvances in many parts of western Scotland, which may reflect climatic oscillations or glaciodynamic changes that occurred during the transition from marine-based to land-based conditions. By the later part of the Late-glacial Interstade (ca. 13-11 ka BP), glaciers had retreated into restricted areas or had disappeared completely. A major glacier readvance occurred during the Younger Dryas (Loch Lomond) Stade (ca. 11-10 ka BP) when an icefield occupied most of the western Highlands and smaller ice masses developed elsewhere on the Mainland and on some Hebridean islands. Initially, retreat of many Loch Lomond Readvance glaciers was interrupted by stillstands or readvances, which may reflect declining precipitation towards the end of the stade. Final deglaciation was more rapid, and was probably forced by rapidly rising temperatures. Copyright © 1996 INQUA/ElsevierScience Ltd

INTRODUCTION

CLIMATIC BACKGROUND

Western Scotland, located on the eastern fringe of the North Atlantic Ocean, occupies an important location as regards global palaeoclimate. Oceanic circulation in the North Atlantic plays a major role in determining poleward energy transfer in the northern hemisphere, and is thought to be a key factor regulating global climate change (Broeker et al., 1985; Rind et al., 1986; Broeker and Denton, 1990; Lehman and Keigwin, 1992). Accordingly, palaeoclimatic data from western Scotland are important for understanding the timing and magnitude of past climate change, and the mechanisms by which climate signals and oceanic circulation patterns are modified and transferred to adjacent landmasses. The Younger Dryas event, for example, was particularly marked in western Britain, and can be used as a yardstick against which Late-glacial climatic signals from other parts of the world can be measured. Over the last three decades a large amount of research has focused on the record of Late Quaternary environmental change in Scotland, including glaciation, periglaciation, sea-level change and biostratigraphy (for recent reviews see Sutherland, 1991; Boulton et al., 1991; Gray and Coxon, 1991; Sutherland and Gordon, 1993; Ballantyne and Harris, 1994). The purpose of this paper is to review evidence for glacier fluctuations in western Scotland from the LGM to final deglaciation in the early Holocene, and to discuss the possible links between the glacial record and climatic, oceanic and glaciodynamic controls. The area covered spans from Kintyre in the south (55°20'N) to Cape Wrath in the north (58°38'N), and includes the mountains of the mainland coastal fringe, the Hebridean islands to the west, and the Island of Arran in the Firth of Clyde (Fig. 1).

Late Quaternary climatic fluctuations in Western Scotland are intimately associated with southward and northward migrations of the average position of the Oceanic Polar Front (Ruddiman and Mclntyre, 1973, 1981; Lamb, 1979; Ruddiman et al., 1980; Peacock and Harkness, 1990; Ko~ et al., 1993). At present, the Oceanic Polar Front lies well to the north of the British Isles, and western Scotland has a mild, maritime climate, strongly influenced by warm surface water carried from lower latitudes by the North Atlantic Drift. January sea-surface temperatures (at the same latitude as Labrador) are as high as 4°C. Dominantly westerly atmospheric circulation over this warm ocean results in the frequent passage of moist air masses across Scotland, producing a variable climate that is often wet and mild, although snowfall is common in the mountains between November and April. Along much of the west coast, mean July temperature is ca. 14°C and mean January temperature ca. 4°C. On the summit of Ben Nevis, Britain's highest mountain (1343 m), mean annual temperature is fractionally above 0°C, and the absolute minimum recorded temperature is - 1 7 . 3 0 ° C (Manley, 1971a; McConnell, 1988). At sealevel, annual precipitation mostly ranges between 1000 and 2000 mm, although local orographic effects exert a large influence on precipitation, and annual totals of 3000--4000 m m occur on some western mountains. Altitudinal precipitation gradients in excess of 4.5 m m year -1 m -1 have been recorded (Ballantyne and Harris, 1994). At present, no glacier ice or permanent snow exists in Scotland, and semi-permanent snow beds survive only in sheltered locations on the highest mountains, most notably at 1160 m in Observatory Gully on Ben Nevis 137

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FIG. 2. Reconstructed palaeotemperatures for the wannest month (upper curve) and the coldest month (lower curve) for the period 15-8 ka BP, based on radiocarbon-dated coleopteran assemblages in lowland Britain. DS: Dimlington Stade; LI: Late-glacial Interstade; YD: Younger Dryas Stade; H: Holocene. After Atkinson et al. (1987).

the Polar Front, as reconstructed from planktonic assemblages in ocean cores (Ruddiman and McIntyre, 1981).

FIG. 1. Western Scotland, showing the location of places mentioned in the text.

LAST GLACIAL MAXIMUM

(Manley, 1971b; Moran, 1988; Ballantyne and Harris, 1994). Days of snow-lie increase nearly linearly with altitude, indicating that year-round snow cover would be expected to occur at 2285 m, for a level unshaded surface (Lockwood, 1982). This figure does not represent a present-day glacier equilibrium line altitude (ELA), however, because glacier location is strongly influenced by local patterns of accumulation and ablation related to snow-blowing, avalanching and topographic shading. When these factors are taken into account, the threshold glaciation level over northern Britain lies at roughly 1670--1735 m, well above the highest summits (Lockwood, 1982). Although these figures are notional values, and at best give only an approximation of the glaciation limit, they provide benchmarks for comparison with reconstructed Pleistocene glacier ELAs, allowing ELA depression at different times to be estimated. The best available pre-Holocene palaeotemperature estimates for the British Isles are those based on fossil coleopteran assemblages (Fig. 2; Atkinson et al., 1987). These data are for lowland Britain, however, and palaeotemperature estimates based on the dimensions of former glaciers suggest that during the Loch Lomond (Younger Dryas) Stade mean July sea-level temperatures in Western Scotland were ca. 6-7°C, some 3-4°C lower than estimates derived from coleopteran data (Sissons, 1980; Ballantyne, 1989a; Ballantyne and Harris, 1994). Comparative paiaeotemperature data for other periods of the Late-glacial in western Scotland are not available. The timing and magnitude of temperature changes indicated by the coleopteran record are in phase with oscillations of

Radiocarbon and Uranium-series dates from Inchnadamph and the Isle of Lewis (Fig. 1) show that those areas were ice free until ca. 25 ka BP, and that the LGM was therefore after this date (yon Weymarn and Edwards, 1973; Lawson, 1984; Sutherland and Walker, 1984; Atkinson et al., 1986; Sutherland, 1991). Timing of the maximum extent of the ice is not known, although some evidence suggests that it may have been earlier than the date of 18 ka BP for the ice sheet maximum in eastern England (Sutherland, 1984a, 1991). The limits of the ice sheet in western Scotland mostly lay on the continental shelf beyond the present coast (Fig. 3). Submerged morainal banks marking the probable ice limit have been identified on the shelf south of St. Kilda (Selby, 1989; Peacock et al., 1992), but St. Kilda itself appears to have escaped ice sheet glaciation, and supported only a small valley glacier (Sutherland et al., 1984). North of St. Kilda, the limit of grounded ice possibly extended northeast across the shelf, either crossing the present coast at the northern tip of Lewis (Sutherland and Walker, 1984) or passing north of that island (Bowen et al., 1986; Stoker and Holmes, 1991). The vertical dimensions of the last ice sheet have been reconstructed for some areas by mapping periglacial trimlines that delimit the transition between ice-scoured lower ground and periglacially-weathered summits. Recent data from the mainland and the islands of Skye and Harris (Ballantyne, 1990, 1994, this volume; Ballantyne and McCarroll, 1995) yield consistent patterns of ice-surface morphology, which complement information on ice flow directions on lower ground. These results have important implications for modelling studies, which

Glacier Fluctuations in Western Scotland

139

provides a rough figure as the basis for discussion and future research.

ICE SHEET RETREAT

FIG. 3. Reconstructed Last Glacial Maximum ice sheet limits in Great Britain. Based on Bowen et al. (1986) and Hall and Bent (1990).

hitherto have been poorly constrained by geomorphological evidence (e.g. Boulton et al., 1977, 1985). An increasing body of evidence (including trimline altitudes) shows that the last ice sheet consisted of several dispersal centres drained by ice streams. Ice domes existed over the Outer Hebrides, the Inner Hebridean islands of Skye and Mull, and a number of centres on the mainland (e.g. Bailey et al., 1924; Harker, 1901; yon Weymarn, 1979; Peacock, 1984, 1991; Sutherland, 1984a; Thorp, 1987; Lawson, 1990). Patterns of erratic dispersal indicate that the positions of some ice domes and ice streams shifted through time (Sutherland, 1984a; Lawson, 1990). Detailed spatial and temporal patterns of the dynamics of the ice sheet remain unclear, however, and await future research. No empirically-determined estimate has been made of the ELA of the last ice sheet in western Scotland. Indeed, given the considerable uncertainty surrounding the extent and vertical dimensions of the ice sheet in many areas, no such estimate is possible at present. However, a reconstructed ELA is available for the tiny St. Kilda glacier, which is assumed to have existed at the LGM (Sutherland et al., 1984). The inferred value is based on the maximum altitude of lateral moraines of the former glacier, which extend up to 150 m a.s.1. Sutherland (1984b) described a submarine rock platform around St. Kilda with a backing cliff ca. - 1 2 0 m below sea level, and proposed that it marked the Late Devensian sea-level. If correct, this suggests a revised ELA of 270 m above contemporary sea level for the St. Kilda glacier (Sutherland et al., 1984). The reconstructed ELA of the St. Kilda glacier does not, of course, represent a direct estimate of the ELA of the last ice sheet, but it at least

It has been argued that ice sheet deglaciation in Scotland was uninterrupted by major stillstands or readvances, with the exception of the Wester Ross Readvance (Sutherland, 1984a, 1991). The limit of the Wester Ross Readvance is marked by an extensive, sharp-crested end moraine that can be traced from Applecross in the south to the shores of Loch Broom in the north (Fig. 1; Robinson and Ballantyne, 1979; Sissons and Dawson, 1981; Sutherland, 1984a). There is no marked drop in the altitude of raised shorelines at the ice limit, such as would indicate a prolonged stillstand of the ice margin during this period of rapid isostatic uplift. Instead, the marine limit declines inland along the sea-lochs inside the moraine (Sissons and Dawson, 1981), suggesting the gradual opening of calving bays during relative sea-level fall. No direct dates are available for the Wester Ross Readvance, although it predates the onset of organic sedimentation at Loch Droma, on the former ice-shed, which has been dated to 12,810:t:155 BP (Kirk and Godwin, 1963). This date, however, is probably too young, and the Wester Ross Readvance is currently believed to date to around 13.5 ka BP (Ballantyne et al., 1987). A great deal of evidence for pauses or readvances of the ice sheet margin also exists elsewhere on the west coast. Dawson (1982) identified an end moraine complex in central Islay (Fig. 1) which coincides with an eastward drop in the marine limit of ca. 12 m. This evidence was dismissed by Sutherland (1984a) as incomplete and unsubstantiated, but subsequent investigation of the area by A.G. Dawson and the author suggests that the Central lslay Moraine may represent a significant oscillation of the ice sheet margin. At Otter Ferry on the shores of Loch Fyne (Fig. l), the ice margin halted (or readvanced) while local sea-level fell by at least 20 m, an event that has been radiocarbon dated to ca. 13 ka BP (Sutherland, 1984a, 1991). Similarly, in the area between Loch Fyne and Oban, significant falls in relative sea-level occurred while the ice sheet margins retreated by only a few kilometres (Gray and Sutherland, 1977). For the area from Oban to Wester Ross, the relationship between ice-sheet deposits and raised shorelines is less well known. Substantial glacimarine deltas, apparently associated with halts in deglaciation, occur at Loch Don on Mull (Benn and Evans, 1993), near mouths of several mainland sea-lochs, and on the Islands of Skye and Raasay. On Skye, several lines of evidence suggest a stillstand or expansion of local ice after withdrawal of mainland ice, but before the Loch Lomond Stade. Because much of this evidence has not been published elsewhere, it will be described here.

Ice Sheet Retreat Stages on Skye

Evidence for a pre-Loch Lomond Stade ice readvance on Skye is found on the low ground fringing the Cuillin

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D.I. Benn

FIG. 4. Reconstructedice surface contours(in metres) and flowlines for the Last Glacial Maximumfor the Island of Skye and adjacent areas, based on mappingby C.K. Ballantyne and the author. Nunataks shownin black.

Hills in the central part of the island (Figs 4 and 5). The clearest evidence is in Glen Drynoch, where glacitectonically-folded silts and sands crop out in degraded moraine ridges at ca. 15 m O.D. (NG 425308). To the SE of the Cuillin Hills, in Strath Suardal, glacitectonic structures and patterns of erratic dispersal record a reversal of ice flow direction, with westward-flowing

FIG. 5. Provisionalmap of ice sheet stillstand or readvancelimits on the Island of Skye and adjacent Mainland. There is no evidence that the ice limits shownwere contemporaneous.

mainland ice being succeeded by eastward-flowing local ice (Le Coeur and Kuzucuoglu, 1992). In lower Glen Brittle, an abrupt landward drop in the marine limit from ca. 18 m to below 7 m has been interpreted as evidence for a stillstand of local ice during ice sheet deglaciation (Walker et al., 1988). Drops in the marine limit, from ca. 30 to ca. 20 m, also occur in the vicinity of Braes and Broadford in eastern Skye, indicating that an ice lobe remained in the sheltered sounds between Skye, Scalpay and Raasay while relative sea-level fell in the surrounding open waters. The limit of the ice lobe is particularly clear at Braes and on the nearby coast of southern Raasay, where there are large raised glacimarine deltas with surfaces at ca. 30 m a.s.1. Additionally, a readvance of ice between Skye, Scalpay and Raasay is recorded by icethrust subaqueous outwash at Suisnish in southern Raasay (NG 557342). Extensive raised deltas at altitudes of ca. 30 m a.s.1, also occur at Kyleakin in eastern Skye (Walker et al., 1988), and at Plockton on the adjacent Mainland, and may mark the position of the mainland ice sheet at around the time local ice stood at Braes and Broadford. Ballantyne (1988) described boulder ridges and moraines near Dunan, in eastern Skye, which he interpreted as the lateral limit of an ice-sheet readvance. Subsequent work in the area (Benn, 1990, 1991) showed that the features are more likely to be a medial moraine formed between confluent ice streams. Collectively, the above evidence suggests that ice in the Cuillin Hills remained dynamically active after deglaciation of the open waters between Skye and Wester Ross, and the withdrawal of mainland ice from all or most

Glacier Fluctuations in Western Scotland of the island. It is proposed that, during ice-sheet deglaciation, a large calving bay opened up between Skye and the Mainland, separating the Skye accumulation area from Mainland ice. Ice retreat continued until the margins stabilised in sea-lochs or narrow sounds between Skye and adjacent islands. Glacitectonic structures record readvances of the ice margin in some localities. No dating evidence is currently available, and it is uncertain whether the evidence records a single, synchronous readvance of the Cuillin ice centre or localised and asynchronous icefront oscillations. The altitude of raised shorelines on Skye and the mainland, however, suggests that the ice margin positions on Skye may have been broadly contemporaneous with the Wester Ross Readvance.

Causes of Late-glacial Ice Sheet Fluctuations

There is no reason why the Wester Ross Readvance should be granted special status as the only significant interruption of ice sheet retreat in western Scotland. The evidence described above shows that stillstands or readvances occurred along most of the west coast, usually at or near points where the ice sheet margin became grounded above contemporary sea-level. These pauses or oscillations in the ice margin can be explained by one or both of two factors: (1) internal changes in ice sheet dynamics; and (2) climatic forcing. Ice Sheet Dynamics

Many of the retreat stages noted above are located very close to the contemporary sea-level, often at narrow points between islands or constrictions in sea-lochs. Such sites would have acted as pinning points for tidewater outlet glaciers, reducing losses by calving and stabilizing the ice margin, as first proposed by Mercer (1961). Additionally, the withdrawal of ice from deep troughs and sounds would have reduced the importance of ice streams as drainage routes delivering inland ice to calving margins (cf. Hughes, 1992). The reduction of calving losses and ice-stream 'draw-down' would have altered the mass balance of ice-sheet catchments, perhaps allowing periods of positive mass balance until the margins adjusted to terrestrial conditions. In such cases, readvances (such as the Wester Ross Readvance) could occur along parts of the ice sheet margin. Another dynamic mechanism that may account for the expansion of ice in south and east Skye after withdrawal of mainland ice is the removal of constraints imposed by confluent ice. At the ice sheet maximum, eastward and southward flow from the Skye accumulation centre was prevented by the presence of thick ice draining from the Mainland (Fig. 4). The recession of this ice by rapid calving losses would have allowed Skye ice to drain radially away from the high ground. This mechanism, however, could only apply in the case of Skye, and is inappropriate elsewhere. Climatic forcing

It is also possible that pauses or oscillations of the ice

141

sheet margins reflect climatic events. A climatic cause was proposed by Sissons (1981), who suggested that the wester Ross Readvance occurred in response to an increase in snowfall as the average position of the atmospheric Polar Front migrated northward at ca. 13.513 ka BP. Alternatively, intervals of positive ice-sheet mass balance could be due to drops in temperature. Evidence for brief cold episodes superimposed on general climatic warming have been recognised in biostratigraphic records for the Late-glacial period prior to the Loch Lomond Stade, both in Britain and the North Atlantic region as a whole (Atkinson et al., 1987; Walker et al., 1994; Levesque et al., 1993). A single 'revertence episode' is recorded in pollen profiles from the islands of Mull, Skye and Lewis (Lowe and Walker, 1986; Walker and Lowe, 1990; Edwards and Whittington, 1994) and from Oban on the mainland (Tipping, 199l). Currently available dates indicate that each of these revertence episodes occurred within the period ca. 12,300-11,800 BP, apparently too late to correlate with known or inferred ages of the Wester Ross Readvance or the Otter Ferry Stage (see above). A broader perspective on the Late-glacial climate of the North Atlantic region is provided by sea-floor sediments and Greenland ice cores. Both oceanic and ice core records show that the last glaciation was characterised by a series of warm-cold oscillations - - Dansgaard-Oeschger events - - superimposed on longer-term cooling cycles (Dansgaard et al., 1993; Bond et al., 1993). Severe stadial conditions at the end of each long-term cycle culminated in the discharge of large amounts of icebergs (Heinrich events) followed by a rapid return to interstadial conditions (Bond et al., 1992). This sequence of events has been linked to the advance of Laurentide ice streams into the Labrador Sea, resulting in calving along tidewater margins and the delivery of massive volumes of ice into the North Atlantic (Bond et al., 1992, 1993; Andrews et al., 1994). It is not known whether the Dansgaard-Oeschger events reflect climatic forcing at sub-Milankovitch timescales or cyclic internal instabilities of the Laurentide Ice Sheet, although current thinking favours a climatic cause (Bond et al., 1993; Bond and Lotti, 1995). In either case, the reduction of the salinity of the North Atlantic caused by the influx of icebergs was probably sufficient to shut down thermohaline circulation, thus influencing the climate of the region as a whole (Bond et al., 1993; see also Broeker and Denton, 1990). It may be hypothesised, therefore, that the DansgaardOeschger events are present in the Scottish ice sheet record. The most recent prominent Heinrich event (HI) has been dated to ca. 14 ka BP, a time when glacier readvances occurred in north and South America, Scandinavia and New Zealand (Broeker and Denton, 1990; Clapperton, 1993, 1995). It is suggested that this event may correlate with the Wester Ross Readvance and other readvances of the ice sheet margin in western Scotland. A rigorous dating program will be necessary to test this suggestion, and establish the place of Scottish ice sheet oscillations in their dynamic regional context. It is not known whether Scotland was completly

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D.I. Benn

deglaciated prior to the expansion of glaciers in the Loch Lomond Stade. Sissons (1976) argued that, because radiocarbon dates imply that much of Scotland was ice free by ca. 13 ka BP, and coleopteran evidence from lowland Britain indicates warm conditions at that time, the remaining ice may have vanished as early as 12.5 ka BP. Sutherland (1980, 1984a), however, has warned that the basal radiocarbon dates from freshly-deglaciated terrain are not reliable, and must be interpreted with caution. He argued that evidence from Loch Fyne and adjacent areas indicates the survival of Highland ice masses until at least 12.5 ka BP, and probably throughout the Late-glacial Interstade. Ice is also thought to have survived in the Northwest Highlands throughout the interstade (Ballantyne et al., 1987). More data are needed to establish the controls on ice sheet deglaciation in western Scotland. In particular, precise dating of particular ice-margin positions and local sea-level curves are needed to determine their relationship (or lack thereof) to climatic forcing. This will undoubtedly require a large research effort, and the application of modern dating techniques.

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FIG. 6. Loch Lomond Readvance glacier limits in Scotland, based on numerous sources referenced in text. Nunataks in the area of the Highland icefieldare not shown.

THE LOCH LOMOND (YOUNGER DRYAS) STADE The Loch Lomond Stade (Younger Dryas), conventionally dated to 11-10 ka BP, marked a return to severe cold conditions (Fig. 2; Atkinson et al., 1987). The stade is clearly represented in numerous bog and lake cores by a mineral-rich layer containing sparse tundra-type pollen, which contrasts with the organic-rich sediments of the prior Late-glacial Interstade and the succeeding early Holocene (e.g. Walker and Lowe, 1990). The onset of the Loch Lomond Stade has been radiocarbon dated to around 11,300-10,600 BP, and the end dated to 10,200-9900 BP (Walker et al., 1994). This chronology agrees well with limiting dates for glaciation, obtained from organic material underlying and overlying glacial deposits (Sutherland, 1986; Gray and Coxon, 1991). During the Younger Dryas, the Oceanic Polar Front migrated southwards to 35°N, isolating Britain from the warming influence of low-latitude air masses (Ruddiman et al., 1980; Ruddiman and McIntyre, 1981). Estimates based on glacier dimensions of the mean July temperature in western Scotland at that time are 6-7°C (Sissons, 1979a, 1980; Ballantyne, 1989a), comparable with figures of ca. 10°C for England reconstructed from beetle remains (Fig. 2; Atkinson et al., 1987). A great amount of research has focused on mapping the limits of Loch Lomond Stade glaciers in Scotland (Sutherland, 1984a; Gray and Coxon, 1991). The maximum ice limits are marked in many areas by lateral and frontal moraines, ice-contact slopes at the upvalley limit of sandar, or deltas, but in some places the limits are indistinct and have been interpolated (e.g. Thorp, 1986). On some lithologies it is also possible to determine the upper ice limits using trimline evidence (Thorp, 1981; Ballantyne, 1989a). A large icefield, drained by numerous

outlet glaciers, occupied the western Highlands from Loch Lomond in the south to Wester Ross in the north (Fig. 6). The limits of the icefield are well constrained by morphological mapping and biostratigraphic evidence in the areas from the Firth of Clyde to Callander, from Loch Rannoch to the Great Glen, and in Loch Linnhe to the north of Oban (Fig. 6; Thorp, 1986; Walker et al., 1992). Elsewhere on the mainland biostratigraphic control is weak, and the glacier limits are largely based on morphological evidence alone. For the Northwest Highlands, the icefield limits shown in Fig. 6 are those proposed by Bennett and Boulton (1993a, b), and are probably correct to within a few kilometres in most places. In the Southwest Highlands, from the Firth of Clyde to Oban, the limits have not been mapped in detail and are rather speculative. Small independent glaciers occupied cirques around the margins of the main icefield and in the mountains of the far north (Lawson, 1986). Independent ice fields and small valley glaciers occupied the islands of Arran (GemmeU, 1973), Mull (Gray and Brooks, 1972; Benn and Evans, 1993); Rum (Ballantyne and Wain-Hobson, 1980) Skye (Walker et al., 1988; Ballantyne, 1989a; Benn et al., 1992) and Harris (Sutherland, 1984a; C.K. Ballantyne, pers. commun.). Glacier limits in the central part of the Island of Skye are shown in Figs 7 and 8. There was no glacier ice on Islay or Jura, but a fossil rock glacier in the mountains of Jura has been assigned to the Loch Lomond Stade by Dawson (1977). Many outlet glaciers along the west coast terminated in the sea, and their maximum extents and dynamics may have been strongly influenced by calving processes (Greene, 1992). Equilibrium line altitudes have been calculated for many Scottish Loch Lomond Stade glaciers and icefields,

Glacier Fluctuations in Western Scotland

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FIG. 8. Loch Lomond Readvance glacier maxima and subsequent retreat positions in the Cuillin Hills, Skye. 1: Ice margin positions; 2: Interpolated glacier limits; 3: Interpolated glacier retreat positions; 4: Erratic trains; 5: Ground over 300 m (omitted within glacier limits for clarity); 6-9: Pollen sites deglaciated by successive local pollen assemblage zones. Reproduced from: Benn et al. (1992), Journal o f Quaternary Science, 7, 125-144.

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HG. 9. Generalizedcontours(in metres)of reconstructedELAsfor Loch LomondReadvance(YoungerDryas)glaciers in the ScottishHighlands. Modified from Sissons (1980).

The method used in Scotland is that applied by Sissons (1976), which finds the area weighted mean altitude of the reconstructed glacier surface: this assumes a former Accumulation Area Ratio (AAR) of 0.5. This method yields higher reconstructed ELAs than estimates based on AARs of 0.65, as used in other parts of the world (e.g. Meierding, 1982), and was shown to yield slight overestimates of modern glacier ELAs by Sutherland (1984c). Due to uncertainties surrounding the surface form of the Highland icefield in many areas, the most reliable ELAs are for the smaller ice masses of the northern Highlands and the Islands of Skye and Rum (BaUantyne and Wain-Hobson, 1980; Lawson, 1986; Ballantyne, 1989a). Generalised ELAs for Loch Lomond Stade glaciers in the Highlands are shown in Fig. 9. There is a marked rise in ELAs from the western and southern margins of the Highlands, reaching a maximum in the Cairngorms, a pattern that has been attributed to a pronounced precipitation gradient (Sissons, 1980; Tipping, 1985). The decline in precipitation away from the west coast was apparently more marked during the Loch Lomond Stade than at the present time, probably due to the chilling effect of the west Highland icefield on incoming moist air masses, elevating snowfall in the west and creating a precipitation shadow in its lee. The termination of the Loch Lomond Stade in Britain was associated with rapid northward movement of the Oceanic Polar Front and a marked rise in temperature (Coope, 1970, 1977; Ruddiman and McIntyre, 1973, 1981; Bard et al., 1987; Atkinson et al., 1987). Temperature increases in Britain were apparently very rapid, around 1.7-2.8 ° per century (Atldnson et al., 1987). During the 1970s, the rapidity of this environmental change was taken as evidence that deglaciation at the end of the stade was exceedingly rapid, a view reinforced by the then prevalent interpretation of apparently chaotic, 'hummocky' moraines within the limits of many Loch Lomond Stade glaciers as stagnation deposits (e.g. Sissons, 1979b). Collectively, the morphological and biostratigraphical evidence seemed to indicate that the

In recent years, detailed studies of moraine genesis have led to a revision of this model. It has been shown that many areas of so-called 'hummocky moraine' are actually nested lateral and frontal moraines recording oscillations of the ice margins during retreat (Eyles, 1983; Benn, 1992, 1992; Bennett, 1990, 1994; Bennett and Boulton, 1993a, b). Regional mapping of moraine patterns show that deglaciation was interrupted by numerous stillstands and minor readvances, indicating a protracted period of active glacier retreat (Benn et al., 1992; Bennett and Boulton, 1993b). Benn et al. (1992) recognised a two-phase pattern of deglaciation in the CuiUin Hills of Skye, consisting of an initial phase of icemargin oscillation followed by a final phase of more rapid, uninterrupted retreat and localised stagnation (Fig, 8). A biostratigraphic framework for deglaciation, and further evidence for time-transgressive, active glacier retreat was provided by studies of pollen records inside the glacier limits (Benn et al., 1992; cf. Walker and Lowe, 1985). Two models have been proposed to reconcile the evidence for glacier oscillations during deglaciation with the biostratigraphic evidence for rapid temperature increases at the beginning of the Holocene. First, Benn et al. (1992) showed from pollen records that deglaciation on Skye was well advanced before the rapid early Holocene thermal increase, and probably began during the later part of the Loch Lomond Stade. They proposed that deglaciation was initiated by a shift to more arid, cold conditions during the stade, and that only the final phase of uninterrupted glacier retreat reflects increasing temperatures at the beginning of the Holocene. According to this model, therefore, the two-phase deglaciation pattern on Skye reflects a change in forcing mechanisms during glacier retreat, the initial phase driven by a decline in snowfall and reduced accumulation, and the second driven by increased ablation. Support for the idea that the Younger Dryas in Northwest Europe consisted of a moist, cold period followed by a more arid (but still cold) period is provided by Late-glacial pollen records, which sho~ an increase in steppe and halophytic taxa, such as Artemisia, during the stade (Walker and Lowe, 1990; Walker et al., 1994; Benn et al., 1992 and references therein). An alternative view of glacier response to climate change at the end of the Loch Lomond Stade has been proposed by Bennett and Boulton (1993a) and Bennett (1994). They argued that the Highland icefield may have been sufficiently large to dampen the effects of the early Holocene thermal amelioration, and that glaciers could therefore have remained active throughout deglaciation despite increasing regional temperatures. This view is not incompatible with the idea that forcing mechanisms may have changed through time, although it implies that the climatic signal in the moraine record may be less clear than suggested by Benn et al. (1992). It should be noted however, that much of the mapping on which Bennett's and Boulton's views was based was conducted at a

Glacier Fluctuations in Western Scotland regional scale, and may contain many errors in detail (Greene et al., 1994). Moreover, there is at present no evidence to show when retreat of the Highland icefield began relative to the Loch Lomond Stade/Holocene transition, or how the pattern of deglaciation is related to precipitation and temperature changes. Resolution of these problems is an important goal for the future. THE HOLCENE IN WESTERN SCOTLAND There is no evidence for Neoglaciation in Scotland. Sugden (1977) suggested that some cirque glaciers may have formed in the Caimgorm Mountains during the Little Ice Age of the 17th, 18th and early 19th Centuries, but subsequent biostratigraphic studies and radiocarbon dating showed that glaciers have not existed in the Cairngorms since the Younger Dryas (Rapson, 1985). Nor is there any evidence for Little Ice Age glaciers in the western Highlands, where precipitation is higher. A ridge of bouldery debris occurs at over 900 m in the northeast cirque of Ben Nevis, but this has been shown to be an active Holocene snow avalanche impact rampart (Ballantyne, 1989b). Climatic conditions in Scotland during the Little Ice Age were neither severe enough, nor sufficiently prolonged, to permit renewed glacier growth. There is, therefore, a striking contrast between the non-glacial conditions of the late Holocene and the Loch Lomond Stade when substantial icefields and numerous cirque glaciers occupied the Scottish Highlands. This contrast is markedly different from the pattern in certain other montane areas of the world, where Late Holocene moraines lie only a few kilometres from moraines of probable Younger Dryas age. This contrast is partly due to the fact that the Scottish Highlands were simply not high enough to intersect the regional snowline during the Little Ice Age, but probably also reflects differences in global circulation patterns between the Younger Dryas and Late Holocene. During the Holocene, the Oceanic Polar Front in the North Atlantic did not move far enough south to cause sufficient cooling for glacier renewal (Lamb, 1979). FUTURE RESEARCH

Relative and absolute differences in the magnitude of Late-glacial ice advances thoughout the world have the potential to highlight former climatic gradients and circulation patterns, and to provide insight into the factors underlying climatic change. The glacial record in western Scotland is particularly important in this respect due to Scotland's exposure to changes in oceanic and atmospheric circulation in the North Atlantic region. An important outstanding research problem is to determine the extent, timing and significance of oscillations of the ice-sheet margin for the period 25-12 ka BP. In particular, a firm chronology needs to be established for ice-margin oscillations and sea-level change, so that the Scottish record can be set in the wider context of

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ocean-atmosphere-cryosphere changes in the North Atlantic region. In particular, the relationships between ice-margin fluctuations, Dansgaard-Oeschger events and terrestrial biostratigraphic 'revertence' episodes need to be established. A related problem is determining the form and dynamics of the last ice sheet. Trimline mapping is yielding a detailed picture of the ice sheet surface at the LGM, allowing mapping of former ice domes, catchment areas and ice streams. When taken together with geomorphological and sedimentological studies of the former glacier bed, such morphological reconstructions can provide vital input for ice-sheet models by defining basal boundary conditions. Without empirical constraints on ice sheet form and bed characteristics, climatic inferences based on numerical models can be very misleading. Finally, controls on deglaciation during and following the Loch Lomond Stade merit further study. The overall pattem of glacier recession is now fairly well established for much of western Scotland, but the links between glacier behaviour and climatic forcing are far from clear. Establishing these links will require the collection of high-resolution multi-proxy palaeoclimatic records (including coleoptera) combined with accurate dating control on moraine sequences. ACKNOWLEDGEMENTS I am grateful to Professors C.K. Ballantyne and C.M. Clapperton for discussions on the glacial record of Scotland, and particularly for pointing out the possible wider significance of Scottish ice-sheet fluctuations, and their thorough reviews of this paper.

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