Sedimentology and luminescence ages of Glacial Lake Humber deposits in the central Vale of York

Proceedings of the Geologists’ Association 120 (2009) 209–222

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Sedimentology and luminescence ages of Glacial Lake Humber deposits in the central Vale of York Della K. Murton a,*, Steven M. Pawley b,1, Julian B. Murton c a

School of Earth and Ocean Sciences, Cardiff University, Main Building, Park Place, Cardiff CF10 3YE, UK Department of Geography, Royal Holloway, University of London, Egham, Surrey TW20 0EX, UK c Permafrost Laboratory, Department of Geography, University of Sussex, Brighton BN1 9QJ, UK b

A R T I C L E I N F O

A B S T R A C T

Article history: Available online 4 October 2009

The sedimentary sequence through the Hemingbrough Formation exposed at two sites in the central part of the Vale of York, south of the Escrick moraine ridge, is described and used to reconstruct the palaeoenvironmental history of Glacial Lake Humber. Interbedded wave ripples and laminated silts and clays at both sites indicate that Lake Humber was characterised by fluctuating water levels, often no deeper than wave base. Optically stimulated luminescence ages of 21.0  1.9, 21.9  2.0, and 24.1  2.2 kyr returned from two wave-rippled sandy beds within the glaciolacustrine sequence at Hemingbrough, c. 10 km south of the Escrick moraine ridge, provide the first direct chronological determination for the low-level phase of Lake Humber. As these beds are principally attributed to glacial meltwater emanating from the Vale of York ice lobe of the British Ice Sheet, when its margin was at or near the Escrick moraine ridge, this corroborates the interpretation that this ridge marks the LGM ice limit. ß 2009 The Geologists’ Association. Published by Elsevier Ltd. All rights reserved.

Keywords: Lake Humber Glaciolacustrine Luminescence dating Last Glacial Maximum

1. Introduction In Britain the main Late Devensian glacial episode, coincident with Marine Isotope Stage 2 (MIS 2), is termed the Dimlington Stadial, based on stratigraphic observations in eastern England (Rose, 1985). Cosmogenic nuclide surface-exposure dating and radiocarbon dating in the UK indicate that the British Ice Sheet (BIS) attained its maximum size c. 22.0 kyr (Bowen et al., 2002), during the Last Glacial Maximum (LGM) Chronozone c. 19.0– 23.0 kyr (Mix et al., 2001). Initial deglaciation had begun by 21.4  1.3 kyr, with more extensive ice wastage after 17.4  0.4 kyr. Sometime during this Late Devensian glaciation a major proglacial lake – Lake Humber – formed on the southern margin of the BIS in the Vale of York, in eastern England (Clark et al., 2004a) (Fig. 1). The location of the Dimlington Stadial/LGM limit of the BIS in the Vale of York has been subject to debate (see reviews in Evans et al., 2001; Catt, 2007). Ice emanating from eastern Pennine valleys and the Stainmore Gap flowed southward along the vale before terminating either at the Escrick moraine ridge (e.g. Kendall and Wroot, 1924; Charlesworth, 1957; Ford et al., 2008) or as much as 50 km farther south, near Doncaster (e.g. Clark et al., 2004b;

* Corresponding author. Tel.: +44 0 29 208 74830; fax: +44 0 29 208 74326. E-mail address: [email protected] (D.K. Murton). 1 Present address: Alberta Geological Survey, 4999-98 Avenue, Edmonton, Alberta T6B 2X3, Canada.

Catt, 2007; Sejrup et al., 2009). The latter limit is based on interpretation of a discontinuous NW–SE trending sand and gravel ridge between Thorne and Wroot as an ice-marginal feature marking the western edge of a transient surge of the Vale of York ice into Lake Humber (Gaunt, 1976a, 1994). However, this limit is disputed by Straw (2002), who noted the paucity of diamicton and deglacial features between Escrick and Thorne-Wroot, and the lack of ice-marginal landforms along the eastern margin of the Vale of York south of the Escrick moraine ridge. Recently, a till dated by optically stimulated luminescence (OSL) to after 23.3  1.5 kyr and before 20.5  1.2 kyr from Ferrybridge, west of Doncaster (Fig. 1), has been attributed to a separate advance of the Vale of York ice at least 15 km beyond the Escrick limit, but prior to the deposition of the Thorne-Wroot sands and gravels (Bateman et al., 2008). To clarify the Late Devensian glacial history in the central part of the Vale of York this paper examines the sedimentology of Lake Humber deposits and provides luminescence ages on them in the region south of the Escrick moraine ridge. If ice advanced southwards to this ridge and ploughed into the northern part of Lake Humber (Ford et al., 2008), then the glaciolacustrine sediments to the south of the ridge should provide a record of ice-marginal conditions. Accordingly, our objectives are: (1) to present some of the first details about Lake Humber sediments; (2) to interpret the processes that deposited them; and (3) to date some of the sediments. From these, we provide some constraints to the local glacial history. First, we summarize the existing reconstructions of the lake and the challenges of interpretation.

0016-7878/$ – see front matter ß 2009 The Geologists’ Association. Published by Elsevier Ltd. All rights reserved. doi:10.1016/j.pgeola.2009.09.001

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Fig. 1. (A) Regional map showing physiographic features and sites referred to in the text. The dashed boxes refer to the two sites shown in Fig. 2 that are the focus of this study. (B) LGM ice limits and distribution of mapped glaciolacustrine deposits in eastern England. Adapted from Ford et al. (2008). Stratigraphical context is presented in the text. Published radiocarbon ages converted into calendar years using Fairbanks 0107 (Fairbanks et al., 2005) available at http://www.radiocarbon.ldeo.columbia.edu/research/ radcarbcal.htm .

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2. Glacial Lake Humber 2.1. Summary of previous work Glacial Lake Humber (Lewis, 1894) developed in the southern and central parts of the Vale of York, south of the Escrick moraine ridge (Gaunt, 1994; Ford et al., 2008), and in the Ancholme Valley (Gaunt et al., 1992), as the BIS advancing across Holderness restricted drainage from the proto-Ouse and proto-Trent catchments through the Humber Gap (Fig. 1). Within the Vale of York, southward advance of the BIS ice caused glacio-isostatic depression towards the north. Lake Humber is thought to have existed in two distinct phases (Gaunt, 1976b): an earlier, brief, high-level phase at c. 30 m O.D., and a later, prolonged, low-level phase at c. 8 m O.D. These were separated by lake drainage and subaerial exposure of the lake floor. Drainage of the lake is thought to have occurred eastward through the Humber Gap and intermittently southeastward through the Witham Gap at Lincoln. More detailed summaries of the lake history are given by Gaunt (1981), Bateman and Buckland (2001) and Catt (2007). The high-level lake phase is inferred from discontinuous patches of rounded, but predominantly locally-derived sands and gravels at c. 30 m O.D. on the Magnesian Limestone escarpment of the eastern Pennines (Edwards, 1937). The sands and gravels are thought to define a former beach of Lake Humber (Gaunt, 1976b; Bateman et al., 2008), although Edwards (1937) attributed them to marine submergence. Recent exposures at Ferrybridge, showed a sigmoidal variation in downslope thickness of the gravels, which graded into a clast-free, moderately sorted medium sand (Bateman et al., 2008). However, in view of the nearsurface hillslope location of many of these deposits, it is likely that reworking by periglacial slope processes has removed or modified some of them (cf. Carruthers, 1947; Frederick et al., 2001). Temporary drainage of Lake Humber to 4 m O.D. or below and subaerial exposure of the lake floor have been inferred from a

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brecciated and desiccated or ground-ice cracked layer in a borehole from near Carlton, and from ventifacts and intraformational cryoturbation structures in the region (Gaunt, 1981, 1994). The low-level phase of Lake Humber is recorded unequivocally by sands and laminated silts and clays of the Hemingbrough Glaciolacustrine Formation (Ford et al., 2008)—hereafter referred to as the Hemingbrough Fm (cf. Thomas, 1999). Originally termed the 25-Foot Drift (Edwards, 1937), the Hemingbrough Fm is up to 24 m thick and rests directly on bedrock or sands and gravels of the Basal Glaciofluvial Deposits (Ford et al., 2008). It comprises from the base upwards: the Park Farm Clay Member, the Lawns House Farm Sand Member and the Thorganby Clay Member. Gaunt (1994) inferred from the persistent and even lamination of the Hemingbrough Fm that it was deposited in quiescent depositional conditions in standing water. He suggested a lake level not much higher than the upper limit of the formation at 8 m O.D. The age relationship between the Hemingbrough Fm and the sands and gravels at c. 30 m O.D. remains difficult to establish in the absence of radiometric dates from both deposits. Gaunt (1976b) attributed them both to deposition in Lake Humber, with the Hemingbrough Fm deposited mostly during the low-level phase, after the Late Devensian maximum ice advance in the Vale of York and Holderness; the lowest part of the clay and lacustrine sands beneath it, however, may be coeval with the sands and gravels of the initial high-level phase (Gaunt, 1994). However, Straw (1979, 1991) suggested that the high-level phase may be of Early Devensian age. Central to Gaunt’s two-stage model of Lake Humber is blockage of the Humber Gap, initially by an ice dam, and subsequently by morainic debris (Gaunt et al., 1992). However, evidence for two distinct lake levels is not always found. Only the low-level phase is inferred from stratigraphic observations at South Ferriby cliff and the Horkstow Moraine at Eastfield Farm, in the Humber Gap (Frederick et al., 2001). Other proglacial lakes developed as the as the Vale of York ice lobe retreated northward from the Escrick and York moraine ridges

Fig. 2. Location maps for the sites at Hemingbrough and Glade Farm. Scale and orientation are given by National Grid co-ordinates. Contours show elevations above Ordnance Datum (m O.D.).

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(Ford et al., 2008). Lake Elvington formed between them, and is recorded by 3–6 m of laminated silt and clay, with numerous sand beds, assigned to the Elvington Glaciolacustrine Fm. Continued ice retreat impounded Lake Alne to the north of the York moraine ridge; within it accumulated c. 2.5 m of laminated silt and clay, with occasional sand beds, of the Alne Glaciolacustrine Fm. The surface elevations of the two formations are 10–12 m and 14–17 m O.D., respectively. It is uncertain whether Lake Alne co-existed with, or developed after, Lake Humber disappeared. The demise of Lake Humber occurred either by gradual silting up of the lake (Gaunt, 1981, 1994) or drainage by breaching of its morainic plug through the Humber Gap (Frederick et al., 2001). Its demise was followed by fluvial deposition of sand leve´es on the emergent lake bed south of the Escrick moraine ridge, with subsequent fluvial incision initiating the contemporary drainage network (Gaunt et al., 1992; Gaunt, 1994). The age of Lake Humber has proved difficult to pinpoint without directly dating the lacustrine sediments of the Hemingbrough Fm. Although the lake is thought to have existed sometime between c. 23 and 14 kyr BP (reviewed in Bateman et al., 2008), both LGM and younger Devensian ages have been proposed (Fig. 1B). A maximum age of 26,163  2046 cal yr BP has been inferred from a bone fragment reworked into or below the base of high-level Lake Humber deposits on the eastern margin, at Brantingham (Gaunt, 1974), while sand attributed to the high-level lake phase itself at Caistor has provided a TL age of 22.7  1.4 kyr (Bateman et al., 2000). Both ages are broadly consistent with radiocarbon ages of 22,043  497 and 21,734  372 cal yr BP obtained from moss within the Dimlington Silts, which underlie the Skipsea Till at Dimlington (Penny et al., 1969); ice from the North Sea Basin blocked the Humber Gap and deposited this till. In contrast, comparable littoral deposits from the western margin of Lake Humber, at Ferrybridge, returned an OSL age of 16.6  1.2 kyr (Bateman et al., 2008), suggesting that the high-level phase was significantly later than the LGM, occurring during the Killard Point Stadial of Ireland (McCabe and Clark, 1998) and Greenland Stadial 2a (Lowe et al., 2008). Minimum ages from peat at or near the base of blown sand overlying the glaciolacustrine sequence of 12,965  180 cal yr BP from Armthorpe (Gaunt et al., 1971), 12,879  168 cal yr BP from near York (Matthews, 1970) and 12,398  109 cal yr BP from Cawood (Jones and Gaunt, 1976) coincide with, or slightly precede, Greenland Stadial 1 (the Younger Dryas). 2.2. Field sites Two field sites were selected for study on the basis of the largest vertical exposures of Lake Humber sediments currently available south of the Escrick moraine ridge. The sites were recently active clay pits located at Hemingbrough [National Grid reference SE 674317], and adjacent to Glade Farm [SE 622402] (Figs. 1A and 2), c. 10 and 1 km south of the ridge, respectively. Although this region was visited as part of Xth INQUA Congress in 1977 (Catt, 1977), hitherto only generalised accounts of the sediments exposed at Glade Farm (Carruthers, 1947) and Hemingbrough (Catt, 1977; Gaunt, 1976b, 1994) have been published. The type site for the Hemingbrough Fm (Thomas, 1999) is from an earlier working [SE 675316] at the current field site. 3. Methods 3.1. Sedimentary analysis Open-face exposures of the sedimentary sequences at Hemingbrough and Glade Farm were logged between July 2004 and August 2007, and continuous monoliths collected to aid identification of small-scale sedimentary structures. In the laboratory, the monoliths were cleaned by scalpel, photographed and logged.

Sedimentary units were identified on the basis of composition, sedimentary structures, colour and particle size. Moist colour was defined using a Munsell Color soil chart. Sediments for particle-size analysis, determined using laser diffraction on a Malvern Mastersizer 2000, were collected from strata > 10 mm thick to avoid cross-bed contamination. Clast lithologies were identified from washed samples, although there were insufficient numbers for rigorous lithological analysis. 3.2. Optical stimulated luminescence (OSL) dating Samples for OSL dating were collected by hammering opaque plastic tubes into the section face. Three samples were collected from Hemingbrough, and one from Glade Farm. In the laboratory sand-sized quartz was extracted under subdued orange light conditions using standard preparation procedures. Carbonates and organic matter were firstly removed respectively using HCl and H2O2 washes, and then the samples were wet sieved to isolate grains in either the 90–125 or 125–180 mm size fractions. Heavy minerals were separated in sodium polytungstate (2.89 gm/cm3) and the remaining sediment was subsequently agitated in 40% HF solution for 60 min to etch the alpha radiation contaminated outer rind of the quartz. Finally, the quartz-rich fractions were agitated in 40% H2SiF6 solution for 5 days in order to dissolve any residual feldspar grains followed by shaking in concentrated HCl for 1 h to remove fluoride precipitates. Environmental dose rates were determined from radioisotope concentrations as measured in the laboratory from sediment taken from the ends of the sampling tubes using inductively coupled plasma mass spectrometry/atomic emission spectrometry (ICPMS/AES). Field measurements of the gamma radiation dose were also made using an Ortec MicroNomad NaI detector. The conversion of radioisotope concentrations into dose rates used the factors of Adamiec and Aitken (1998) with the absorption of beta particles with respect to sediment grain size accounted for using the factors of Fain et al. (1999). A cosmic dose rate contribution was also calculated from the altitude, latitude, and longitude of the section as well as the thickness and estimated density of the overburden (Prescott and Hutton, 1994). Dose rates were attenuated for sediment water content assuming present-day values as measured in the laboratory. Luminescence measurements were performed on a Risø OSL/ TL-DA-15 system using blue-light LED stimulation (470 nm, c. 45 mW/cm2) with the ultraviolet emissions detected by a 9235QA photomultiplier tube fitted with a 8 mm-thick Hoya U-340 optical filter. Laboratory irradiation used the 90Sr/90Y b-source that is mounted within the Risø system and delivering a dose rate of 8.5 Gy/min. Quartz grains were mounted onto 9.8 mm diameter steels with the inner 9 mm part of the disc covered with a monolayer of c. 1000 grains using silkospray. OSL measurements were performed for 40 s whilst the samples were held at 130 8C with the first 0.4 s of the OSL decay curve being integrated and a background subtracted from the last 4 s of the stimulation. The purity of each aliquot with respect to feldspar contamination was checked using infra-red stimulation (performed at 130 8C) which preferentially depletes the feldspar signal over quartz (e.g. Bannerjee et al., 2001). Despite HF and H2SiF6 etching, a small residual signal (752 counts/0.4 s) resulting from IR stimulation was present in sample HEMA02, comprising 21% of the regenerated blue-light stimulated signals (3555 counts/0.4 s). As a precautionary measure, all OSL measurements were preceded by an IR bleach to reduce any residual feldspar signals to a low level. Equivalent dose (De) was estimated using the SAR (single aliquot regenerative dose) protocol (Murray and Roberts, 1998; Murray and Wintle, 2000, 2003). The OSL resulting from a small test dose (Tx) applied after each natural and/or regenerative

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Fig. 3. Stratigraphic logs for Hemingbrough (A) and Glade Farm (B) showing the dominant sedimentary structures. OSL sample HEMA03 was sampled from an exposure of H1, 20 m west of the logged section.

measurement (Lx) was used to normalise any luminescence sensitivity changes. Four to five regenerative doses were used to produce the sensitivity-corrected growth curve (Dx = Lx/Tx, x = 1, 2, 3) with the equivalent dose calculated by interpolation of the natural luminescence signal (N = Ln/Tn). Recuperation of the OSL signal was monitored by the OSL resulting from a zero gray dose point (D0 = 0 Gy) and the success of the sensitivity correction was checked by using a repeated dose point (D1) to obtain the recycling ratio. Aliquots were rejected from the analysis if levels of recuperation were greater than 5% of the natural signal (D0/ LnTn > 5%) or if the recycling ratio exceeded 1.0  0.1. The De of at least 12 aliquots was determined for each sample with uncertainties on each estimate being calculated using the Monte Carlo method. These replicate measurements were combined into a single De estimate for each sample using an uncertainty-weighted mean (central age model, Galbraith et al., 1999). 4. Stratigraphy and sedimentology 4.1. Description 4.1.1. Hemingbrough At Hemingbrough clay pit, the upper 10.1 m of the Hemingbrough Fm was examined in several excavated vertical faces. The sequence observed comprises three sedimentary units, HB1– HB3 (Fig. 3A).

Unit HB1 (5.6 m thick) consists of laminated clayey silt interbedded with massive to faintly stratified clayey silt. The laminated clayey silt – as observed in the monoliths – is characterised by wavy parallel, planar parallel or lenticular lamination (Fig. 4A). The laminae are <1 mm to several mm thick, and commonly alternate between light-coloured silt or fine sand and dark-coloured clayey silt, together forming distinctive couplets. Some light-coloured laminae have sharp lower contacts and grade smoothly upwards from silt into dark clayey silt. In other cases, the dark clayey silt laminae sharply overlie – and are often more laterally continuous than – the light-coloured silty laminae (Fig. 4A). The massive to poorly stratified clayey silt beds are dark greyish brown (10YR 3/2) (Fig. 4B), with a median particle size of 6.3 mm (n = 66). They are 30–210 mm thick, laterally continuous, sometimes graded and irregularly spaced between the laminated parts of unit HB1. Inclined or recumbent folds, or convolute lamination are common in the unit, and often occur near laminae that are discontinuous or irregular and have diffuse contacts, or where laminated parts merge into massive or faintly stratified parts (Fig. 4B). Intraformational faults and occasional flint pebbles are also present in unit HB1. The unit grades upwards into unit HB2. Unit HB2 (3.3 m thick) comprises massive silt that contains distinctive sandy interbeds or laminae. The silt is dark greyish brown (10YR 4/2) or brown (10YR 4/3), and slightly sandy, with a median particle size of 24.5 mm (n = 34); it forms beds 20–210 mm

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Fig. 5. Sandy beds within massive silt, unit HB2. (A) Convolute lamination forms the central part of bed H4. Wavy and flaser bedding (wave ripples) in bed H5 are overlain by silty drapes. Unit HB3 caps the sequence. (B) Peaked symmetrical ripple form sets on the surface of bed H1 exposed on a horizontal surface represent wave ripples that are characteristic of very shallow, near-emergent conditions in Lake Humber.

Fig. 4. Monoliths of laminated clayey silt interbedded with massive to faintly stratified clayey silt, unit HB1. (A) Laminated clayey silt, showing wavy parallel, planar parallel and lenticular lamination. The laminae typically form couplets of light-coloured silt or fine sand and dark-coloured clayey silt. (B) Inclined fold within faintly stratified to massive clayey silt. Scale bar in centimetres. Images digitally manipulated to clarify the sedimentary structures.

thick that are generally massive and brecciated. Some silty beds, however, contain remnants of faint horizontal parallel lamination, lenticular lamination or irregular and diffuse lamination; and some beds contain silty intraclasts that are massive or laminated. The interbeds or laminae within unit HB2 comprise silty sand to sandy silt. They are very dark greyish brown (10YR 3/2) and show a gradation of structures between (a) isolated ripple forms, (b) continuous horizontal laminae 1–10 mm thick and (c) beds 10– 180 mm thick. The ripple forms and laminae are common throughout the unit, whereas the beds tend to become thicker towards the top of it; the most prominent beds are numbered H1– H5 (Fig. 3A). The particle size of the beds tends to increase upward through the unit, with median values of 34.5 mm and 84.1 mm for H1 and H5, respectively. Internally, the sandy beds are massive to laminated, the laminae including wavy to undulating horizontalsubhorizontal forms 1 mm to a few mm thick. Wavy and flaser lamination occur within H5, with silt drapes above symmetrical, round-crested to peaked ripple form sets (Fig. 5A). The internal lamination is often poorly defined, and convolute lamination is common (H4 in Fig. 5A). The lower contacts of the sandy beds are

sharp and more or less horizontal, sometimes with silt intraclasts above them. The upper contacts commonly have well developed ripple form sets with symmetrical, straight to slightly sinuous crests that have peaked or rounded forms. Ripple crests oriented at 084–2648 were observed in a horizontal exposure of H1 (Fig. 5B). Similar ripple form sets locally occur as lenticular lamination within some of the host silt. Ball-and-pillow structures of sand are common throughout the unit. Unit HB2 is sharply overlain by unit HB3. Unit HB3 (0.7–1.2 m thick) consists of laminated to massive clayey silt. The silt is very dark greyish brown (10YR 3/2). When exposed in dried and weathered quarry faces, the silt sometimes shows a wavy parallel lamination, with laminae about one to several mm thick; such faces are finely brecciated. By contrast, moist and fresh exposures reveal no evidence of sedimentary structures within the silt. The monoliths show massive to faintly laminated mottled silt to silty sand. 4.1.2. Glade Farm At Glade Farm, the sedimentary sequence examined in vertical excavated faces is up to 8.8 m thick, and comprises four units, GF1– GF4 (Fig. 3B). At the base of the sequence, unit GF1 (2.6 m thick) consists of laminated clayey silt, with minor amounts of sand. The stratification types observed in the field are planar parallel and lenticular lamination, with subsidiary wavy lamination and bedding. The lenticular and wavy types are characterised by round-crested to peaked symmetrical ripple form sets of silt or sandy silt interspersed within clayey silt. Prominent wavy bedding in the form of three 10–25 mm-thick sandy silt beds draped by clayey silt laminae occurs between depths of 7.8 and 8.0 m (Fig. 3B); beneath some symmetrical ripple form sets at the top of

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the beds, the cross lamination is bidirectional (Fig. 6). A 20 mmthick silty sand layer occurs at depth of 8.6 m. Overall, the laminated clayey silt is similar in colour, particle size and lamination types to unit HB1. The monoliths from unit GF1 reveal details of the lamination. The laminae typically occur as distinct couplets of (1) yellowish brown (10YR 5/4) to brown (7.5YR 5/3) silty or silty sand that are overlain by (2) very dark greyish brown (10YR 3/2) to dark brown (7.5YR 3/2) clayey silt (Fig. 7A). The lighter-coloured silty laminae comprise continuous, discontinuous and diffuse types. Continuous silty laminae are usually <1–2 mm thick and have upper and lower contacts that are sharp and planar to undulating; some laminae, however, are graded, both within the light-coloured silt and extending into the overlying dark clay. Discontinuous silt laminae (i.e. lenticular lamination) vary in width from a few millimetres to a few centimetres and are sharply overlain by a clay drape (Fig. 7A).

Diffuse silty laminae are irregular, fade out laterally and are often associated with intraformational folds (Fig. 7B). In contrast, the darker-coloured clayey laminae are generally continuous and vary in thickness from <1 mm to several millimetres. Some silty or clayey laminae are disrupted by diapirs, flame structures and load casts, and intraformational normal faults occur throughout unit GF1. The unit grades upward into unit GF2. Unit GF2 (4.1 m thick) comprises generally massive clayey silt. The unit is very dark greyish brown (10YR 3/2), with a median particle size of 3.7 mm. Some lenticular and wavy lamination occur in the upper c. 0.5 m of unit GF2, similar to that in unit GF1. Curved and polished fracture surfaces (slickensides) – some marked with grooves and ridges – are abundant throughout the unit; of 10 slickensides measured, the average angle of dip was c. 288 and the direction of dip towards 1908. The vertical spacing between slickensides varies from a few millimetres to several centimetres. Unit GF3 (0.8–1.2 m thick) comprises clayey silt with sandy interbeds or laminae. The clayey silt is dark to very dark greyish brown (10YR 4/2 to 3/2) and forms beds 5–150 mm thick that are faintly laminated and commonly brecciated. The sand to silty sand is dark yellowish brown (10YR 4/4), massive to laminated, and occurs as individual ripple forms, laminae 2–10 mm thick or beds 10–110 mm thick (Fig. 8A and B), similar to those in unit HB2. Lamination types include (a) cross lamination; (b) subhorizontal, slightly wavy parallel lamination; and (c) wavy lamination; and they are often defined by alternations of silty sand and clayey silt. Some of the lamination, however, is irregular and discontinuous, with signs of incipient convolution lamination. Intraclasts of clayey silt occur along the base of some sandy beds (Fig. 8A). Upper and lower contacts are sharp planar, irregular or undulating. Ripple form sets on top of the sandy beds are symmetrical and roundcrested to peaked. Wavy and lenticular lamination are common throughout unit GF3, and some flaser bedding and wavy bedding occur near the top of it; ripple form sets comprise both round-

Fig. 7. Laminated clayey silt, unit GF1. (A) Parallel lamination comprising lightcoloured (siltier) and dark-coloured (clayier) couplets. Some silty laminae are graded, whereas others have sharp upper contacts. Lenticular lamination and normal faults are also present. (B) Diffuse, irregular or folded silty laminae with a sequence of parallel lamination and lenticular lamination. Scale bar in centimetres. Images digitally manipulated to clarify the sedimentary structures. The dark band two thirds of the way down (A), and at the top of (B) are cracks that developed during air drying.

Fig. 8. Clayey silt with sandy beds, unit GF3. The sandy beds are massive to laminated, some with intraclasts of clayey silt above their sharp lower contacts (lower sand bed in A). The clayey silt is massive to faintly laminated (B). Scale bar in centimetres.

Fig. 6. Wavy bedding from one of three prominent sandy beds, 7.8–8.0 m depth, unit GF1. Symmetrical ripple form sets with peaked to rounded crests (wave ripples) are draped by clayey silt. Bidirectional cross lamination occurs beneath some ripple forms. Above the wavy bedding is parallel lamination and lenticular lamination.

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4.2. Interpretation The depositional processes and environment associated with the sediments at Hemingbrough and Glade Farm are interpreted together because the sites are located only 10.5 km apart in the same drained-lake basin and exhibit similar sedimentary characteristics. 4.2.1. Laminated clayey silts The laminated clayey silts in units HB1 and GF1 are characteristic of fine-grained glaciolacustrine deposits (cf. Gilbert, 1975; Ashley, 1988). The light- and dark-coloured couplets (Fig. 7A) are attributed to rhythmic deposition in a proglacial lake (Smith and Ashley, 1985). Couplets with a sharp contact between the silty lamina and the overlying clayey lamina are interpreted as annual rhythmites (varves): the silty lamina records inwashing of coarser sediment – probably as overflows or interflows issuing into the lake from the mouths of rivers or glacial streams – during spring and summer, whereas the clayey lamina develops from suspension settling in quiet-water conditions beneath lake ice during winter. In contrast, couplets that smoothly grade upward from a silty lamina to a clayey lamina probably indicate episodic turbidity currents: as the current velocity diminishes, so does the flow competence, allowing progressively finer particles to settle from suspension (Kuenen and Menard, 1952). Ripple form sets within the laminated clayey silts record wave action in shallow water of the lake. Significantly, no clear examples of climbing-ripple cross lamination or other forms of current ripples were identified in the field sections or monoliths. Instead, wave ripples are indicated by the peaked symmetrical ripple forms – some with bidirectional cross lamination (ripple on left side of Fig. 6) resulting from wave orbital motion (Reineck and Singh, 1975, Figs. 16 and 20) – and by the wavy and flaser bedding/ lamination (Fig. 3B). Peaked symmetrical ripples characterise very shallow, near-emergent conditions, while round-crested ones may characterise deeper water (Collinson et al., 2006) or they may result from reworking of wave ripple crests during the process of emergence (Reineck and Singh, 1975). Fig. 9. (A) Sand of unit GF4, between 0.7 and 1.6 m marks on the staff. A discontinuous pebbly sand-sandy gravel layer – between 1.1 and 1.2 m – truncates stratification in the underlying sand. Location of 3 monoliths is shown to the right of the staff. The lower 30 cm of the unit – between 0.7 and 1.0 m – comprise interbedded sand and clayey silt. (B) Wavy bedding and flaser bedding associated with rounded symmetrical ripple forms (wave ripples) in the lower part of the unit GF4. Trowel for scale.

crested and peaked forms. The lower contact of the unit is sharp and undulating or, where it cuts down into unit GF2, concave-up, forming a channel structure at a depth of 2.1 m. Overall, unit GF3 becomes increasing sandy towards the top, grading upward into unit GF4 (Fig. 3B). Unit GF4 (0.9 m thick) comprises medium to fine-grained sand (Fig. 9A). The lower c. 30 cm of the unit consists of interbedded sand and clayey silt, the stratification of which occurs as wavy bedding and flaser bedding associated with symmetrical rounded ripple forms (Fig. 9B). Above the basal 30 cm, the sand becomes faintly laminated to massive. Convolute lamination and irregular silty streaks occur in some parts of this unit, and occasional pebbles are scattered throughout it. A prominent but discontinuous subunit of pebbly sand to sandy gravel up to 10 cm thick occurs in the middle of the unit (Fig. 9A). The pebbly sub-unit is massive and has a sharp lower and upper contact. It contains rounded to subangular clasts, up to a few centimetres in maximum dimension, of Carboniferous Sandstone and chert, with occasional Magnesian Limestone pebbles.

4.2.2. Massive to faintly stratified silt and clayey silt The massive to faintly stratified silt and clayey silt in units HB1– HB3 and GF2–GF3 is attributed in part to resedimentation of laminated clayey silt. Resedimentation is indicated by remnants of parallel lamination similar to undisturbed lamination in the laminated clayey silt of units GF1 and HB1. In addition, silty intraclasts – some of them laminated – within unit HB2 indicate erosion from pre-existing laminated sediments. Resedimentation of lake-bottom muds may have taken place by subaqueous slumping and/or mudflows. In addition, some massive beds probably formed by steady deposition from suspension of abundant silt and clay mixed throughout the water column. Similar massive beds from distal sediment cores in water depths of >5 m and <15 m in proglacial Sunwapta Lake, Alberta, Canada, have been attributed to an absence of strong diurnal cycles of meltwater inflow and wind (Gilbert and Shaw, 1981). Lastly, some massive beds may record wave disturbance or mixing of sediment by wind-generated currents in shallow lake waters (Smith and Ashley, 1985). 4.2.3. Sandy beds The sandy beds in units HB2 and GF3 are attributed mainly to wave action. Although turbidity currents can deposit sandy beds in glacial lakes (Gilbert and Shaw, 1981), the apparent absence of normal grading and ripple-drift cross lamination indicative of suspension settling and traction (Jopling and Walker, 1968) argues against this process. Alternatively, concentrated or hyperconcen-

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trated density flows (Mulder and Alexander, 2001) generated by slumping of soft sediment on the front of the Escrick moraine ridge might account for the massive nature of some sandy beds (cf. ‘grain flow’ sands in Shaw, 1977), but such flows would have had to have crossed the almost flat bottom of the lake to reach the basin centre, near Hemingbrough. Although this process cannot be discounted, a more likely process is wave-induced sediment movement. Wave action is indicated by the occurrence of wave ripples on top of and within some sandy beds. As discussed above, the peaked symmetrical ripple forms suggest shallow, near-emergent conditions. Analogous conditions are those within sandy tundra lakes on the Arctic Coastal Plain of Canada and Alaska. For example, waverippled sands are abundant on beaches and in the shallow-water zone (up to c. 1–2 m deep) above wave base in the Tuktoyaktuk Coastlands of western Arctic Canada, where waves rework the sandy lake floor during windy spells in summer; the wave-rippled sands, sometimes with fairweather silty drapes above ripple forms, have a sheet-like geometry (Murton, 1996, 2001). Shallow-water conditions are also consistent with the presence of convolute lamination in units HB2 and GF3; liquefaction just beneath the sediment surface is often triggered by the movement of waves over the water surface above (Allen, 1985) or by a rising and falling water table (Collinson et al., 2006) associated with lake-level changes.

Lake Humber sediments occurred during emergence of the lake shallows—which would favour permafrost aggradation (cf. Mackay, 1997). It would also have occurred during freezing of lake ice to the lake bottom in shallow water whose depth was less than the maximum thickness of winter ice (usually c. 2 m)—which would permit freezing of saturated sediments beneath the ice (cf. Burn, 1990) and favour preservation of permafrost under shallow water, as is common beneath arctic lakes (Shilts and Dean, 1975). The fine brecciation observed in dried faces through unit HB3 may record either desiccation and/or freezing-induced consolidation. Shearing took place within the massive clayey silt of unit GF2. The slickensides in this unit indicate that the sediment had significant cohesion when shearing occurred. Unless the sediment was quite deeply buried and consolidated prior to erosion of the overburden, a more likely cause of consolidation was freezing of this frostsusceptible clayey silt (Williams and Smith, 1989), as discussed above. The cause of the shearing, however, remains to be determined. Possibilities include: (1) transmission of compressive stress from a glacier into frozen proglacial sediments (cf. Etzelmu¨ller and Hagen, 2005), if the Vale of York ice lobe pushed south into partially frozen (warm) permafrost in the clayey silt, when the ice margin was at or near the Escrick moraine ridge; (2) shrink-swell behaviour of clay that accompany drying-wetting cycles, as is common in vertisols; or (3) mass movement of some type.

4.2.4. Sand unit GF4 The sand unit GF4 represents either a fluvio-aeolian deposit or a shallow-water lake deposit, or a combination of them. On stratigraphic grounds, unit GF4 could be interpreted as part of a fluvio-aeolian cover sand sequence (Skipwith Sand Member of the Breighton Sand Fm; Ford et al., 2008). After lake drainage, the top of the glaciolacustrine deposits may have been reworked by fluvial and aeolian processes, with clays and silts being interbedded with sands and passing up into fluvial sands with some gravel lags and evidence for downcutting, before finally passing up into a mixed aeolian and fluvial sequence. This interpretation assumes that the Lawns House Farm Sand Member and the overlying Thorganby Clay Member have been eroded away at the Glade Farm pit, probably when or just before the Breighton Fm sands were deposited. This fluvio-aeolian interpretation, however, is inconsistent with the occurrence of wave ripples – indicated by wavy bedding and flaser bedding – in the basal 30 cm of the unit (Fig. 9B). Furthermore, we have not observed any evidence of peat within or beneath unit GF4, in contrast to descriptions of the Skipwith Sand Member reported in Ford et al. (2008, plate 4). Instead, the basal 30 cm are clearly shallow-water lake deposits. Above this, the limited sedimentary structures observed are inconclusive about depositional conditions.

5. OSL dating results

4.2.5. Deformation structures Soft-sediment deformation frequently occurred within sediments beneath the lake floor, producing convolute lamination, ball-and-pillow structures, diapirs, flame structures and load casts (Allen, 1985; Collinson et al., 2006). Triggers of deformation probably included wave action and subaqueous slumping. Smallscale slumping is inferred from inclined and recumbent folds, and from irregular, diffuse lamination. Slumping may have been caused by rapid loading of sand onto fine-grained sediments (cf. Gilbert and Shaw, 1981) or by lake-floor instability. Freezing of sediments beneath the emergent floor or shallows of the lake occurred intermittently. The brecciation within freshly exposed faces through silty deposits of units HB2 and GF3 resembles brecciation associated with reticulate cryostructures in clays and silts within permafrost or seasonally frozen ground (Murton and French, 1994). By analogy with modern Arctic conditions, winter freezing of the upper decimetres or metres of

5.1. Luminescence characteristics Four samples for OSL dating were collected from the study sites, with a single sample being taken from Glade Farm in unit GF4 (Fig. 3B) and three samples taken from Hemingbrough in unit HB2 (Fig. 3A). Representative examples of OSL decay curves for these samples are illustrated in Fig. 10A and show the presence of a rapidly decaying signal typical of the quartz fast component which is most commonly used for dating (Wintle and Murray, 2006), but the emissions being of a relatively weak intensity. In order to confirm that the initial part of the OSL signal used in dating is dominated by the quartz fast component, individual decay curves for each sample were summed to increase the signal-to-noise ratio. By curve fitting (using the Levenberg–Marquardt algorithm), the data could be resolved into three exponential decays termed the ‘fast’, ‘medium’, and ‘slow’ components (Fig. 10B) which represent emissions from traps with different optical decay rates (e.g. Smith and Rhodes, 1994; Bailey et al., 1997). The mean photo-ionisation cross-section of the most rapidly decaying component was calculated to be 2.54  0.08  10 17 cm 2 which closely matches values for the quartz fast component (Jain et al., 2003). This component dominates the initial 0.4 s integral of the OSL signal in these samples, forming 88  6% of total luminescence after background subtraction. Prior to each OSL measurement, pre-heating was necessary to remove charge from thermally unstable traps (Aitken, 1998). The dependency of the SAR procedure on pre-heat temperature was assessed using a dose recovery pre-heat plateau test with groups of three aliquots being bleached twice with blue light for 500 s separated by a 10 ks pause. A laboratory dose approximately equivalent to the sample’s natural dose was subsequently administered to the aliquots and De was measured using the SAR protocol at different pre-heat temperatures. Pre-heat and cut-heat temperatures were all >220 8C because LM-OSL (linear modulation) revealed the presence of the thermally unstable ‘ultra-fast’ component which can be completely removed by a cut-heat of 220 8C (Choi et al., 2003) (Fig. 11A). Good dose recovery, recycling ratios, and low levels of recuperation were found to occur for pre-heat temperatures of between 220 and 280 8C (Fig. 11B) and a pre-heat temperature of 240 8C was used in subsequent measurements. Under these condi-

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Fig. 10. Characteristics of the OSL signal in the study samples. (A) Typical OSL decay curves for sample HEMA02 resulting from natural and regenerative dose measurements, showing the rapidly decaying ‘fast’ component signal, recorded over 40 s using blue-light stimulation with the sample held at 130 8C. (B) Component analysis of sum of 13 individual OSL decay curves (sample GLAD01) resulting from a regenerative dose of 322 Gy. Photo-ionisation cross-sections (PICs) which are proportion to the decay rates of each component are also shown.

tions, the mean measured/given dose ratio for all four samples is 1.05  0.03, demonstrating that the SAR protocol can recover a known laboratory dose in these samples to within 5% of unity. 5.2. Equivalent doses and OSL ages A summary of the De estimates is shown in Table 1 alongside the stratigraphic position of the samples and environmental dose rate data. De estimates obtained for the quartz fast component by curve fitting are also shown because this signal is most rapidly bleached in sunlight. In the three samples taken from Hemingbrough, precise De estimates were obtained with the growth of the sensitivity-corrected OSL signal showing no evidence of saturation (Fig. 12), resulting in sample ages of 21.0  1.9, 21.9  2.0, and 24.1  2.2 kyr (weighted mean = 22.2  0.5 kyr)

Fig. 11. (A) LM-OSL (linear modulation) signal for sample HEMA03 with the stimulation intensity ramped from 0 to 90% power over 80 s, resulting in a highresolution measurement of the rapidly decaying OSL components. The thermally unstable ultra-fast component which is undesirable for dating is present at low cutheat temperatures but is removed by heating to 220 8C. (B) Pre-heat plateau dose recovery test with successfully recovery of a known laboratory dose of 49 Gy, good recycling ratios, and low levels of recuperation observed for pre-heat temperatures of 220–280 8C.

when using the fast component. The De of the sample taken from Glade Farm, however, was close to the point of saturation and resulted in a considerably older age of 105.3  10.2 kyr. The samples that were taken from Hemingbrough were collected from waverippled sand beds which require shallow water depths (above wave base) which would result in a high propensity for adequate solar resetting of the OSL signal. In contrast, the depositional environment relating to the deposition of the GF4 sand unit at Glade Farm is not certain. If partial bleaching is considered a possible cause of age

Table 1 Summary of sample details, dosimetry, equivalent doses, and sample ages using the fast component estimates of De. Dose rates are based on the conversion of the radioisotope concentrations, gamma dose rate from NaI spectrometry, cosmic dose rate calculated from the depth of the samples and an internal dose to quartz of 0.06  0.02 Gy/ka. Sample details Sample Hemingbrough HEMA01 HEMA02 HEMA03 Glade Farm GLAD01

Dosimetry

Equivalent dose

Depth (m)

K (%)

ICP-MS/AES U (ppm)

Th (ppm)

wc (%)

Dose rate (Gy/ka)

De (0.4 s)

De (fast)

90–125 90–125 90–125

1.5 4.0 4.0

1.8 1.6 1.5

1.9 2.4 2.0

6.2 8.0 7.0

20 23 23

2.47  0.22 2.49  0.23 2.17  0.20

54  1 52  1 55  1

54  1 52  1 52  1

125–180

0.7

1.6

2.1

6.9

23

2.54  0.24

261  27

Grain size (mm)

267  5

Age (kyr) 21.9  2.0 21.0  1.9 24.1  2.2 105.3  10.2

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the ages at Hemingbrough appears to indicate that these samples are unlikely to be significantly affected by partial bleaching. 6. Discussion 6.1. Shallow-water conditions

Fig. 12. Typical SAR growth curve for sample HEMA03 showing no evidence for signal saturation and resulting in precise A estimates.

overestimation, single-grain measurements are now commonly performed in OSL dating, allowing more completely reset grains to be identified (e.g. Olley et al., 2004). However, further investigation at a single-grain level is not considered practical for the samples in this study owing to the low intensity of the OSL signals even when using large aliquots containing >1000 grains. Although being less conclusive, investigation of De as a function of stimulation time, termed De(t) plots have also been used to detect evidence for partial bleaching by the presence of a rising De trend as the integration width of the OSL signal is widened, resulting in a greater contribution from the more difficult to bleach medium component (Bailey, 2003). The samples taken from Hemingbrough showed the presence of a De(t) plateau over the first 2.4 s of the stimulation where the fast and medium components dominate (Fig. 13). The interpretation of the De(t) plot at Glade Farm, however, was complicated by a strongly falling (>34 Gy) De(t) trend reflecting a substantial contribution from a thermally unstable slow component (Jain et al., 2003). Nevertheless, the absence of rising De(t) trends and consistency of

The latter stage of Lake Humber was characterised by shallow, fluctuating water levels. Shallow water is indicated by the abundant wave ripples within the Park Farm Clay Member (units HB1, HB2, GF1 and GF2) and unit GF4. Fluctuating lake levels are inferred from the alternation of sandy wave ripples and clayey silts in this member, recording alternating shallow water and deeper and/or quieter water conditions, respectively. The location of Hemingbrough in the central part of the lake basin indicates that shallow-water conditions were not confined to the basin margin, but were basin-wide. Farther southeast, at Eastfield Farm near Winteringham, close to the Humber Gap (Fig. 1A), similar ripples – with straight crests, bidirectional foresets and clay drapes – from laminated sands above Skipsea Till have also been attributed to deposition at the edge of a standing body of water that was possibly the low-level phase of Lake Humber (Frederick et al., 2001); if so, then Lake Humber in the vicinity of the Humber Gap was also shallow after ice retreat. The widespread occurrence of shallow, fluctuating lake levels may explain the lack of welldefined strandline deposits associated with the low-level phase of Lake Humber reported by Edwards et al. (1940). 6.2. Sand sources A major influx of sand into Lake Humber occurred from the margin of the Vale of York ice lobe at the Escrick moraine ridge. Although the Park Farm Clay Member underlies the Escrick and York moraine ridges (Ford et al., 2008), and continues northward into the Harrogate district (Cooper and Burgess, 1993), the overlying Lawns House Farm Sand Member forms a lobe extending south from the Escrick ridge for up to 15 km, thinning gradually

Fig. 13. De(t) plots for the study samples, showing the presence of flat or slightly falling trends. No evidence of partial bleaching is present, indicated by rising De(t) values.

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(Ford et al., 2008). About 1 km south of the ridge, at Glade Farm, the substantial thickness of sand (0.9 m) and the coarse pebbly subunit within unit GF4 are consistent with deposition close to a substantial source of relatively coarse-grained sediment. Wave ripples within the basal 30 cm of unit GF4 suggest shallow water close to the ice margin. At both Glade Farm and Hemingbrough the sandy beds progressively increase in frequency upwards through the stratigraphy, suggesting an increasing sand supply from a nearby source. Even though the sequence at Hemingbrough contains no single, thick sand layer directly comparable to the Lawns House Sand Member, the prominent sandy beds (H1–5 in unit HB2) are regarded as distal equivalents. The sand was probably introduced to the lake as a combination of outwash and slumped material. Significantly, the thick sandy unit GF4 at Glade Farm lies directly south of the Crockey Hill meltwater system, which is defined by sandy and gravelly deposits of the Crockey Hill Esker Member running north–south between the York and Escrick moraine ridges (Ford et al., 2008). Erosion of the lake bed by flowing water is indicated by the sharp and locally channelled nature of the lower contact of the sandy beds observed in the present study and by the incorporation of rip-up clasts of underlying silt. Slumping of sediment on the Escrick moraine ridge would have been likely during ice-marginal conditions, triggered perhaps by rapid deposition of soft unconsolidated sediment and fluctuating lake levels, which transferred large amounts of sediment into the lake (cf. Zolitschka, 1996). Additional sand influxes to the central part of Lake Humber supplied by inflow from Pennine drainage down the Wharfe glacial diversion or River Aire, from streams draining the Yorkshire Wolds (Fig. 1A) and from periglacial slope processes cannot be discounted. Although these sources would always have been present during deposition of the Park Farm Clay Member, the observed upward increase in the amount of sand at both Hemingbrough and Glade Farm (Fig. 3) is consistent with the dominant sand source derived from the Escrick moraine ridge, when the ice margin was at or near it. Thus, dates on the sand layers are taken to approximate the age of the moraine. 6.3. LGM ice limit The weighted-mean OSL age of 22.2  0.5 kyr from Hemingbrough firmly places the low-level phase of Lake Humber within the LGM. The emplacement of the sandy beds also corresponds, within error, to the initial pulse of deglaciation of the BIS identified by Bowen et al. (2002) at 21.4  1.3 kyr. This is contrary to two ice advances, proposed by Bateman et al. (2008), first, to at least 15 km beyond the Escrick moraine ridge after 23.3  1.5 kyr but before 20.5  1.2 kyr, and subsequently to Wroot-Thorne. The absence of geomorphic features associated with glacier surging (cf. Smith, 1990) or diamicton in the sediments observed at both Hemingbrough and Glade Farm is consistent with the age model proposed here. The Hemingbrough OSL ages and sedimentary interpretations corroborate the conceptual model of Ford et al. (2008, Fig. 4), which denotes the Escrick moraine ridge as the LGM ice margin in the Vale of York. It is more difficult, however, to reconcile the Hemingbrough OSL ages of 21.0  1.9, 21.9  2.0, and 24.1  2.2 kyr with radiocarbon ages of 22,043  497 cal yr BP and 21,734  372 cal yr BP from the Dimlington Silts (Penny et al., 1969), since the latter ages imply that Holderness was ice-free at this time; thus, what was blocking the Humber Gap? Numerical simulations of the BIS (Boulton and Hagdorn, 2006) during the Dimlington Stadial reconstruct fastflowing ice streams down the east coast of England and in the Vale of York, although their timing is rather loosely constrained. If all of the above ages are correct, then their concordance, within error, may indicate that surging of the east coast ice stream onto Holderness occurred c. 22 ka or shortly afterwards, immediately after the period

of ice-free conditions. This advance, we suggest, was more or less coeval with advance of the Vale of York ice to the Escrick moraine ridge. Another factor, however, that may have contributed to ponding of proglacial water during the low-level phase of Lake Humber was glacio-isostatic depression towards the north. 6.4. High-level lake phase The weighted-mean OSL age of 22.2  0.5 kyr for the low-level phase of Lake Humber conflicts with the OSL age of 16.6  1.2 kyr obtained by Bateman et al. (2008) from sand attributed to the earlier, high-level phase. These authors inferred that the high-level phase did not occur until 16.6  1.2 kyr, with the low-level phase and its subsequent rapid drainage complete by c. 14 kyr. One explanation for these conflicting ages is that the dated sand interpreted as distal beach material by Bateman et al. (2008) may record reworking by colluvial processes (cf. Carruthers, 1947). Coeval OSL ages of 17.7  1.2 kyr derived from periglacial deposits at c. 34 m OD at Caistor (Bateman et al., 2000), and 17.5  1.6 kyr at Eppleworth on the dipslope of the Yorkshire Wolds (Wintle and Catt, 1985) are consistent with this alternative interpretation. These three ages broadly correspond with the end of Greenland Stadial 2b, a period of probable active-layer deepening and enhanced periglacial slope processes in southern England (Murton et al., 2003), and the first part of Greenland Stadial 2a (Lowe et al., 2008). 7. Conclusions The uppermost 10 m of glaciolacustrine sediments of the Hemingbrough Fm exposed in the central part of the Vale of York record the depositional conditions during the latter part of the lowlevel phase of Lake Humber. The interbedding of wave ripples in laminated clayey silts suggests that lake levels repeatedly fluctuated between shallow water to emergent conditions that were characterised by wave-worked sands and deeper water characterised by suspension settling of silt and clay. Ground freezing and brecciation of silty deposits occurred intermittently during low lake levels. These interpretations question the concept of a single lake level at c. 8 m associated with the low-level phase, and indicate that quiescent depositional conditions in standing water (Gaunt, 1994) were repeatedly interrupted by high-energy wave reworking. Wave-rippled sandy beds become more prominent towards the top of the Lake Humber deposits and indicate a nearby source of relatively coarse sediment that is attributed to the Vale of York ice lobe, when its margin stood at or near to the Escrick moraine ridge. Three OSL ages obtained from two sandy beds at Hemingbrough provide a weighted-mean age of 22.2  0.5 kyr, placing the lowlevel phase during the LGM. They support the conceptual model proposed by Ford et al. (2008) of a single ice advance into Lake Humber during the LGM. This study demonstrates the applicability of OSL dating as a viable chronological tool for glaciolacustrine sediments deposited in shallow water. Its wider application may further constrain paleoenvironmental change in the Humber Basin. Although the lowermost stratigraphy of the Hemingbrough Fm has yet to be described and analysed in detail, OSL dating may be of benefit to address currently unresolved issues: (1) to constrain more robustly the duration of Lake Humber; (2) to determine whether ice advanced south of the Escrick moraine ridge prior to the LGM. Acknowledgements Julian Slater and Mike Page of Plasm or are thanked for access to Hemingbrough and Glade Farm. The Selby District mapping team from the BGS, led by Tony Cooper, are thanked for earlier

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