3.1. Erosion patterns

The variable impact of glacial erosion across Scotland was identified by Linton (Reference Linton, Miller and Watson1959). Comparisons of maps of non-glacial (Fig. 3) and glacial (Fig. 6) landforms and regolith for the northern Scottish Highlands reveal inverse correlations that can be linked to model results for former glacier basal thermal regimes (Fig. 7). Saprolites, tors and pre-Devensian tills are concentrated in NE Scotland and in the eastern parts of the Northern Highlands; areas which lack roughened or streamlined terrain typical where glacial erosion has been effective. Conversely, these non-glacial features are virtually absent from across the Lewisian lowlands of the Outer Hebrides, apart from N Lewis (Hall Reference Hall, Gilbertson, Kent and Grattan1996), and also from the NW Highlands, with its distinctive cnoc-and-lochain terrains. Similar observations led to the recognition of zones of glacial erosion at regional (Linton Reference Linton1963; Clayton Reference Clayton, Brown and Waters1974; Haynes Reference Haynes and Clapperton1983) and sub-regional (Gordon Reference Gordon1979; Hall Reference Hall1986; Hall & Sugden Reference Hall and Sugden1987) scales across Scotland.

Figure 6 Glacial landscapes and landforms in the northern Scottish Highlands. Glacially eroded depressions from Sutherland & Gordon (Reference Sutherland, Gordon, Gordon and Sutherland1993). Corries from Barr et al. (Reference Barr, Ely, Spagnolo, Clark, Evans, Pellicer, Pellitero and Rea2017). Glacial streamlining and roughening mapped from NEXTMap™ imagery.

Figure 7 Cumulative time that the bed for the last (MIS 3-2) British–Irish Ice Sheet was at pressure melting point (PMP), expressed as a percentage of the total simulation time in experiment E102b2 (Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009). Persistent frozen basal conditions are indicated by black shading (% PMP<2.5 %). Assuming that similar basal temperatures developed beneath Early and Middle Pleistocene ice sheets, comparison with Figures 3 and 6 indicates close correlations between: (i) distributions of areas with persistent cold-based ice and non-glacial landforms; and (ii) areas of more frequent warm-based conditions and landscapes of glacial roughening and streamlining.

The contrasting modes of mountain ice cap and full ice sheet glaciation interpreted from the oxygen isotope record from marine core DSDP 607 (Fig. 1) are also important for understanding patterns of glacial erosion and deposition across Scotland. Because of differences in ice extent, the zones of erosion and deposition beneath mountain ice caps and full ice sheets may not coincide, particularly in areas E of the main ice divide. Glacial landform distribution in western Scotland provides evidence for prolonged erosion by mountain ice caps. Alpine glacial topography is restricted to small areas on Skye and in Lochaber (Haynes Reference Haynes and Clapperton1983). Glacial over-deepened rock basins are distributed mainly in 15–30 km-wide zones on both sides of the main watershed throughout its length (Fig. 2). The basins include radial troughs around Rannoch Moor that were excavated by ice flowing from a large ice dome over the SW Grampian mountains (Linton Reference Linton and Embleton1972). An axial ice-shed zone of low erosion exists along the watershed of the Northern Highlands, bounded by areally scoured landscapes and over-deepened valleys to W and E (Gordon Reference Gordon1979) (Fig. 6). Belts of thick glacial deposits occur around the inner Moray Firth and in the lower Clyde basin that include, at depth, pre-MIS 3 tills (see section 6.2). Landform and sediment distribution appears to conform to the erosional and depositional zones of former mountain ice caps that developed through the Middle and Late Pleistocene. Mountain ice-cap extent may have been strongly influenced by calving at marine limits in the inner Moray Firth, the Forth and Clyde lowlands and within the Inner Hebrides (Sissons Reference Sissons1981).

Areas of ice-roughened and ice-streamlined bedrock also occur far outside the mountain ice cap limits reached during the Loch Lomond Stadial. Ice roughening extends across much of the Outer Hebrides and beyond the present coastline of the Northern Highlands in Sutherland and Easter Ross (Fig. 6). The onset zones for former ice streams in western Scotland also lie mainly outside Loch Lomond Stadial limits and extend across large areas of the sea bed in The Minch and the Sea of the Hebrides (Bradwell et al. Reference Bradwell, Stoker and Krabbendam2008a). Ice streams excavated deep basins in the Mesozoic rocks of The Minch (Sissons Reference Sissons1967) and in the inner Moray Firth (Sutherland & Gordon Reference Sutherland, Gordon, Gordon and Sutherland1993). Streamlined terrains in the Tweed (Everest et al. Reference Everest, Bradwell and Golledge2007) and Forth valleys (Golledge & Stoker Reference Golledge and Stoker2006) and around the Firth of Clyde (Finlayson et al. Reference Finlayson, Merritt, Browne, Merritt, McMillan and Whitbread2010, Reference Finlayson, Fabel, Bradwell and Sugden2014) also continue across adjacent areas of the present sea bed. The main ice sheet depocentres were on the North Atlantic shelf and in the North Sea basin (Holmes Reference Holmes and Gordon1997). Glacial erosion and deposition in areas peripheral to the main ice centres must relate to large ice sheets. Bimodal patterns of glacial erosion and deposition linked to mountain ice cap and ice sheet glaciation have been recognised across Fennoscandia (Kleman et al. 2008) but the smaller, highly dynamic glaciers and ice sheets that formerly covered the more complex topography of Scotland have received less attention (Haynes Reference Haynes1977).

Glacial dissection of mountain areas is strongly developed in the western Highlands (Linton Reference Linton1949; Haynes Reference Haynes1977; Jarman Reference Jarman2007). An absence of inherited cosmogenic nuclides from low elevation sites requires removal of >2.5 m of rock in the last glacial cycle (Fame et al. Reference Fame, Owen, Spotila, Dortch and Caffee2018). In Assynt, rates of valley deepening were rapid, estimated at 2 m/ka during periods of glaciation since 280 ka (Hebdon et al. Reference Hebdon, Atkinson, Lawson and Young1998). The density and interconnectivity of glacial valleys drops towards eastern areas (Haynes Reference Haynes1977) (Fig. 6). Progressive glacial modification of valley systems is evident moving E–W across the central Grampians. Little-modified fluvial valley networks remain in the Tarf basin, but valleys become deeply incised in the Forest of Atholl, with the first stages of glacial breaching of watersheds, and pass westwards at Drumochter into highly dissected mountains with interconnected valley systems and deep glacial breaches and cols (Hall & Jarman Reference Hall, Jarman, Lukas, Merritt and Mitchell2004). Valley-in-valley forms are common in Scotland, products of linear glacial excavation of trenches within broad straths (Hall Reference Hall1991) (Fig. 5). Where glacial valleys are developed from fluvial precursors, as in the Cairngorms (Sugden Reference Sugden1968; Hall & Gillespie Reference Hall and Gillespie2016) and across much of the southern Highland boundary (Linton Reference Linton1940), the timescales for glacial modification likely span the Pleistocene (Fredin et al. Reference Fredin, Bergstrøm, Eilertsen, Hansen, Longva, Nesje and Sveian2013). Large meltwater channels may have been re-occupied in successive glaciations, especially where meltwater flow was constrained by topography in cols, for example the many Nye channels in the eastern Grampians (Clapperton & Sugden Reference Clapperton, Sugden, Gray and Lowe1977), or against hill flanks, for example in the Ochils (Russell Reference Russell1995) and Lammermuirs (Sissons Reference Sissons1961). The presence of pre- MIS 5 and younger deposits in channel floors in Moray, Buchan and Caithness (see section 6) indicates at least localised reoccupation.

Glacial modification of landscapes had a marked vertical component. A morphological gradation, from ice-scoured valley floors to roughened valley flanks and smooth plateaux, is a conspicuous feature of many valley cross profiles in the central and eastern Scottish Highlands. Along the western seaboard, blockfields that predate the last glaciation (Hopkinson & Ballantyne Reference Hopkinson and Ballantyne2014) and which mark zones of very limited glacial erosion are confined to the highest summits (Ballantyne et al. Reference Ballantyne, McCarroll, Nesje and Dahl1997, Reference Ballantyne, McCarroll, Nesje, Dahl and Stone1998). Many blockfields are associated with fragments of older, rounded montane topography, dissected by valleys and ‘cookie-cut' by corries. Only in a few areas, such as Knoydart, does ice-roughening extend to summits and here blockfields are almost absent (Ballantyne Reference Ballantyne2000) (Fig. 3). Blockfields drop in altitude towards eastern Sutherland and Caithness (Ballantyne & Hall Reference Ballantyne and Hall2008; Phillips et al. Reference Phillips and Merritt2008). Blockfields have not been mapped systematically in the eastern Grampians, but tors drop from elevations of 1200–600 m in the Cairngorms to 400 m at the Cabrach, 40 km to the NE (Blyth Reference Blyth1969), and to 100 m in Buchan (Hall Reference Hall2005). When compared to the distribution and height of low-relief upland surfaces lacking in glacial erosion forms (Fig. 3), it is clear that the upper limit of effective glacial erosion also declines from W to E across much of the Highlands. This gradient is similar in trend to that seen in models of basal temperature for the last ice sheet (Fig. 7).

3.2. Erosion depths

Depths of cumulative glacial erosion can be estimated by using non-glacial landforms and land surfaces as reference points and surfaces for erosion. Glacial forms, mainly cnoc-and-lochain terrain, rock basins, valleys and corries, are cut into pre-glacial land surfaces and so, along with any glacial lowering of hilltops and interfluves, represent total glacial erosion through the Pleistocene (Hall et al. Reference Hall, Gillespie, Thomas and Ebert2013b). Minor non-glacial forms and weathering mantles – such as many tors (Hall & Sugden Reference Hall, Sugden and Andre2007) and blockfields (Ballantyne Reference Ballantyne2010b) and some shallow saprolites (Wright Reference Wright1997) – are of Early to Middle Pleistocene age and indicate little or no glacial erosion through later glacial cycles. Non-glacial and glacially roughened slopes have been themselves lowered by weathering and erosion through the Pleistocene (Phillips et al. Reference Phillips, Hall, Mottram, Fifield and Sugden2006; Hopkinson & Ballantyne Reference Hopkinson and Ballantyne2014; Andersen et al. Reference Andersen, Egholm, Knudsen, Linge, Jansen, Pedersen, Nielsen, Tikhomirov, Olsen and Fabel2018).

Increasing glacial erosion depths are evident moving W across the Grampians. In eastern Buchan, where Tertiary gravel deposits survive on ridge tops (Hall et al. Reference Hall, Gilg, Fallick and Merritt2015a), erosion is locally of the order of metres. In the type landscape of selective linear erosion in the Cairngorms (Sugden Reference Sugden1968), glacial erosion of plateau surfaces was negligible (Hall & Glasser Reference Hall and Glasser2003), although non-glacial denudation of bare rock surfaces operated at 2.8–12.0 m/Ma (Phillips et al. Reference Phillips, Hall, Mottram, Fifield and Sugden2006). Most erosion resulted from glacial deepening of valleys by a maximum of 200–350 metres (Sugden Reference Sugden1968; Hall & Gillespie Reference Hall, Binnie, Sugden, Dunai and Wood2016). Estimated average Pleistocene erosion depths across the Dee catchment are estimated at 30–60 m, with a substantial contribution from the removal of saprock and saprolite (Glasser & Hall Reference Glasser and Hall1997). In the western Grampians, inheritance of ¹⁰Be cosmogenic nuclides is highest on ridge tops and declines or is absent at lower elevations. Modelled erosion rates decline eastwards from 40 to 20 m/Ma (Fame et al. Reference Fame, Owen, Spotila, Dortch and Caffee2018).

Glacially streamlined terrain found along parts of the Great Glen, in much of Caithness and in patches along the lower Dee valley towards Aberdeen (Fig. 6) merges laterally with non-streamlined terrain with similar ridge-top elevations, suggesting that erosion has been restricted mainly to the excavation of valleys and depressions. Similarly, in terrain on the Lewisian basement gneiss of the western seaboard, it has long been thought that glacial erosion has been largely restricted to the removal of regolith (Godard Reference Godard1961), with deep erosion of basement only in basins and valleys (Krabbendam & Bradwell Reference Krabbendam and Bradwell2014). On the Ross of Mull (Fig. 8), the cnoc-and-lochain terrain has been interpreted as a glacially stripped and roughened tor field developed on an erosion platform of Pliocene age (Le Coeur Reference Le Coeur1988). Depths of erosion above cnoc summits, however, remain poorly constrained in both streamlined and roughened terrains. The recognition of mega-grooves 0.1–6 km long, 10–100 m wide and 5–15 m deep on the streamlined bed of a former ice stream in Assynt (Bradwell et al. Reference Bradwell, Stoker and Krabbendam2008a) and the floors of The Minch (Bradwell et al. Reference Bradwell, Stoker and Larter2007) and the Sea of the Hebrides (Dove et al. Reference Dove, Arosio, Finlayson, Bradwell and Howe2015) demonstrate that, in zones of fast ice flow, glacial erosion of hard crystalline rock was, at least locally, highly efficient.

Figure 8 Cnoc-and-lochain terrain developed in Caledonian granite, Fionnphort, Ross of Mull: 1 = Cnoc developed in massive granite; 2 = Rock basin excavated in fractured granite; 3 = Major fracture transverse to ice-flow; 4 = Granite monoliths, with >10 m vertical fracture spacing, acting as resistant, stoss-side bastions.

Deep glacial erosion is manifest elsewhere. In the mountains of the western Highlands, glacial erosion has been highly effective, excavating basins and fjords below sea level and dissecting and lowering pre-existing watersheds (Sissons Reference Sissons1967). Glacially polished surfaces above 500 m in Glen Shiel and Glen Nevis show little or no cosmogenic ¹⁰Be inheritance, indicating ≥2.5 m of erosion in the last glacial cycle (Fame et al. Reference Fame, Owen, Spotila, Dortch and Caffee2018). Ice tends to flow towards and stream along pre-existing fluvial valley systems, a confluent flow pattern leading to highly effective linear erosion (Sugden Reference Sugden1968; Nesje & Whillans Reference Nesje and Whillans1994), the main component of the glacial buzz-saw (Hall & Kleman Reference Hall and Kleman2014). Transfluent ice flow in the western Highlands led to breaching of watersheds and to cutting of new valleys (Haynes Reference Haynes1977). Rock basins reach depths of 200–300 m below sea level in lochs Ness and Morar (Sissons Reference Sissons1967), and basins in the Sound of Raasay and Loch Linnhe are of even greater depths (Sissons Reference Sissons1976). In the upper Forth valley, sharp contrasts exist in depths of glacial erosion. Resistant lava plateaux and tapered interfluves developed in softer rocks but somewhat sheltered from glacial erosion in lee locations stand ∼100 metres above the drift-filled floors of the Forth and Teith valleys (Linton Reference Linton1962; Fig. 9). Glacially excavated trenches and rock basins extend below these valley floors to depths below sea level of 100 m E of Stirling and 170 m at Bo'ness (Sissons Reference Sissons1969; Francis et al. Reference Francis, Forsyth, Read and Armstrong1970). Borehole records also show rockhead at 100 m below the Devon valley to the E (Soons Reference Soons1959) In W Fife, the thick Late Carboniferous dolerite sill between the M90 motorway and the Lomond Hills now forms a fragmented scarp that has been breached and lowered by ice flow. Many crag and tail forms in the glacially streamlined terrain of the Central Lowlands are largely erosional features, with crags formed in volcanic plugs and tails formed in sedimentary rock (Burke Reference Burke1969; Sissons Reference Sissons1971; Evans & Hansom Reference Evans and Hansom1996). The difference in elevation of beds on the stoss and lee sides provides a minimum depth of glacial erosion around the crags of 25–100 m (Geikie Reference Geikie1887).

Figure 9 Uneven impact of glacial erosion in the upper Forth valley. Smooth, tapered interfluves (Linton Reference Linton1962) and benches are developed in sandstones of mainly Devonian age (brown dashed lines). Carboniferous lava plateaux, marked by red dashed lines, have been lowered, roughened and weakly streamlined by glacial erosion. The intervening valleys have been over-deepened and thick sediments infill rock basins below the Carse of Stirling (CS) that reach depths of >100 m (Sissons Reference Sissons1967). Abbreviations: GH = Gargunnock Hills. MH = Menteith Hills. OH = Ochil Hills. General direction of ice flow indicated by arrows.

Little attention has been given to the substantial thicknesses of weak Carboniferous to Neogene sedimentary rocks likely removed by glacial erosion from the basins and shelves surrounding Scotland. Major unconformities truncating Early and Middle Pleistocene sediments are mapped in the North Sea (Cameron et al. Reference Cameron, Stoker and Long1987) and on the North Atlantic shelf (Stoker et al. Reference Stoker, Hitchen and Graham1993). Progressively older Cenozoic to Palaeozoic rocks sequences are also truncated to landward (Andrews et al. Reference Andrews, Long, Richards, Thomson, Brown, Chesher and McCormac1990). The mass balance of rock removal and its transfer to the shelf by glacial processes has not been fully quantified, but preliminary investigations suggest that lowering of extensive areas of shelf by glacial erosion was required to produce the large volumes of Pleistocene sediment (Clayton Reference Clayton1996; Glasser & Hall Reference Glasser and Hall1997). Such scalping of inner shelves by glacial erosion and during episodes of low sea level is apparent in the Upper Regional Unconformity off western Norway (Rise et al. Reference Rise, Ottesen, Berg and Lundin2005). Mass transfer from Scotland may have been sufficient to induce isostatic rebound of the land area in response to Pleistocene erosion (Hall & Bishop Reference Hall, Gordon, Whittington, Duller and Heijnis2002) and tilting towards North Sea depocentres (Løseth et al. Reference Løseth, Raulline and Nygård2013) but also awaits further quantification.

3.3. Corries

Corries, or cirques, are striking components of Scottish mountain scenery. Many corries were occupied by small glaciers during the Loch Lomond Stadial, but the small volumes of debris found in end moraines indicate that erosion of the corrie basins has extended over a much longer period (Sissons Reference Sissons1976). Corrie dimensions are moderate, with headwalls generally ranging in height from 150 m to 450 m and with width and length rarely exceeding 1 km (Gordon Reference Gordon1977). Such dimensions, when averaged over Pleistocene time, indicate quite slow erosion, but this is potentially misleading because cirque erosion rates were highly variable through the Pleistocene (Crest et al. Reference Crest, Delmas, Braucher, Gunnell and Calvet2017); and, also, many corries may have been occupied by independent glaciers for only short periods before being overwhelmed by ice sheet advance (Richardson & Holmund Reference Richardson and Holmund1996). Corries with glacially-rounded upper slopes are common in the SW Grampians (Sissons Reference Sissons1976), close to a major centre of ice sheet accumulation. Here, and perhaps elsewhere, the main phases of corrie erosion may have occurred during the many stadial periods of the Early Pleistocene, with long periods of burial beneath the ice sheets of the Middle and Late Pleistocene and more limited erosion in the short intervening phases of mountain glaciation. Some support for this hypothesis is provided by model results, with an apparently short duration of warm-based conditions at high elevations across Scotland beneath the last ice sheet (Fig. 7). The frequency of rock-slope failures and failure scarps within corries in NW Scotland suggests that cirque enlargement during successive glacial-interglacial cycles involved deepening by glacial erosion, headwall retreat by rock-slope failure and removal of RSF debris during subsequent glaciation (Ballantyne Reference Ballantyne, Wilson, Gheorghiu and Rodés2013; Cave & Ballantyne Reference Cave and Ballantyne2016).

Corrie distribution in Scotland is a function of climate, topography and time. Corrie floor elevations are low along the western seaboard but rise across the main Highland watershed towards the eastern Grampians. Snowfall from moisture-bearing North Atlantic air masses fed the corrie glaciers, with precipitation decreasing markedly with distance from the present shoreline (Barr et al. Reference Barr, Ely, Spagnolo, Clark, Evans, Pellicer, Pellitero and Rea2017). Corries are typical of high mountain scenery in Scotland, indicating the effect of decreasing air temperature with altitude. Most corries face NE and N, the sector of least direct solar radiation and hence least ablation (Sissons Reference Sissons1967). Topographic controls on corrie distribution exist at the regional and local scales (Gordon Reference Gordon1977). Corries are largely absent from Caithness, eastern Sutherland, Moray and the central and eastern Southern Uplands because few hills stood above 700 m a.s.l. before glaciation. East of the main watershed in northern Scotland, corries are developed mainly in valley heads on the northern flanks of glens draining to the Moray Firth (Fig. 6). Corrie floor altitudes vary in elevation by 350 m and 400 m in parts of the Grampian Highlands and Cairngorms (Sissons Reference Sissons1976). This broad range reflects local geological and topographical controls (Gordon Reference Gordon1977) but corries with floors at different altitudes may also have been occupied at different times, a reminder that several generations of corries may exist within individual mountain groups (Godard Reference Godard1965; Sugden Reference Sugden1969; Rudberg Reference Rudberg1994).

3.4. Controls on glacial erosion

Complex interactions between geological, topographic, climatic and glaciological factors influenced patterns of ice flow and depths of glacial erosion across Scotland.

• • Ice streams developed on sticky, rigid beds only where lubricated by meltwater (Krabbendam & Bradwell Reference Krabbendam and Bradwell2013). On soft beds, with slippery, fine-grained sedimentary rocks or deformable unconsolidated basal substrates (Boulton & Jones Reference Boulton and Jones1979; Golledge & Stoker Reference Golledge and Stoker2006), basal sliding was facilitated, leading to draw-down of ice towards the Mesozoic basins of the Moray Firth, Minch and Forth Approaches. Thin-skinned glacitectonics triggered by high pore-water pressures contributed to erosion through the thrusting and transport of glacial rafts of rock and sediment (Phillips & Merritt Reference Phillips and Merritt2008; Phillips et al. Reference Phillips, Everest and Reeves2013). Such processes are highly effective along the margins of ice streams (Patterson Reference Patterson1998) and help to account for the high frequency of glacial rafts found along the shores of the inner Moray Firth (Merritt et al. Reference Merritt, Auton, Connell, Hall and Peacock2003) and the Firths of Clyde (Merritt et al. Reference Merritt, Akhurst, Wilkinson, Riding, Phillips, Smith, Finlayson and Dean2014) and Forth (Kendall & Bailey Reference Kendall and Bailey1908).

• • The pre-glacial topography of the western Highlands was already at high elevation and fluvially dissected at the start of the Pleistocene and so provided gradients at the beds of alpine glaciers and ice sheets that were steeper than on the plateaux and lowlands more typical of the topography of eastern areas. Broad straths and lowland corridors established before the Pleistocene also guided ice flow towards major ice streams (Fig. 2).

• • Climatic gradients and shadow effects across Scotland were consistently steeper during the cold stages of the Pleistocene (Golledge & Hubbard Reference Golledge and Hubbard2005) reflecting reduced North Atlantic circulation and atmospheric moisture flux. As per today though, precipitation and temperatures remained higher over the western highlands (Barr et al. Reference Barr, Ely, Spagnolo, Clark, Evans, Pellicer, Pellitero and Rea2017), leading to higher rates of snow accumulation, warmer englacial and subglacial temperatures and enhanced ice flow (Boulton & Hagdorn Reference Boulton and Hagdorn2006; Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009; Patton et al. Reference Patton, Hubbard, Andreassen, Winsborrow and Stroeven2016).

• • Steeper relief, higher accumulation rates and enhanced englacial and subglacial temperatures sustained fast-flowing, wet-based marine outlet glaciers which were active over longer periods in the west. In the east, slower, cold-based glaciers and ice caps generally persisted, with relatively short-lived, or no, phases of fast flow (Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009; Fig. 7). Vertical contrasts in glacial erosion in glaciated uplands reflect greater ice thickness and flow velocity at low points in the landscapes (Glasser Reference Glasser1995) and the fundamental importance of englacial thermal boundaries in providing sharp upper limits to effective glacial erosion (Sugden Reference Sugden1968; Hall & Glasser Reference Hall and Glasser2003; Fabel et al. Reference Fabel, Ballantyne and Xu2012).

Ice extent and dynamics varied in space and time as ice fronts advanced and retreated (Haynes Reference Haynes and Clapperton1983) and as basal thermal conditions changed (Gordon Reference Gordon1979). Simulations for the last ice sheet indicate the importance of long binge and short (10–30 % relative duration) purge cycles. Under climatic cooling, periods of ice sheet inception and thickening were predominantly characterised by extensive, protective cold-based glaciers (Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009). During purges driven by climatic amelioration, warm-based ice became extensive, with ice streams propagating from the shelf edge towards fast-flow onset zones on the present land area (Boulton & Hagdorn Reference Boulton and Hagdorn2006; Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009). Fast flow driven by basal sliding promoted short episodes of rapid glacial erosion (Näslund et al. Reference Näslund, Rodhe, Fastook and Holmlund2003; Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009; Kucher et al. Reference Kuchar, Milne, Hubbard, Patton, Bradley, Shennan and Edwards2012), implying that roughened and streamlined rock landscapes are mainly products of relatively brief phases of erosion when ice streams and outlet glaciers were active, leading to rapid interior drawdown, dynamic thinning and ice mass wastage (i.e., Kucher et al. Reference Kuchar, Milne, Hubbard, Patton, Bradley, Shennan and Edwards2012). The total duration of Pleistocene ice sheet cover was limited (∼500 ka in build-up centres; Fig. 1) and was further reduced at increasing distances from the main ice accumulation centres (Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009). Hence, whilst rates of glacial erosion may be high, the rather brief duration over which efficient glacial erosion operated through the Pleistocene reduced its impact on the landscape.

Scotland's maritime location and its 0.8–1.4 km-high mountains create near-optimal conditions for the rapid development of mountain ice caps in response to moderate (3–6·5°C) falls in summer temperatures (Payne & Sugden Reference Payne and Sugden1990; Golledge et al. Reference Golledge, Hubbard and Sugden2008). The δ¹⁸0 isotope record (Fig. 1) suggests that during the Early Pleistocene, the dominant mode of glaciation was by mountain ice caps. During the Middle and Late Pleistocene, however, mountain ice caps existed mainly during the slow build-up to full ice sheet cover. One model for mountain ice cap and alpine glacier growth in Scotland is provided by the Loch Lomond Stadial, when conditions of rapid temperature fall, steep precipitation gradients to north and east and persistent sea ice brought rapid build-up of ice caps in western Scotland (Hubbard Reference Hubbard1999; Golledge et al. Reference Golledge, Hubbard and Bradwell2010). Markedly different configurations may have developed under conditions with more gradual temperature fluctuations, such as in the period of increasing ice volume at 38–32 ka, when an extensive ice cap is reconstructed over the Cairngorm mountains in mathematical simulations (Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009). Full ice sheet growth required greater temperature falls (>6·5°C) (Payne & Sugden Reference Payne and Sugden1990) of longer duration (3–10 ka) (Hubbard et al. Reference Hubbard, Bradwell, Golledge, Hall, Patton, Sugden, Cooper and Stoker2009) (Fig. 1).

Large ice sheets developed over Scotland from ∼1.2 Ma onwards and extended onto the North Atlantic shelf (Fig. 2). As there is no reliable evidence for the existence of nunataks above the last ice sheet at its maximum (Fabel et al. Reference Fabel, Ballantyne and Xu2012), and models of the last ice sheet indicate ice surface elevation above 1.5 km (Kuchar et al. Reference Kuchar, Milne, Hubbard, Patton, Bradley, Shennan and Edwards2012), then it is likely that earlier ice sheets of equivalent or greater extent also covered all of the Scottish mainland, including its highest summits. Streamlining of bedrock (Bradwell et al. Reference Bradwell, Stoker and Krabbendam2008a) and excavation of glacial valleys (Graf Reference Graf1970) contributed to the progressive adaptation of topography and the glacier bed towards more efficient ice flow through the Pleistocene. Middle Pleistocene ice sheets terminated near to the North Atlantic shelf edge (Stoker et al. Reference Stoker, Hitchen and Graham1993) or against or near to the western edge of the FIS in the North Sea (Fig. 2). This apparent overstep of Early Pleistocene ice limits, also seen in the southern British Isles (Lee et al. Reference Lee, Woods and Moorlock2015), implies that the British–Irish Ice Sheet, unlike the Laurentide Ice Sheet (Balco & Rovey Reference Balco and Rovey2010), did not approach or reach its maximum limits in the earliest Pleistocene. The position of the British Isles below 60°N, where insolation intensity is less sensitive to changes in Earth's obliquity, may have prevented the build-up of large ice sheets under the 40 ka cycles of the Early Pleistocene (Huybers & Tziperman Reference Huybers and Tziperman2008).

The varied glacial landscapes of Scotland can be seen as products of average conditions with two end members. Landscapes of little or no glacial erosion occur where cold-based conditions were established repeatedly and for long periods beneath successive ice sheets and ice caps through the Pleistocene (Fig. 7). Landscapes of deep and extensive glacial erosion are found where streams of fast, warm-based ice flow developed during successive glacial cycles. Most of the landscapes of Scotland, however, fall between these two extremes, with ice sheet models indicating long periods of cold-based ice cover and only brief phases of fast ice flow (Fig. 7). The morphological impact of glacial erosion on these intermediate landscapes was weak (Fig. 10).

Figure 10 View N up Helmsdale towards Griam Mor and Griam Beag in Sutherland, isolated hill masses developed on Devonian conglomerate. Extensive low-relief surfaces have been only weakly dissected by fluvial and glacial erosion during the Pleistocene.

4.1. Chemical weathering before and during the Pleistocene

The record of chemical weathering at and beneath the landsurface in Scotland is long and complex. Whilst the dating of saprolites that lack stratigraphic context is challenging (Hall Reference Hall, Dawson, Jones, Small and Soulsby1993a), most weathering profiles and soils appear to have developed in the near-surface during the Neogene and Pleistocene. A clay-rich, kaolinitic and base-poor saprolite type is found in E Buchan as part of a Palaeogene landscape that includes the Buchan Gravels Formation (Hall Reference Hall1985; Hall et al. Reference Hall, Gilg, Fallick and Merritt2015a). Similar geochemically evolved saprolites found elsewhere in Scotland are probably either exhumed from beneath sedimentary cover (Humphries Reference Humphries1961; Parnell et al. Reference Parnell, Baron and Boyce2000) or hydrothermal in origin, the latter including kaolins on Shetland (May & Phemister Reference May and Phemister1968) and at Pittodrie, Bennachie (Hall Reference Hall, Gordon and Sutherland1993d; Hall et al. Reference Hall, Gilg, Fallick and Merritt2015a). Elsewhere, only gruss-type weathering profiles are found, with high sand and low fines contents and retention of weathering-susceptible primary minerals such as Ca-feldspar and biotite (Hall Reference Hall1985). Gruss localities are concentrated in NE Scotland, the Caithness–Sutherland border and NW Lewis, but are almost certainly more widespread than shown in Figure 3 because large parts of Highland Scotland have not yet been surveyed for weathering remnants. The mineralogy of North Sea sediments (Dypvik Reference Dypvik1983; Nielsen et al. Reference Nielsen, Rasmussen and Thyberg2015) indicates that the gruss weathering type started to develop from the Late Miocene onwards in response to climate cooling. Deep gruss profiles (Fig. 11), which can extend to depths of many tens of metres (Hall et al. Reference Hall, Mellor and Wilson1989b), probably developed mainly in the Pliocene period (5.333–2.588 Ma). Shallow weathering profiles, with depths of a few metres, almost certainly continued to develop in temperate interglacial and cool interstadials during the Early Pleistocene (Hall et al. Reference Hall, Mellor and Wilson1989b). We may suppose this because the total duration of such ice-free intervals was long (Fig. 1), rates of gruss development can be fast (Dethier & Lazarus Reference Dethier and Lazarus2006) and grussification of clasts is typical in pre-MIS 5 weathered till units in Scotland (Connell et al. Reference Connell, Edwards and Hall1982; Bloodworth Reference Bloodworth, Auton, Firth and Merritt1990) and northern England (Boardman Reference Boardman and Boardman1985). Where Middle to Late Pleistocene till units incorporate and overlie weathered rock, the period of weathering must predate deposition of the tills (Connell et al. Reference Connell, Edwards and Hall1982; Gordon Reference Gordon, Gordon and Sutherland1993a; Hall Reference Hall, Gordon and Sutherland1993b) (Fig. 11). On rocks with high susceptibility to granular disintegration or oxidation, near-surface weathering may be entirely of post-glacial origin. For example, the rapid breakdown of some quartz dolerites on exposure to air (Orr Reference Orr1979) may account partly for the extensive development of onion-skin weathering in Carboniferous sills and dykes in the Midland Valley (Henderson Reference Henderson1893; Walker Reference Walker1935).

Figure 11 Gruss pocket at Cairngall Quarry, Mintlaw, Buchan, developed in medium-grained biotite granite below a thin till cover. Gruss weathering profiles in this area reach known depths of >60 m (Hall Reference Hall1985).

Prior to glaciation, deep weathering covers were widespread, but weathering penetrated most deeply in fractured, sheared and geochemically-susceptible rocks beneath basin and valley floors, whilst adjacent hills were developed in resistant and largely fresh rock types such as quartzite, acid granites and slate (Godard Reference Godard1962; Hall Reference Hall1986). One effect of glacial and non-glacial erosion was to progressively thin and finally remove gruss from wide areas through the Pleistocene (Clark & Pollard Reference Clark and Pollard1998; Krabbendam & Bradwell Reference Krabbendam and Bradwell2014; Hall et al. Reference Hall, Sarala and Ebert2015b). Particularly in western areas of high glacial erosion, the very limited survival of gruss means that tills of the last, Late Devensian glaciation are composed of clasts and matrix material released by the breakage and comminution of fresh rock debris. In contrast, in eastern areas where saprolite pockets remain, the tills may incorporate large amounts of former saprolite and soil materials (Wilson et al. Reference Wilson, Bain and Duthie1984).

4.2. Mountain-top detritus

Mountain-top detritus (MTD) refers to frost-riven regolith of diverse character found on Scottish hills (Ballantyne Reference Ballantyne, McCarroll, Nesje, Dahl and Stone1998). The character of MTD is closely controlled by rock type and fracturing and by the time available for its formation. Three main types of MTD have been recognised (Ballantyne Reference Ballantyne1984): (i) openwork blockfields; (ii) clast-rich, sandy diamictons; and (iii) clast-rich, silty and fine sand diamictons. Blockfields are typical of rocks with widely spaced fractures, such as quartzite and granite; sandy diamictons are developed mainly on sandstones and psammites; and silty diamictons are common on greywackes and mica schists. MTD generally forms only shallow regolith, reaching depths of 0.1–1.6 m, with surface concentrations of larger clasts and more matrix fines at depth. Clast angularity, local derivation and vertical sorting indicate an origin by frost riving and frost heaving, processes that have operated effectively at high elevations since the last deglaciation and probably earlier also (Ballantyne & Harris Reference Ballantyne and Harris1994).

On many Scottish mountain summits, there is evidence that MTD predates the last ice sheet. Where glacial erosion by the last ice sheet has left bare bedrock surfaces, MTD occurs only on densely-fractured rocks (Fig. 12). MTD has not reformed on more massive rock types, indicating that it is slow to develop on these rocks (Fabel et al. Reference Fabel, Ballantyne and Xu2012). Where glacial erosion has been ineffective in removing regolith, MTD may contain assemblages of clay minerals, including gibbsite and kaolinite, that are distinct from those found on the same rock types on lower slopes (Ballantyne Reference Ballantyne1994a). Many hill summits in the British Isles and Ireland show a distinct upslope limit to MTD (Fig. 3). Earlier interpretations of this limit as a glacial trimline at the upper limit of the last ice sheet (Ballantyne Reference Ballantyne, McCarroll, Nesje and Dahl1997) have given way in recent years to a recognition that the limit marks an englacial boundary between warm- and cold-based ice within the last ice sheet (Fabel et al. Reference Fabel, Ballantyne and Xu2012; McCarroll Reference McCarroll2016). Detailed mapping has shown that the elevation of this trimline drops away from the main watershed towards the Outer Hebrides and to the inner Moray Firth (Ballantyne Reference Ballantyne2010a).

Figure 12 Mountain top detritus on the Red Cuillin, Skye. Fine- to medium-grained granite is broken into a thin cover of MTD with many small, angular clasts in a granular sand matrix. The summit was probably over-topped by the last ice sheet but exposed as a nunatak in the Loch Lomond Stadial (Small et al. Reference Small, Rinterknecht, Austin, Fabel and Miguens-Rodriguez2012). Estimated erosion rates of 30–40 mm/ka are based on ¹⁰Be cosmogenic inventories (Fame et al. Reference Fame, Owen, Spotila, Dortch and Caffee2018).

The age of the MTD found above trimlines on mountain plateaux is poorly constrained. MTD may have thickened in the post-glacial period, but the main phase of its formation generally predates the last ice sheet (Hopkinson & Ballantyne Reference Hopkinson and Ballantyne2014). Development of MTD requires that summits and plateaux are free of permanent snow and glacier ice and exposed to frost action and weathering. Hence, blockfields may have formed or thickened during MIS 5-3 or during earlier interstadial and interglacial phases of the Early and Middle Pleistocene. On relict, non-glacial plateau surfaces in northern Scandinavia, ¹⁰Be and ²⁶Al inventories are consistent with formation of the present MTD layer through the Middle and Late Pleistocene (Goodfellow et al. Reference Goodfellow, Stroeven, Fabel, Fredin, Derron, Bintanja and Caffee2014a). In western Norway, cosmogenic nuclides indicate low erosion rates (4–6 m/Ma) (Andersen et al. Reference Andersen, Egholm, Knudsen, Linge, Jansen, Pedersen, Nielsen, Tikhomirov, Olsen and Fabel2018). The residence times of regolith on the Cairngorm plateaux are short, with maximum erosion rates for exposed rock surfaces ranging from 2.8 to 12.0 m/Ma, and tor emergence rates after stripping of regolith at 11–35 m/Ma (Phillips et al. Reference Phillips, Hall, Mottram, Fifield and Sugden2006). There is also evidence, however, of low Middle to Late Pleistocene erosion rates. Cosmogenic exposure ages indicate that weathering pits deeper than 5–10 cm predate the last glacial cycle (Hall & Phillips Reference Hall and Phillips2006b). Weathering pits up to 20 cm deep occur on large tabular blocks set within MTD. Blockfields that have been disrupted locally by ice flow and pass downslope into stripped, joint-bounded granite bedrock surfaces display weathering pits up to 12 cm deep. The weathering pits indicate formation and stripping of MTD in the Cairngorms in or before MIS 6. Evidence of a longer weathering history for plateau surfaces is also given by pockets of thick gruss-type saprolite of likely Pliocene to Early Pleistocene age found on high ground in the central and eastern Grampians (Hall & Mellor Reference Hall and Mellor1988; Phillips et al. Reference Phillips, Hall, Mottram, Fifield and Sugden2006; Hall Reference Hall, Sugden and Andre2007). However, for thin blankets of MTD to be maintained at even low erosion rates, formation of new detritus from frost weathering of bedrock is required over 100 ka timescales (Marquette et al. Reference Marquette, Gray, Gosse, Courchesne, Stockli, Macpherson and Finkel2004). Ballantyne (Reference Ballantyne2010b) has proposed a dynamic model of long-term blockfield evolution in which Neogene regolith cover was gradually removed by subaerial surface lowering during ice-free periods and progressively replaced by the products of frost-wedging under periglacial conditions, with most existing MTD formed entirely within the last 135 ka (Hopkinson & Ballantyne Reference Hopkinson and Ballantyne2014). Where MTD is older, the uppermost trimlines mapped on Scottish mountains may predate the last ice sheet, a suggestion consistent with the more restricted extent of the last ice sheet compared to earlier ice sheets on the North Atlantic shelf (Ballantyne et al. Reference Ballantyne, Fabel, Gheorghiu, Rodés, Shanks and Xu2017)

4.3. Tors

Tors are small rock knobs produced by long-term differential weathering and erosion. At lower latitudes than Scotland, tors may be exhumed from thick saprolites, but almost all Scottish tors sit on hill and ridge tops and have no close association with deep weathering. Rock compartments with low fracture spacing have emerged as tors through the repeated formation and stripping of thin regolith similar in properties to that which exists on surrounding slopes today (Hall & Sugden Reference Hall, Sugden and Andre2007). This debris mantle, in areas such as the Cairngorms, is of complex origin, with elements that derive from hydrothermal alteration (Hall & Gillespie Reference Hall and Gillespie2016), multiple phases of chemical weathering (Hall & Mellor Reference Hall and Mellor1988), granular disintegration due to stress release (Phillips et al. Reference Phillips, Hall, Mottram, Fifield and Sugden2006) and frost weathering (Ballantyne Reference Ballantyne1994b). Studies in other humid, unglaciated granite terrains indicate rates of regolith formation of >5 m/Ma (Bierman Reference Bierman1994; Duxbury et al. Reference Duxbury, Bierman, Portenga, Pavich, Southworth and Freeman2015), comparable to known rates of MTD production. Such rates, when combined with evidence of significant Holocene chemical weathering in the Cairngorms (Soulsby et al. Reference Soulsby, Chen, Ferrier, Helliwell, Jenkins and Harriman1998) and Late Devensian frost weathering and mass movement (Ballantyne Reference Ballantyne1994b), suggest that much of the 0–2 m of regolith found around tors today has formed mainly in the Late Pleistocene. Thin regolith may be largely renewed in ice-free intervals during each interglacial- glacial cycle.

In areas of Pleistocene glaciation, tors are absent from zones of areal scouring but may be present in landscapes of alpine glaciation (Small et al. Reference Small, Anderson, Repka and Finkel1997), selective linear erosion (Sugden Reference Sugden1968; Sugden & Watts Reference Sugden and Watts1977) and in areas of little or no erosion (André Reference André2004). Tor landscapes in Scotland (Fig. 3) also include isolated hill masses around which the last British ice sheet streamed at the Last Glacial Maximum, for example the hills of South Uist (Ballantyne & Hallam Reference Ballantyne and Hallam2001), northern Arran (Godard Reference Godard1969) (Fig. 13) and Ben Loyal in Sutherland (Godard Reference Godard1965). Like blockfields, the existence of tors was formerly attributed to these areas standing as nunataks above the last ice sheet (Godard Reference Godard1965). The existence of glacial erratics and glacially displaced tor blocks, however, leaves no doubt that these and other summits were ice-covered. It appears that the tors have survived beneath protective covers provided by cold-based glaciers (Sugden Reference Sugden1968). Where no sliding of ice takes place across the glacier bed, then delicate features including tors and regolith may be preserved (Rapp Reference Rapp, McCann and Ford1996). Evidence of localised glacial transport from the tor site is provided by displaced tor blocks and boulder trains (Hall & Phillips Reference Hall and Phillips2006a). Displaced tor blocks may reach sizes of 100 m³ (Fig. 14) and show the distinctive weathered surfaces typical of Cairngorm tor summits (Hall & Phillips Reference Hall and Phillips2006b). In the Cairngorms, such criteria can be used to trace the movement of former tor blocks, allowing distances and directions of transport to be identified. Tors formed of smaller blocks may be partly or wholly demolished to form boulder trains (Hall & Phillips Reference Hall and Phillips2006a).

Figure 13 Tors developed in the Northern Arran Granite emplaced at ∼60 Ma (Dickin et al. Reference Dickin, Moorbath and Welke1981). Note the truncation of inclined, sub-parallel joints by the glacial slopes of Glen Sannox, the exposures of thin granular regolith and the weathering pits on granite surfaces.

Figure 14 Glacially-transported tor block, eastern Ben Avon in the eastern Cairngorms. The tor in the background has lost superstructure to glacial entrainment. Extensive spreads of sandy MTD, with small blocks, are developed on the Cairngorm Granite.

Tors in glaciated regions have been referred to as pre-glacial landforms (e.g., Sugden Reference Sugden1968). The term is used in two senses – pre-Pleistocene or pre-glaciation, the latter where the tor predates one or more phases of glaciation. The advent of cosmogenic isotope analysis has allowed constraints to be placed on the exposure ages and erosion rates for rock surfaces on tors. Cosmogenic isotope data confirm that Cairngorm and Caithness tors are older than the last interglacial (Phillips et al. Reference Phillips, Hall, Mottram, Fifield and Sugden2006; Ballantyne & Hall Reference Ballantyne and Hall2008). The oldest Cairngorm tor surface at Clach na Gnùis has a minimum exposure and burial age of 675 ka (Phillips et al. Reference Phillips, Hall, Mottram, Fifield and Sugden2006). This tor summit carries weathering pits 1 m deep and is unmodified by glacial erosion, yet the accumulated nuclide inventory on its top indicates that this tor is not a pre-Pleistocene landform. Existing tors in the Cairngorms are dynamic landforms which have attained their present forms through the Middle Pleistocene.

Where tors or even tor roots survive in formerly glaciated areas, then the total glacial erosion since tor emergence has been confined to modification or removal of the protuberance, without significant lowering of surrounding surfaces. In glaciated tor fields, such as the Cairngorms (Gordon Reference Gordon, Gordon and Sutherland1993b), this means that glacial erosion has been only a few metres (Hall & Phillips Reference Hall and Phillips2006a). On spurs that merge down slope into larger roche moutonnée forms, the survival of tor roots identifies a zone of very limited glacial erosion on the upper part of the spur (Sugden et al. Reference Sugden, Glasser and Clapperton1992). Tors are therefore useful indicators of terrain that has experienced minimal glacial erosion. In such terrain, known erosion rates indicate that the contribution of erosion from non-glacial processes exceeds that from glacial processes through the Pleistocene. In the Cairngorms, at least three separate phases of glacial modification of tors are recognised on the basis of exposure ages (Phillips et al. Reference Phillips, Hall, Mottram, Fifield and Sugden2006). On these and other granite hills in Scotland, the granite surfaces of tor plinths carry weathering pits >10 cm deep, formed during or before the last interglacial, requiring that tor removal by glacial processes occurred in MIS 6 or earlier (Hall & Phillips Reference Hall and Phillips2006b).

4.4. Exfoliation or sheet joints in granite terrain

Exfoliation or sheet joints are joint sets with orientations subparallel to the present or former ground surface that occur at depths of 1–100 m in granite intrusions (Ziegler et al. Reference Ziegler, Loew and Moore2013). Exfoliation or sheeting develops where high maximum principal compressive stresses up to a few tens of MPa exist at shallow depths, oriented subparallel to the ground surface, and exceed the least surface-normal principal stress, a product of overburden thickness including topographic effects (Holzhausen Reference Holzhausen1989). Extension and opening of exfoliation joints in the near-surface is a response to topographic curvature (Martel Reference Martel2017). Exfoliation joints guide weathering and erosion and have been identified in Scotland as an important control on the morphology of non-glacial forms, including domes (Hall et al. Reference Hall, Gillespie, Thomas and Ebert2013b), tors (Goodfellow et al. Reference Goodfellow, Skelton, Martel, Stroeven, Jansson and Hättestrand2014b) and blockfields (Hall & Glasser Reference Hall and Glasser2003), and glacial forms, including valleys (Glasser Reference Glasser1997), cirques (Haynes Reference Haynes1968) and roches moutonnées (Gordon Reference Gordon1981; Sugden et al. Reference Sugden, Glasser and Clapperton1992). In areas where Pleistocene glacial erosion has led to changes in topography, distinct sets of exfoliation joints can be related to palaeo-topography (Ziegler et al. Reference Ziegler, Loew and Moore2013). In the Cairngorms, at least two generations of sheet joints are recognised oriented sub-parallel to the slopes of plateaux and to glacial valleys and cirques (Glasser Reference Glasser1997). Here and in other granite terrains in Scotland (Godard Reference Godard1969) (Fig. 13), sub-horizontal joint sets are also present that are not related to present slopes and so may relate to pre-glacial relief development.

6.1. The Kirkhill sequence

The Pleistocene sequence in the Kirkhill area (Hall & Jarvis Reference Hall, Jarvis, Gordon and Sutherland1993) is unique in the British Isles terrestrial record in providing a detailed record of at least three complex stadial–interstadial–interglacial cycles (Fig. 16).

• • In the oldest cycle believed to date to MIS 8 (245–303 ka, Merritt et al. Reference Merritt, Auton, Connell, Hall and Peacock2003), glaciation and the deposition of till was possibly succeeded by a phase of weathering, prior to erosion of the till surface, perhaps by the meltwater that deposited the overlying glacifluvial and glacideltaic gravels and sands. After deglaciation of the site, periglacial conditions were established, with frost shattering of bedrock and solifluction and later deposition of fluvial gravels. Following a phase of erosion, a podzolic soil formed under humid temperate conditions on the gravel surface.

• • The second cycle commenced with a climatic deterioration marked by development of arctic soil features and the onset of gelifluction. This was followed by glaciation of the site, with deposition of till. Subsequently, a further soil developed on this till surface and its pedogenic characteristics imply a return to interglacial conditions. This cycle is assigned to MIS 6 – MIS 5e (see below).

• • The youngest cycle (Devensian) opened with deposition of a sequence of periglacial deposits, the subsequent development of permafrost, marked by ice-wedge casts, and the eventual deposition of three further till units. Subsequent deglaciation involved minor glacifluvial deposition and was accompanied, or succeeded, by a phase, or phases, of periglacial activity until soil formation began at the start of the Holocene interglacial.

Figure 16 Kirkhill and Leys schematic Middle to Late Pleistocene stratigraphy (after Merritt et al. Reference Merritt, Auton, Connell, Hall and Peacock2003). British Geological Survey © UKRI 2018.

The oldest event known in the Kirkhill/Leys area is the cutting of a linear depression into weathered basic igneous rock at Leys Quarry. The Leys Till Formation, containing quartzite erratics glacially transported from the W, was laid down on its floor. The till contains grussified basic igneous clasts and appears to have been weathered after deposition (although local incorporation of previously weathered bedrock and subsequent limited glacial transport also occurred). The overlying Denend Gravel Formation comprises 3–5 m of coarse, felsite-dominated, pebble to boulder gravel, with interstratified sand units. The Denend Gravel was deposited by meltwater moving towards the W and SW, part of a valley sandur extending from an ice margin lying relatively close to, and E of, Leys. At Leys Quarry, high-angle, glacideltaic foresets are preserved, indicating the ponding of a glacial lake in the Ugie valley (Fig. 17) at this time. It appears that ice blocked the valley E of Leys, though it is unclear if this was ice flowing from the Moray Firth and curving S, or ice having flowed N up the western North Sea margin. The set of channels at Kirkhill Quarry in which these gravels and younger sediments accumulated may have been cut by meltwater flow at this time or earlier. Post-depositional disturbance of the gravels is widespread, with folding, faulting and development of wedge structures as the result of melt-out of glacier ice blocks originally buried within the gravels (Hall & Connell Reference Hall and Connell1986).

Figure 17 SW face of Leys Quarry in the 1990s showing high-angle, glacideltaic foreset gravels of the ?MIS 8 Denend Gravel Formation. Flow of water was towards the W and SW. Note extensive Fe and Mn staining and local cementation of the sediments. Excavations below this gravel unit exposed the Leys Till resting on bedrock.

At Kirkhill, an angular felsite rubble up to 2 m thick rests on channel floors. Clasts in the rubble show a well-developed downslope fabric, indicating transport by avalanche or gelifluction. Silt cappings indicate incipient pedogenesis (Connell & Romans Reference Connell, Romans and Hall1984). This Kirkton Gelifluctate Bed is, in turn, overlain by the Pitscow Sand and Gravel Formation, with up to 4 m of mainly horizontally stratified sands, with gravel lags marking former channel floors. This unit is interpreted as a periglacial fluvial or proglacial glacifluvial deposit. Cross-stratification points to transport from the E. Syn-depositional ice-wedge casts indicate deposition of the basal beds under permafrost conditions. The declining numbers of angular clasts in the upper Pitscow beds at Kirkhill, together with the absence of ice-wedge casts and other periglacial features, may indicate the subsequent amelioration of climate.

The first cycle ended with an important phase of pedogenesis, with formation of the Kirkhill Palaeosol Bed on the eroded upper surface of the Pitscow Sand and Gravel (Fig. 18). A conspicuous bleached horizon, up to 19 cm thick with bleached and softened felsite clasts, was found throughout Kirkhill Quarry, and in the NE face of Leys Quarry. This represents the former Ea horizon of a podzolic soil, with a lower mottled Bs horizon locally enriched in organic carbon and free iron as a result of pedogenic translocation. An iron pan is also found towards the base of the profile (Connell et al. Reference Connell, Edwards and Hall1982). Micromorphological analysis of the palaeosol has also revealed structures typical of soils of arctic environments (Connell & Romans Reference Connell, Hall, Romans and Hall1984). Superimposition of periglacial soil features on an earlier interglacial soil horizon seems to have produced a composite soil horizon. At both Kirkhill and Leys, Fe and Mn staining associated with localised cementation occurs extensively to depths of several metres in underlying gravels and sands (Fig. 18). These indicators of former waterlogging or fluctuating groundwater levels contrast sharply with the leached horizon represented by the podzol, though the timing and number of redox cycles involved in Fe and Mn mobilisation is uncertain. At Kirkhill, the truncated Kirkhill Palaeosol Bed is overlain by thin, poorly stratified and weakly organic sands, the Swineden Sand Bed. These sands probably represent slope-wash deposits and are penetrated by a series of frost-cracks (Connell et al. Reference Connell, Hall, Romans and Hall1984). Beneath the sands lies a 1–5 cm-thick bed of black to brown, weakly laminated organic mud that drapes the truncated palaeosol in the E face. The organic mud contains pollen of Poaceae, together with a marked arboreal component of mainly Pinus and Alnus, together with charcoal. The overlying sands show a reduction in arboreal pollen and an increase in grasses and Calluna, possibly reflecting the establishment of an open, treeless environment (Connell et al. Reference Connell, Edwards and Hall1982). Sampling of an equivalent organic sequence in the W face at Kirkhill suggests that two components exist in the pollen spectra (Lowe Reference Lowe1984). Open, grassland types represent pollen contemporaneous with sedimentation of the sands. Recycled pollen, with a significant arboreal component, is perhaps derived from older soil horizons at the site. The stratigraphical position of the palaeosol indicates that it developed during a pre-Ipswichian interglacial, probably MIS 7 (245–186 ka; Merritt et al. Reference Merritt, Auton, Connell, Hall and Peacock2003).

Figure 18 South Face 1 of Kirkhill Quarry in the late 1970s: 1 = floor of meltwater channel cut in a felsite dyke; 2 = Pitscow Sand and Gravel Formation; 3 = Kirkhill Palaeosol Bed; 4 = Camphill Gelifluctate Bed; 5 = Rottenhill Till; 6 = Corsend Gelifluctate Bed; 7 = Hythie Till.

The start of the second cycle is seen at Kirkhill Quarry, where up to 2.5 m of periglacial mass movement deposits, named the Camphill Gelifluctate Bed, are developed above the organic Swineden Sand Bed. The bed includes features indicating multiple phases of periglacial activity, including episodes of gelifluction, cryoturbation and development of an ice-wedge polygonal network. A luminescence date of 142±19 ka BP (Duller et al. Reference Duller, Wintle and Hall1995), if correct, falls within MIS 6. The West Leys Sand and Gravel Formation rests in shallow channels scoured into this periglacial land surface. These thin, probably glacifluvial gravels and sands show planar cross-beds, indicating water flow from the NW, and probably represent outwash or subglacial meltwater deposits. The overlying Rottenhill Till Formation overlies with a weak unconformity these earlier deposits (Fig. 19). Its clast lithology (including boulders of Strichen granite) and fabric indicate deposition by ice moving from the NW. Whilst only three sites are known where the Rotten Hill till is found, all show NW–SE ice flow. It is possible that this flow direction was being “steered” by ice in the Moray Firth impinging on Buchan, much as it did during both early and late stages of the Middle–Late Devensian glaciation (Merritt et al. Reference Merritt, Connell and Hall2017). After deglaciation, the Fernieslack Palaeosol Bed developed on the surface of the underlying Rottenhill Till Formation. This second major phase of soil development is marked by grussification of basic igneous and granite clasts, soil horizon development and clay mineral weathering and translocation. Soil characteristics are typical of the B- and C-horizons of gleyed brown earths developed on similar parent materials in eastern Scotland during the Holocene (Connell & Romans Reference Connell, Romans and Hall1984) and are in contrast to the podzolic soil (Kirkhill Paleosol Bed) which developed on the more acidic, free draining, parent materials available at the site during the MIS 7 interglacial. Support for this interpretation of the Fernieslack Palaeosol is provided by the presence of Alnus, with Betula and coryloid grains from a single sample in the profile (Connell et al. Reference Connell, Edwards and Hall1982). On the simplest interpretation, and in the light of the stratigraphically lower luminescence date, the Fernieslack Soil probably formed during the Ipswichian Interglacial (MIS 5e).

Figure 19 Kirkhill Quarry North Face in the 1980s: 1 = podzolic Kirkhill Palaeosol Bed; 2 = unconformable, sub-horizontal, base of the brown Rottenhill Till (within which the Fernieslacks Palaeosol Bed developed); 3 = the black Corse Diamicton (incorporating glacitectonically deformed pale coloured sand) overlying the Rottenhill Till; 4 = Hythie Till. The section is approximately 15 m high.

The latest (Devensian) cycle began with cryoturbation and gelifluction, followed by incipient pedogenesis and ice-wedge growth recorded by successive units within the Corsend Gelifluctate Bed. Advance of glacier ice occurred whilst ground ice persisted (Connell et al. Reference Connell, Hall, Romans and Hall1984). The Corse Diamicton Formation was first exposed at Kirkhill as a large mass of dark grey clay diamicton consisting largely of reworked Late Jurassic/Early Cretaceous mudstone. The overlying Hythie Till Formation has a clast lithology dominated by psammite and pelitic schist, with basic igneous rocks and felsite. A clast of rhomb porphyry typical of those found in the Oslo Graben was recovered from this unit and a second clast was recovered from talus below the NE face at Leys. The stratigraphic position of Scandinavian erratics in Buchan, and their presence in tills of inland origin, indicate that transport of these rock types into the inner Moray Firth probably occurred before MIS 5 (Hall & Connell Reference Hall, Connell, Ehlers, Gibbard and Rose1991), perhaps in the Anglian/Elsterian (MIS 12) (Merritt et al. Reference Merritt, Auton, Connell, Hall and Peacock2003). The dominance of quartzitic and basic igneous rocks in the Hythie Till Formation indicates a western provenance. The unit can be correlated on the basis of stratigraphy and lithology with tills along the valleys of the North and South Ugie Water, which also have strong W–E fabrics (Hall & Connell Reference Hall, Connell, Ehlers, Gibbard and Rose1991). The succeeding East Leys Till Formation is a very dark grey, mud-rich diamicton. Clast lithologies are dominated by gneissose quartzite and psammite, with chalk, pelites and reddish brown (possibly Devonian) sandstone. Striated shell fragments appear below a near-surface Holocene weathering horizon and include specimens of Macoma sp. and Astarte sp. The matrix contains reworked Late Jurassic/Early Cretaceous dinoflagellate cysts, Mesozoic and Tertiary pollen, and Quaternary foraminifera. Together these organic remains indicate derivation of the matrix from Mesozoic mudstone, Palaeogene sediments and Quaternary glacimarine and marine muds in the Moray Firth basin. Southward transport by ice from the Moray Firth is supported by the presence of chalk clasts and by the transport of dark gneissose rocks comparable to the Inzie Head Gneiss of the Fraserburgh area. The East Leys Till represents the final phase of glacial deposition recorded at the site and is succeeded by thin glacifluvial sands and gravels. At Leys, there is evidence of a late phase of permafrost conditions from involutions with stone pillars, near-surface concentrations of frost-heaved clasts, erected and frost-cracked pebbles and associated development of ice-wedge casts (Hall & Connell Reference Hall and Connell1986).

The Kirkhill/Leys sequence provides wider insights into the sequence of environmental change in Scotland because of its peripheral position in the lowlands of Buchan. In the last glaciation, ice flow into this area came via three pathways: the inner Moray Firth, the western North Sea and the eastern Grampians. In the two early cycles of the Kirkhill sequence, only the Denend Gravel appears to relate to ice blocking the North Sea coast in a manner similar to the ponding of Glacial Lake Ugie in the Late Devensian (Merritt et al. Reference Merritt, Connell and Hall2017). The earlier Leys Till and the later Rottenhill Till were deposited by eastern Grampian ice moving eastward toward the North Sea, a pattern of ice flow similar to that for the Hythie Till around the maximum of the Late Devensian. Each of these tills sits with only weak unconformity on older deposits indicating that, in the last three glacial cycles, eastern Grampian ice was effectively non-erosive in this area. The consequent layer cake stratigraphy includes evidence for two pre-Holocene interglacial phases, in which significant weathering and soil development took place. Cold conditions were established in Buchan in at least two separate phases before the last interglacial, with ice wedges marking intervals with permafrost development (Fig. 16). With the three younger units also formed under periglacial conditions, it is clear that periglacial processes operated repeatedly in the lowlands of eastern Scotland through the Middle and Late Pleistocene (Connell & Hall Reference Connell, Hall and Boardman1987). The Devensian sediments in the Kirkhill and Leys area are thin, with older sediments taking up most of the accommodation space within the once 10–15 m-deep channels at these sites. Monitoring of gas pipeline trenches in the 1980s showed that Devensian glacial deposits are also thin across much of central Buchan yet still provide evidence of multiple events from attenuated stratigraphic units (Merritt et al. Reference Merritt, Auton, Connell, Hall and Peacock2003).

6.2. Deposits elsewhere in Scotland that predate MIS 5

At a few, widely scattered sites across Scotland, organic deposits or weathering horizons dated or attributed to MIS 5 are underlain by older sediments (Fig. 15). Till units found low in stratigraphic sections have also been assigned to MIS 6 (Saalian) or older glaciations. On the plain of Caithness, old glacial deposits sit in lows within glacially streamlined terrain, indicating that bedrock erosion largely predates the last glacial cycle (Hall & Riding Reference Hall and Riding2016).

On Shetland, at Fugla Ness, a till lies below a peat layer that probably formed in the last interglacial. This till has a strong NW–SE fabric (Hall et al. Reference Hall, Gordon, Whittington, Duller and Heijnis2002), a similar direction to movement of ice within the Shetland ice cap during the last glaciation (Hall Reference Hall2013). In Caithness, tills deposited by ice moving out of the inner Moray Firth and the Northern Highlands occur below multiple sequences referred to the Late Devensian glaciation (Gordon Reference Gordon, Gordon and Sutherland1993d; Hall & Riding Reference Hall and Riding2016). Amino acid racemisation ratios for marine shells now incorporated in Late Devensian tills in Caithness and Orkney suggest that the oldest included shells date from MIS 7 and 9, respectively (Bowen & Sykes Reference Bowen and Sykes1988). These ages imply reworking of shells from former marine or glacimarine sediments in the inner Moray Firth that included beds of MIS 6 age or older (Hall & Riding Reference Hall and Riding2016).

On northern Lewis, a raised shore platform is found at 6–12 m O.D. (von Weymarn Reference von Weymarn1979; Gordon Reference Gordon, Gordon and Sutherland1993f). Weathered till occurs in pockets on its surface and contains erratic clasts of Torridonian sandstone and Cambrian quartzite, indicating ice movement from the Scottish Mainland (Peacock Reference Peacock1984) during MIS 6 or earlier. This rock platform and one at a similar elevation on Barra and Vatersay (Selby Reference Selby1987) are overlain by beach gravels that probably date from MIS 5 high sea levels (Hall Reference Hall, Gilbertson, Kent and Grattan1996). In Assynt, a cave at Creag nan Uamh has yielded Uranium-series disequilibrium ages for speleothems of 122+/−12 ka, 143+13/−16 ka, 181+24/−18 ka and 192+53/−39 ka, indicating the establishment of ice-free conditions during the last interglacial and MIS 6 (Lawson Reference Lawson, Gordon and Sutherland1993).

East of Inverness, organic deposits and palaeosol/weathering horizons of MIS 5 age have been described from Allt Odhar and Dalcharn. The weathered, sandstone-rich Dearg Till Formation at Dalcharn is older than the Devensian as it underlies an interglacial palaeosol and deeply weathered Craig an Daimh Gravel (Walker et al. Reference Walker, Merritt, Auton, Coope, Field, Heijnis and Taylor1992; Merritt & Auton Reference Merritt, Auton, Merritt, Auton and Phillips2017). The Moy Peat Bed and underlying gravel at the Allt Odhar site also overlie an older sandstone-rich diamict, the Suidheig Till Formation. Another weathered, sandstone-rich diamict, the Cassie Till Formation, occurs nearby at the base of the sequence at Clava. These older tills are presently exposed within several valleys draining towards the River Nairn (Fletcher et al. Reference Fletcher, Auton, Highton, Merritt, Robertson and Rawlin1996) (Fig. 20). In lower Strath Spey, the podzolic Teindland Palaeosol, probably of last interglacial age, is underlain by beds of sand, gravel and till (Fig. 21). Erratic clasts in the Red Burn Till below the palaeosol indicate ice movement from the NW (Hall et al. Reference Hall, Whittington, Duller and Jarvis1995). Each of these underlying tills probably relates to MIS 6 glaciation, as does weathered till found near Portsoy (Gordon Reference Gordon, Gordon and Sutherland1993a; Peacock & Merritt Reference Peacock and Merritt2000). These sequences hold an important record of Middle and Late Pleistocene environmental change preserved along the southern margin of ice streams moving towards and out of the inner Moray Firth (Merritt et al. Reference Merritt, Auton, Connell, Hall and Peacock2003).

Figure 20 Longitudinal profile of the Allt Carn a'Ghranndaich valley upstream from Clava, E of Inverness, showing multiple, buried and locally weathered Middle Pleistocene till and gravel units (after Fletcher et al. Reference Fletcher, Auton, Highton, Merritt, Robertson and Rawlin1996). British Geological Survey © UKRI 2018.

Figure 21 Teindland Quarry in 2000: 1 = Teindland Lower Sand; 2 = Teindland Buried Soil (MIS 5e); 3 = Teindland Upper Sand; 4 = truncated and glacitectonically sheared units beneath overlying Teindland Till.

In the central Grampians, glacilacustrine sediments have been overridden by later ice movements and glacitectonically disturbed (Merritt Reference Merritt, Lukas, Merritt and Mitchell2004; Smith et al. Reference Smith, Merritt, Leslie, Krabbendam and Stephenson2011). These deposits locally rest on the Pattack Till Formation, typically a pale to moderate yellowish brown diamict that includes boulders of grey porphyritic granodiorite sourced from the SW (McMillan et al. Reference McMillan, Hamblin and Merritt2011). When compared to the longer sequences found in the Inverness area, where intervening interglacial and interstadial deposits have been located, the colour, iron staining and weathered condition of many clasts in the Pattack Till suggest that it has been exposed to full interglacial conditions and is pre-Devensian in age. The unit is overlain locally by heavily iron-stained, gravelly deposits that were probably laid down as moraines and ice-proximal fans during retreat of the penultimate ice sheet. This area is transitional between terrains of selective linear glacial erosion, with survival of deep weathering on plateaux (Hall & Mellor Reference Hall and Mellor1988), and more extensively ice-moulded topography where fast-ice flow was directed along the main straths (Smith et al. Reference Smith, Merritt, Leslie, Krabbendam and Stephenson2011). The preservation of pre-MIS 5 glacigenic sediments in Glen Pattack on the margin of a zone of ice streaming suggests that older sediments may occur in similar situations elsewhere in the Spey catchment.

The most widespread occurrences of Middle Pleistocene glacial deposits occur in Buchan (Hall & Connell Reference Hall, Connell, Ehlers, Gibbard and Rose1991). Deeply buried tills are recorded in numerous boreholes in the Ellon (Merritt Reference Merritt1981) and Peterhead (McMillan & Aitken Reference McMillan and Aitken1981) areas. Around Kirkhill, the distinctive Fernieslack Palaeosol and the Rottenhill Till (see above) have been traced over a wide area (Hall et al. Reference Hall, Mellor and Wilson1989a). At two sites W of Peterhead, Toddlehills and Savock, complex glacigenic sequences are preserved (Gemmell et al. Reference Gemmell, Murray and Connell2007; Connell Reference Connell2015). Both sites lie adjacent to the S–N-trending Aldie–Laeca and Dens channel systems (Merritt et al. Reference Merritt, Auton, Connell, Hall and Peacock2003, map 7; Fig. 13) that were covered by eastward advancing ice early during the late Middle–Late Devensian glaciation (Merritt et al. Reference Merritt, Connell and Hall2017). Correlatives of the Rottenhill Till Formation (dated to MIS 6 at the Kirkhill/Leys sites) have been recorded at both sites (Figs 22, 23) and are overlain by fluvial sands at Toddlehills with a mean OSL age of 85 ka (Gemmell et al. Reference Gemmell, Murray and Connell2007), again consistent with deposition of the till during MIS 6. Later, coastal ice advancing northward skirted the area but delivered meltwater northward through the channels into Glacial Lake Ugie. The sequences at Toddlehills escaped erosion at that time as it was on the western flank of the West Dens channel.

Figure 22 Toddlehills Quarry in 2006, N face. Section is ca.6 m high. Weathered bedrock (1) is overlain by a till unit (2), correlated with the MIS 6 Rottenhill Till Formation at Kirkhill. OSL dated (mean 85 ka) sands are obscured by made ground in mid-section (3). Late Devensian Hythie Till Formation can be seen higher in the section beneath vegetated made ground (4).

Figure 23 Savock Quarry in 2014. Section is ca.5 m high: 1 = Middle Pleistocene sand and gravel channel fill; 2 = brown, possibly weathered till correlated with the MIS 6 Rottenhill Till Formation at Toddlehills, Leys and Kirkhill; 3 = upper clast-rich till correlated with the MIS 3-2 Hythie Till.

Highly mottled, red mud was exposed in a temporary pit [NGR NK 033409] on the north slope of the Moss of Cruden in 1984. A sample from its less weathered base contained palynomorphs typical of the Kimmeridge Clay and also Quaternary forms (L. A. Riley pers. comm. 1984), indicating an early phase of glacial transport from the Moray Firth. North of Ellon (Fig. 15), the Pitlurg Till was deposited by ice moving S along the present North Sea coast from the Moray Firth. The age of this unit is uncertain, but amino acid ratios from included shells indicate that the till cannot be older than MIS 6. The Bellscamphie Till is the oldest unit identified in the Ellon area. It was deposited by ice moving from the NW or W, probably prior to MIS 5 (Hall & Jarvis Reference Hall and Jarvis1995). In western Buchan, the Crossbrae Farm Peat bed of probable MIS 5a or 5c age overlies the Crossbrae Till, likely to be of MIS 6 age (Whittington et al. Reference Whittington, Connell, Coope, Edwards, Hall, Hulme and Jarvis1998).

Further south, old tills may be present at Aberdeen (Hall & Connell Reference Hall, Connell, Ehlers, Gibbard and Rose1991), a grey till probably derived from the south (Bremner Reference Bremner1934) and a weathered brown till derived from the west (Synge Reference Synge1956, Reference Synge1963). Most of these tills likely date from MIS 6 and indicate a similar interplay of ice masses in Buchan, as occurred in the last glaciation. At Nigg Bay, immediately south of Aberdeen city, a complex glacigenic sequence is exposed in the southern cliffs (Gordon Reference Gordon, Gordon and Sutherland1993e). The upper two tills are believed to date to ice advances during the Late Devensian. At the base of the exposed sequence, gravels and sands (the Ness Sand and Gravel Member) have recently provided an OSL age estimate of ∼63 ka (Gemmell et al. Reference Gemmell, Murray and Connell2007), suggesting they relate to outwash deposition during an MIS 4 glaciation. Onshore and offshore borehole data indicate that a palaeo-valley form exists immediately N of the cliff section that has been eroded down to −30 m O.D. (Merritt at al. Reference Merritt, Auton, Connell, Hall and Peacock2003, fig. 46). It is unclear if this palaeo-channel represents a former course of the River Dee cut during lowered sea level events or has been partially excavated by subglacial processes. It is also unclear if this older valley preserves sediments that pre-date MIS 4. At Ellon, N of Aberdeen, borehole evidence also indicates a rock-cut channel of the River Ythan descending to at least -20 m O.D. (Merritt Reference Merritt1981) and, again, it is possible that this deep channel preserves pre-Devensian sediments.

South of Stonehaven, near Inverbervie, the Benholm Clay Formation is another old glacigenic sediment, laid down before peat formation in MIS 5 (Fig. 24) (Auton et al. Reference Auton, Gordon, Merritt and Walker2000). The Benholm Clay includes striated and polished marine shell fragments, Carboniferous, Early and Late Cretaceous and Palaeogene microfossils, and clasts of limestone, coal, chalk and flint derived from offshore in the North Sea basin (Auton et al. Reference Auton, Gordon, Merritt and Walker2000). It was laid down by ice that moved onshore during a post-Anglian and pre-Devensian glacial episode. Amino acid (D/L) ratios of ca.0.34 have been obtained from fragments of the marine mollusc Arctica islandica, which indicate that they are of MIS 9 age, or older (Auton et al. Reference Auton, Gordon, Merritt and Walker2000). The clay was possibly derived from overridden sediments at the base of the Coal Pit Formation, or might represent the onshore feather edge of a till occurring in the Fisher Formation, generally considered to be of Saalian (MIS 6–10) age (Gatliff et al. Reference Gatliff, Richards, Smith, Graham, McCormac, Smith, Long, Cameron, Evans, Stevenson, Bulat and Ritchie1994).

Figure 24 Sheared lens of the Burn of Benholm Peat Bed within olive grey shelly clay and diamicton (Benholm Clay Formation) at Burn of Benholm. The shelly deposits are overlain by red-brown diamicton of the Late Devensian Mill of Forest Till Formation.

In the lower Clyde and the upper Forth lowlands, the Geological Memoirs refer to borehole records that show thicknesses of Pleistocene sediment that locally exceed 100 m (Clough et al. Reference Clough, Hinxman, Wilson, Crampton, Wright, Bailey, Anderson, Carruthers and Grabham1925; Francis et al. Reference Francis, Forsyth, Read and Armstrong1970; Paterson et al. Reference Paterson, Hall, Stephenson and Forsyth1990; Forsyth et al. Reference Forsyth, Hall and McMillan1996; Hall et al. Reference Hall and Mellor1998). Only the upper part of these sediment accumulations is of Late Devensian age (Browne & McMillan Reference Browne and McMillan1989). Tills of likely MIS 4 age, the Ballieston Till Formation in the Clyde basin, and the Littlestone Till Formation in Ayrshire, occur widely (Finlayson et al. Reference Finlayson, Merritt, Browne, Merritt, McMillan and Whitbread2010), but the stratigraphy and age of older sediments below these tills remain poorly known (Sutherland Reference Sutherland and Walker1984).

6.3. Significance of the Middle Pleistocene terrestrial stratigraphic record

The Pleistocene stratigraphical record in Scotland includes glacigenic deposits and periglacial features that represent cold stages back at least to MIS 8, together with terrestrial sediments with a floral, faunal or pedological record of interstadial and interglacial conditions. Whilst Middle Pleistocene sediments are sparse or absent across wide areas of Scotland and are poorly dated generally, the stratigraphic record is important for the direct information it can provide on the pattern and sequence of environmental change on the land area. The increasingly detailed and nuanced interpretations of palaeoenvironmental proxies for the modern, post-glacial and late glacial periods in Scotland can be applied also to the fragmentary evidence of earlier periods. Periglacial sediments also provide a wealth of data on palaeo-temperatures and palaeo-precipitation (Ballantyne & Harris Reference Ballantyne and Harris1994) and constrain former ice extent and thickness (Fabel et al. Reference Fabel, Ballantyne and Xu2012). Marine sediments and landforms from MIS 5 and perhaps earlier periods indicate former sea levels and wave environments (Smith et al. Reference Smith, Barlow, Bradley, Firth, Hall, Jordan and Long2019).

Glacial deposits older than MIS 5 allow constraints to be placed on former ice extent, dynamics and flow paths that can be compared with those of the last glacial cycle. The basal sediments of the Kirkhill sequence represent an interglacial-cold interstadial- stadial cycle that probably predates MIS 7. The sequence includes glacial and glacifluvial deposits that derive both from inland and from the North Sea coast. Patterns of ice flow in eastern Buchan were similar to those in later phases in MIS 6 and MISs 3-2 (Merritt et al. Reference Merritt, Connell and Hall2017). In MIS 6, ice in northern Shetland was moving towards the NW, as in the Shetland ice cap in MIS 3-2 (Hall Reference Hall2013). Cnoc-and-lochain terrain had already formed on the Lewisian gneisses (Chapelhowe Reference Chapelhowe1965). In Caithness, till units record movements of inland and Moray Firth ice. The tills sit in bedrock depressions that are part of the streamlined bedrock terrain of the plain of Caithness and indicate that streamlining of bedrock was present before MIS 5 (Hall Reference Hall2013). In NW Lewis, the presence of Torridonian and Cambrian clasts in a weathered till indicate the operation of the Minch Ice Stream in MIS 6. Carry of these and other erratics further S may indicate that the mainland ice sheet crossed parts of the Outer Hebrides in MIS 6 (Flinn Reference Flinn1978a). Across Moray and into lower Strath Spey, till fabrics and erratics for likely MIS 6 tills indicate similar pattern of ice flow as for MIS 3-2 (Merritt et al. Reference Merritt, Connell, Hall, Peacock, Merritt, Connell and Bridgland2000) In Buchan, MIS 6 tills appear to have all been derived from inland. Extensive W–E flow is indicated that is consistent with granitic, schistose, basic igneous and red sandstone clast assemblages found in a Saalian till in the Fladen area derived from NE Scotland (Sejrup et al. Reference Sejrup, Aarseth, Ellingsen, Lovlie, Reither, Bent, Brigham-Grette, Jansen, Larsen and Stoker1987). Combined evidence from blockfields and tors and glacial sediments indicates that the configuration, thermal regime and ice flow during MIS 6 were broadly comparable to those of the last ice sheet.