Ice stream influence on West Greenland Ice Sheet dynamics during the Last Glacial Maximum

Ice stream influence on West Greenland Ice Sheetdynamics during the Last Glacial MaximumDAVID H. ROBERTS,1* ANTONY J. LONG,1 BETHAN J. DAVIES,1 MATTHEW J. R. SIMPSON2 and CHRISTOPH SCHNABEL31 Department of Geography, Durham University, Durham, UK2 Department of Earth Sciences, Durham University, Durham, UK3 NERC Cosmogenic Isotope Analysis Facility, Scottish Enterprise Technology Park, East Kilbride, UK

Roberts, D. H., Long, A. J., Davies, B. J., Simpson, M. J. R. and Schnabel, C. 2010. Ice stream influence on West Greenland Ice Sheet dynamics during the Last GlacialMaximum. J. Quaternary Sci., Vol. 25 pp. 850–864. ISSN 0267-8179.

Received 22 April 2009; Revised 28 September 2009; Accepted 30 September 2009

ABSTRACT: This paper investigates the processes governing bedrock bedform evolution in ice sheetand ice stream areas in central West Greenland, and explores the evidence for a cross-shelf ice streamat the Last Glacial Maximum (LGM). To the east of Sisimiut the formation of streamlined bedforms withhigh elongation ratios and high bedform density has been controlled by geological structure andtopography in slow-flowing ice sheet areas. At the coast, the effects of regional flow convergence,caused by coastal fjord orientation, routed ice into the Sisimiut/Itilleq area where it formed an icestream onset zone. This funnelled ice into an offshore trough (Holsteinsborg Dyb), resulting in asouthwesterly regional ice flow direction and the formation of a topographically routed ice stream(Holsteinsborg Isbrae). To the south of this, striae and bedform evidence show that local valley glaciersinitially flowed east to west across the coast, but were later redirected by the Itilleq Fjord ice whichturned southwestward due to diffluent flow and deflection by Holsteinsborg Isbrae. Roches mouton-nees in this area have low elongation ratios and high bedform density, but do not provide unequivocalsupport for ice streaming, as they are a product of both bedrock structure and changes in ice flowdirection, rather than enhanced flow velocities. Cosmogenic surface exposure ages limit maximum icesheet surface elevation to ca. 755–810m above sea level in this region. Such ice thickness enabledHolsteinsborg Isbrae to reach the mid/outer continental shelf during the LGM, and to contribute to theformation of a trough mouth fan and the Outer Hellefisk moraines. Initial deglaciation across thisregion was driven by rising sea level and increasing air temperatures prior to the Bølling Interstadial atca. 14.5 cal. ka BP. Between 12 and 10 cal. ka BP both increased air and ocean temperatures post theYounger Dryas, and peak sea-level rise up to the marine limit, caused accelerated thinning andmarginal retreat through calving, although dating evidence suggests ice streams remained along theinner shelf/coast boundary until at least ca. 10 cal. ka BP, their longevity maintained by increased icethickness and ice discharge. Copyright # 2009 John Wiley & Sons, Ltd.

KEYWORDS: Greenland Ice Sheet; Last Glacial Maximum; ice sheet dynamics.

Introduction

Recent research in formerly glaciated high-latitude regions hashighlighted the significant influence of ice streams on ice sheetdynamics during both advance and retreat phases. AroundAntarctica and the Arctic Basin, ice streams formed majorarteries evacuating inland ice to the shelf edge, controlling ice,fresh water and sediment flux to the oceans and significantlyinfluencing ice sheet mass balance and response times toenvironmental forcing (MacAyeal, 1993; Andrews andMaclean, 2003; O Cofaigh et al., 2003; Sejrup et al., 2003;Stokes et al., 2005). Such ice stream systems leave distinctive

geomorphic footprints with onset zones marking convergentflow and accelerating ice velocities, trunk zones withmegascale glacial lineations and marginal shear zones markingpeak velocity ice stream corridors, and trough mouth fans or tillwedges forming down ice sediment depocentres (Stokes andClark, 1999; Canals et al., 2000; Lowe and Anderson, 2002;Clark and Stokes, 2003; Ottesen et al., 2008).

In Greenland, however, little is known of the influence of icestreams on overall ice sheet dynamics, and there is a severepaucity of information detailing ice stream behaviour offshoreduring the Last Glacial Maximum (LGM) (Sisimiut Stade;Funder, 1989). Onshore there have been studies of thecontemporary behaviour of Jakobshavns Isbrae, Helheim andKangerdlussuaq glaciers, all of which are topographicallycontrolled by over-deepened troughs (Echelmeyer and Harri-son, 1990; Iken et al., 1993; Truffer and Echelmeyer, 2003;Howat et al., 2005; Joughin, 2006; Luckman et al., 2006), but

JOURNAL OF QUATERNARY SCIENCE (2010) 25(6) 850–864Copyright � 2009 John Wiley & Sons, Ltd.Published online 29 December 2009 in Wiley InterScience(www.interscience.wiley.com) DOI: 10.1002/jqs.1354

*Correspondence to: D. H. Roberts, Department of Geography, Durham Uni-versity, Science Laboratories, South Road, Durham DH1 3LE, UK.E-mail: [email protected]

only Roberts and Long (2005) andWeidick and Bennike (2007)have attempted to reconstruct ice stream behaviour throughoutthe LGM, reporting the thickening and advance of JakobshavnsIsbrae on to the shelf and its subsequent retreat (Long andRoberts, 2003; Roberts and Long, 2005). Bathymetric datasuggest that up to six regional cross-shelf troughs funnelled icestreams on to the western sector of the shelf (Fig. 1), but the onlyreports of ice stream activity offshore come from fragmentarygeomorphological evidence on the northeast and southeastGreenland shelf (Lykke-Anderson, 1998; Solheim et al., 1998;O Cofaigh et al., 2004; Roberts, et al., 2008; Evans et al., 2009).Our understanding of conditions beneath ice streams has

been mainly constructed from palaeo-ice stream imprints, withsoft-bedded, geomorphic evidence providing the major sourceof information (Stokes and Clark, 1999), but there have beensome limited analyses of rigid-bed, bedrock bedforms in areasof fast ice flow. In many cases fjords act as tributaries, oftenfeeding topographically confined ice streams on to the shelf,where convergence and coalescence lead to the formation oflarge cross-shelf composite ice streams (e.g. Ottesen et al.,2008). In such settings, Lowe and Anderson (2002) and OCofaigh et al. (2002) have reported the occurrence of cruderock drumlins, crag and tail forms and bedrock cut meltwaterchannels at the interface between onset and ice stream trunkzones, particularly in areas underlain by crystalline bedrockand devoid of soft sediment sources. Ottesen et al. (2008) alsoreport similar features in northern Norwegian fjords withstreamlined bedrock drumlins and crag and tails resulting fromhigh basal shear stresses and enhanced basal erosion, often besthighlighted in areas with strong geological structure withclosely spaced lithologies of contrasting competence. Bradwellet al. (2008) have also reported well-developed, elongatedbedrock ridges and megagrooves in northwestern Scotlandresulting from ice stream tributary flow, which they attribute toenhanced basal abrasion, with only a minor bedrock structuralinfluence.In some instances, inferences have beenmade about ice flow

velocities and ice/bed interface conditions using rigid bedform

morphology. In British Columbia, Evans (1996) used differ-ences between roches moutonnees (RMs) and whalebacks(WBs) to reconstruct evidence for bed separation beneath ice,which he used to infer slow versus fast ice flow. This work wasdeveloped by Roberts and Long (2005), who reported acorrelation between higher densities of WBs and increased icethickness beneath the fast-flowing margins of JakobshavnsIsbrae. Roberts and Long (2005) also hypothesised that lowbedform elongation ratios (ELRs) (<5:1) beneath isbraes in rigidbed settings result from fixed basal perturbations and higher bedroughness which prevent the basal ice from maintaining astable flow pathway. This contrasts with other recent work inrigid bedding settings which reported high ELR values (>5–25:1), inferring basal ice flow pathways can remain stable overrough beds (Bradwell et al., 2008), although inmany areas ELRscan also be lowered by secondary plucking during deglacia-tion, or a later glacial event (Roberts and Long, 2005).Here we investigate the evidence for ice sheet and ice stream

activity in central West Greenland during the LGM. First, weassess topographic and geological controls on regional ice flowpathways and consider the evidence for the inception of an icestream in this area. We then test previous assertions thatbedform morphology can be used to distinguish zones of slowversus fast ice flow, and use bedform long axes, striae andplucking data to reconstruct changing ice flow phases. Finally,we consider the influence of the ice stream on local glacierdynamics and regional ice sheet behaviour during the LGM anddeglaciation.

The glacial history of West Greenland

The earliest glacial deposits in West Greenland belong to theFiskebanke glaciation (MIS 6; 130 ka), but these are onlypartially preserved on coastal mountains (Kelly, 1985; Funder,1989). During the LGM ice sheet elevation was �500–810m

Figure 1 A topographic and bathymetric overview of the west coast of Greenland (source: IBCAO). Note the bathymetry of the continental shelf,which is cross-cut by a series of E–W-trending troughs that terminate in some instances in trough mouth fans. The area highlighted is denoted in Fig. 2

Copyright � 2009 John Wiley & Sons, Ltd. J. Quaternary Sci., Vol. 25(6) 850–864 (2010)

ICE STREAM INFLUENCE ON W GREENLAND ICE SHEET DYNAMICS DURING THE LGM 851

above sea level (a.s.l.) (Rinterknecht et al., 2009; Roberts et al.,2009), with pre-LGM nunataks characterised by block field andtors (Kelly, 1985; Roberts et al., 2009), and LGM terrain heavilyabraded by warm-based ice. Based on trimline evidence, Kelly(1985) proposed a relatively thin LGM ice sheet, which onlyreached the mid-shelf and the Inner Hellefisk moraines (Brettand Zarudzki, 1979; Funder, 1989) (Fig. 2). In contrast, Tarasovand Peltier (2002) suggested ice was 1100–1300m thick in theKangerlussuaq area, an estimate confirmed by Bennike et al.(2002) and Weidick et al. (2004), who reconstructed thicknessto �1500m in the southwest, inferring burial of many coastalmountains and ice advance to the comparatively narrow shelfedge. A new ice sheet model also projects the western sector ofthe Greenland Ice Sheet (GIS) at least to a mid-shelf positionand perhaps as far as the Outer Hellefisk moraines at the shelfedge (Simpson et al., 2009). The extent and behaviour of icestreams during the LGM, in particular large shelf-basedcomposite ice streams systems, remain poorly understood.

Roberts and Long (2005) and Weidick and Bennike (2007)proposed the coalescence of several outlet glaciers withJakobshavns Isbrae as it flowed westward across Disko Bugt,and the subsequent formation of a large composite ice streamthat traversed the outer shelf through the Outer EgedesmindeDyb to form a distinctive troughmouth fan, but other analogousoffshore features in West Greenland remain unexplored (Figs 1and 2).

Funder and Hansen (1996) suggested a two-stage model fordeglaciation, with initial ice sheet retreat across the shelf inWest Greenland occurring between 16.8 and 11.1 cal. ka BP asthe marine-based portion of the ice sheet was destabilised byrising global sea level, and the ice margin reaching the presentcoast by ca. 10 cal. ka BP (Funder and Hansen, 1996; Bennikeand Bjorck, 2002, Long and Roberts, 2003). This is supportedby dates on marine shells in the Sisimiut area which fix iceretreat to the coast to 10.4–10.2 cal. ka BP (Weidick, 1972;Kelly, 1973, 1979; Fig. 3). The second stage was typified by

Figure 2 The central west coast of Greenland. Note the position of the present ice sheet margin and fjord systems trending E–W. The shelf area iscross-cut by several deep troughs which extend towards the shelf edge. Note also the position of the Inner and Outer Hellefisk moraine complexes(Brett and Zarudzki, 1979)


852 JOURNAL OF QUATERNARY SCIENCE

Figure 3 An overview of the Søndre Strømfjord andNordre Isotoq region. Note the regional topographywith deep fjords running E–W andNE–SW ascontrolled by the regional geology. Themountainous areas of Nordre Isotoq and Sukkertoppen can be seen to the north and south respectively.Much ofthe ice-free terrain between the present ice margin and the coast is heavily glacially scoured. Existing 14C deglacial dates (cal. ka BP) for the coast arealso shown (source: Weidick, 1972; Kelly, 1973, 1979) (satellite image source: Google Earth – � 2009 Google Image IBCAO; � 2009 EuropaTechnologies; Image � 2009 Terra Nova)

Figure 4 A geological overview of the study area. The bedrock geology of this region is Precambrian Archaen gneiss to the south, withPalaeoproterozoic rocks to the north. This alkaline province of southern West Greenland includes swarms of alkaline ultramafic dykes describedas kimberlites and lamproites (Larsen, 1991). The foliation and dyke swarms exhibit a strong ENE–WSW trend



retreat inland and with the ice reaching the present ice sheetmargin position by ca. 8–7 cal. ka BP (Kelly, 1985; Bennike andBjorck, 2002; Rinterknecht et al., 2009).

Study site

The Sisimiut/Itilleq area lies �200 km to the west of the presentice margin between Søndre Strømfjord and Nordre Isotoq(Fig. 3). This region has been dissected over multiple glacialcycles by SW-flowing outlet glaciers from the main ice sheet(Fig. 3). To the northeast and southeast deep troughs (NordreIsortoq, Søndre Strømfjord) dissect the landscape and themountains reach elevations up to 1500m a.s.l. The mountainsharbour high-level trimlines, blockfield and tors lying aboveglaciated terrain (Kelly, 1985; Roberts et al., 2009). Thebedrock geology of this region is Precambrian Archaen gneissto the south, with Palaeoproterozoic rocks to the north. Thisalkaline province of southwest Greenland includes swarms ofalkaline ultramafic dykes described as kimberlites andlamproites (Larsen, 1991). The foliation and dyke swarmsexhibit a strong ENE–WSW trend (Fig. 4).Three large fjords – Kangerluarssuk, Ikertooq and Amerdloq –

are located to the north of Itilleq Fjord, with Ikertooq coincidingwith a major thrust and high strain zone that runs WSW fromthe ice sheet margin through to Holsteinsborg Dyb (Fig. 4), adeep WSW-trending trough that runs 60 km into the middleshelf before shallowing to a trough mouth fan at the shelf edge(Fig. 2). South of Itilleq Fjord, the Qingartaq coastal plain/strandflat is composed of glacially scoured terrain that rarelyrises above 100m a.s.l. East of this, the terrain rises steeply intoa mountainous alpine landscape (�1000–2000m a.s.l.)situated to the west of Sukkertoppen ice cap (Fig. 3). Thesemountains feed valley glaciers westward down on to the coast.

Methodology

Base maps were established using 1:150 000 scale aerialphotographs and supplemented with field mapping of glaciallandforms across the Qingartaq coastal plain and adjacentvalley areas. Satellite images (source: Google Earth) were usedto provide regional topographic and geomorphological over-views. Bedrock bedforms and striae were mapped usingestablished criteria for the identification of subglacial erosional(Gordon, 1981; Glasser andWarren, 1990; Sugden et al., 1992;Roberts and Long, 2005).Bedrock ridges and bedforms were classified as mega

(1000m length), macro (100m length), meso (100–10mlength) or micro (<10m length) in size. Length, width andheight measurements were taken to calculate ELRs.Where RMsor WBs formed dense swarms of bedforms they have beenreferred to as multi-swale complexes (Roberts and Long, 2005).Bedform density refers to number (n) of bedforms km�2, whiletransverse wavelength (TW) refers to the lateral distancebetween bedform crests. Measurements on the orientation ofbedform long axes were also taken and where bedform longaxes were modified by changing ice flow direction theorientation of plucked faces and striae were recorded. Meanresultant bedform orientations were computed using Geoori-ent. Joint dip and dip direction data have been plotted usingequal-area stereonets.Recessional moraines andmarine limits were identified using

established field criteria (Donner and Jungner, 1975; Lindenand Moller, 2005; Long et al., 1999, 2008). Trimlines were

identified as limits separating glacially eroded terrain fromshattered periglacial blockfield areas (Kelly, 1985; Ballantyne,1997). To establish the glacial history of the area and testwhether trimlines represented former ice sheet surface altitude,englacial thermal boundaries (with upper blockfields beingprotected by cold ice) or postglacial weathering limits, sampleswere collected for 10Be and 26Al exposure surface dating toestablish exposure and burial histories. All cosmogenic datesreferred to herein are reported in Roberts et al. (2009). Whenreferring to exposure ages directly, they are presented in ’ka’,meaning thousand years before the date of sample collection(AD 2006). Radiocarbon dates are reported as calibrated ages(cal. ka BP). Altitudes are expressed relative to present sea level(a.s.l.) and approximated to within �10m using a handheldglobal positioning system.

Results

Regional ice flow patterns: Sisimiut to thepresent ice margin

The landscape to the east of Nordre Isotoq and Ikertooq ischaracterised by glacially streamlined, elongated megascalebedrock hills and ridges separated by elongated depressionsand ribbon lakes (Figs 3 and 5). Their trajectory follows aW/SWpath closely mirroring the underlying gneissic geologicalstructure, which is heavily foliated, thrust and intruded byultramafic dykes (Fig. 4). In a number of areas, closely spacedridges are often separated by crude runnels or megagrooves(Fig. 5(b), (d) and (e)). These streamlined hills and megascaleglacial lineations are often up to 5 km long, but rarelymore than200m wide. Their elongation ratios vary between 2 and 40:1,with an average of 10:1 (n¼ 50) (Table 1). Bedform densityvaries with the underlying strata and degree of ridgefragmentation, but in some areas density peaks at 25–30bedforms km�2. The average transverse wavelength of bed-forms in this area is 280m, varying between �100 and 550m.

The overall SW pattern of ice flow is further enhanced at thecoast by the convergent nature of Kangerluarssuk, Ikertooq andAmerdloq Fjords, and the confinement of this ice flow by themountains bordering Itilleq Fjord and Nordre Isotoq (Fig. 3).This topographic configuration focused ice flow towardsHolsteinsborg Dyb. Ikertooq Fjord has a convergent easternend splitting into three tributaries which capture ice flow acrossa 20 km corridor. Kangerluarssuk and Amerdloq Fjords borderIkertooq Fjord south and north, and together all three act astributaries funnelling ice over a 40 km wide corridor towardsHolsteinsborg Dyb.

To the south of this, ice flow was directed SW by SøndreStrømfjord and E–W by Itilleq Fjord. The E–W trend of ItilleqFjord directs ice flow on to the shelf, although the Qingartaqcoast to the south is heavily streamlined and exhibits a suite ofglacially eroded bedforms that show Itilleq Fjord ice flowed SW(Fig. 6).

Qingartaq coast geomorphology

Two ice flow phases are discernible across the Qingartaqcoastal strandflat from striae and bedform evidence. The first setof striae show an E–W trend and relates to an initial advance oflocal valley glaciers fed from Nattoralinnguaq and adjacentvalleys (Fig. 6). There are also small areas of RMs whoselong axis orientations may relate to this early flow phase (Fig. 6).



This evidence is heavily overprinted, however, by a strong NE–SW flow phase which is supported by ubiquitous striaesuperimposed over a dominant suite of SW-trending mega- tomacroscale bedrock ridges superimposed by smaller scale RMcomplexes (Fig. 6).The local pattern of recessional moraines which bend SW

from the mouth of the Nattoralinnguaq valley show that theNattoralinnguaq glacier was deflected, but not completelytruncated, by SW-flowing ice (Fig. 6). These moraines arecomposed of grey glaciomarine sands and silts with in situpairedmarine bivalves (e.g.Hiatella arctica,Mya truncata). Theridges are slope conformable and uniformly spaced, being up to100–300m long, 10m high, 20–30m wide and typically 50mapart. Occasionally they are on-lapped by small deltas. Theyare also formed below the local marine limit, which isdemarcated by a perched boulder limit at �140m a.s.l. Theyhave been interpreted as De Geer moraines (Roberts et al.,2009).

Qingartaq coast bedforms

The largest streamlined bedrock ridges on the Qingartaq coastare over 1 km in length and up to 200mwide and 30–50m high(Fig. 6). Their ELRs range from 1 to 11:1 with an average of 5:1(n¼ 50). They are often heavily fragmented and denuded.Superimposed upon these megascale bedforms are smaller,heavily fragmented meso- and micoscale RMs. The detailedcharacteristics of these have been recorded in six sub-areas(Fig. 6 and Table 1). In all sub-areas RM long axes runpredominantly NE–SW in accordance with regional striae data(Fig. 6 and Table 1), although in sub-area 6 bedform long axesswing more N–S in correspondence with underlying bedrockstructure. Although very fragmented and often plucked in stoss,mid and leeside localities, bedforms tend to form crude RMsrather than WBs, and are abraded up-ice and plucked down-ice. Micro to mesoscale bedforms average 10–14m in lengthand rarely exceed 30m. Widths are typically 3–5m, with

Figure 5 A landscape of areal scour controlled by bedrock structure. Parts (a), (b), (d) and (e) show dominant southwest-trending megascale bedrockridges and megagrooves with high elongation ratios. The ridges, grooves and valleys are caused by the preferential erosion of different lithologieswithin the gneissic bedrock. Part (c) also shows the influence of topographic routing of ice by topography and bedrock structure and the preferentialfocusing of fluvioglacial sediments which aid basal abrasion of megagrooves along predetermined flow lines (satellite image source: Google Earth –� 2009 Google; Image � 2009 Digital Globe; Image � 2009 Terra Nova). This figure is available in colour online at www.interscience.wiley.com/journal/jqs



occasional bedforms reaching 10m wide. Height is generally1–3m. Elongation ratios tend to be less than 5:1 (Table 1).Striae and RM long axes data show a predominant NE–SW

ice flow direction across the Qingartaq coast, and this iscomplemented by the orientation of RM plucked faces. Manyplucked faces are orientated ESE–WNW, transverse to ice flow,

although occasional faces trend N–S (Fig. 6). There is often aclear correlation between RM long-axis orientation andunderlying bedrock structure, and plucked face symmetryand joint orientation. This is well demonstrated in sub-area 3,where the NE–SW andWNW–ESE joint systems are particularlyclear (Fig. 7(a)), and in sub-area 6, where plucked face

Figure 6 Glacially streamlined terrain across the Qingartaq coastal lowlands. Large bedrock ridges and RMs show a dominant SW ice flow directionoverprinting an earlier E–W ice flow phase sourced from the local mountains to the east. The lower panel shows RM long axes, striae, bedrock jointorientations and RM plucked face orientations for all bedforms measured. Localised sets of RM long axes are shown as small rose diagrams (sub-areas1–6). Note the well-developed De Geer recessional moraine train that infers deflection of local glaciers to the southwest by Itilleq Fjord ice



orientations highlight changes in ice flow direction. In sub-area6, two striae sets suggest an E–W (2898) ice flow and a later NE–SW (2358) ice flow (Fig. 7b). RM long axes trendmoreN–S here,pointing to the strong influence of the underlying bedrockstructure and joint system on bedform morphometry, but thesebedforms also exhibit two clear plucked faces (Fig. 7(b)). Thewesterly plucked faces relate to initial ice movement east towest. These have then become sub-parallel to flow, as the iceturned southwestward, whereupon a second set of pluckedfaces trending NNW–SSE was formed.Bedform density along the coast can be approximated at

different scales. Using aerial photograph sources mega- andmacroscale bedforms rarely exceed 25–40 bedforms km�2. Thetransverse wavelength of megascale ridges averages 190m, andvaries between �100 and 450m. From field survey, meso- tomicroscale bedforms can exceed 200 bedforms km�2, particu-larly in multiswale areas that occur over megascale bedforms. Inareas between the bedrock ridges, bedform density is muchlower and, hence, there is a clear coincidence between thepositive relief bedrock ridges and bedform position and density.

Ice sheet thickness

Surface exposure ages limit the Itilleq ice stream surface to amaximum of �750–810m a.s.l. during the LGM, withblockfield terrain above this dating to MIS 6 (Fig. 8 andTable 2; Roberts et al., 2009). Ice thinning from the LGM beganat ca. 21 ka, when ice sheet surface elevation was 570–500ma.s.l. (Fig. 8 and Table 2). This evidence limits maximum icethickness along the Qingartaq coast to �650–710m, as thecoast here rarely exceeds 100m a.s.l. in elevation. Based onbathymetric data, ice thickness in the fjords would have beengreater, with Itilleq Fjord ice a minimum of 1000m thick, andice in Holsteinsborg Dyb at least 1200m thick (Fig. 8).

Discussion

Regional ice sheet processes: areal scour,selective linear erosion and bedform evolution

The terrain between Nordre Isortoq, Søndre Strømfjord and thepresent ice margin is characterised by areal scour and selective

linear erosion (Sugden, 1974), which trends predominantlyENE–WSW (Fig. 3). During successive glacial cycles ice flowtrajectories have preferentially followed this route for a numberof reasons. The first is the simple gravitational spreading of theice sheet in an E–W direction during ice sheet build-up. Thesecond is the control exerted on basal ice flow direction bybedrock structure, with the ENE–WSW orientation of thethrusts, foliation and alkaline ultramafic intrusions within thePrecambrian Archaen gneiss complex guiding overall ice flow(Fig. 4). Lithological and structural contrasts within the bedrockenabled basal ice to preferentially erode the bedrock surfaceinto a series of elongate ridges and lineations, and crudemegagrooves, with more resistant bedrock bands forming theridges.During successive glacial advances the lithological and

structural properties of the bedrock have controlled bedformevolution. In particular, the grooves become pathways forenhanced flow, with an impermeable rigid bed, strain heatingand focused basal meltwater routing inducing basal slidingand increased abrasion across the entire regional ice/bedinterface. The mechanical ability of the ice to abrade thesetroughs is further enhanced by the deposition of fluvioglacialmaterial (Fig. 5(c)), both subglacially and in intervening non-glacial periods, and its subsequent entrainment. This con-centrates and locks debris into predefined flow paths enablingfurther trough erosion (cf. Bradwell et al., 2008). Pressurisedsubglacial meltwater flow may also play a part in trough/groove formation.Even though the development of this well-lineated, fixed

ridge and groove system is not an ice stream-relatedgeomorphic imprint, it does have a bearing on the theoreticalprocess of groove ploughing (Tulacyzk et al., 2001; Clark et al.,2003). The development of this grooved terrain throughsuccessive glacial cycles would allow basal ice to maintaina keeled morphology over kilometres. The sinuous shape ofmany grooves and ridges infers that keels follow pre-existingbedrock structure and topography (Fig. 5(e)), and if they were toencounter a soft-bedded medium down flow such keels couldinitiate glacial lineations.

Ice flow convergence and ice stream onset:Holsteinsborg Isbrae

The overall pattern of ice flow between the present ice sheetmargin and Sisimiut begins to converge at the coast, and there is

Table 1 Bedrock bedform data

Area n Bedform Striae ELR Height(m)

Density TW Comment

Sisimiut to presentice margin

50 245 – 10 200–400 25–30 280 Megascale streamlined bedrock ridges

Qingartaq coast 50 225 – 5 30–50 25–40 190 Mega- to macroscale streamlined bedrock ridgesQingartaq Sub-area 1 30 228 230 4.5 1.5 200 – Meso- and microscale RMs forming multiswale topography

over macroscale streamlined ridgesQingartaq Sub-area 2 30 225 226 4.6 1.2 200 – Meso- and microscale RMs forming multiswale topography




over macroscale streamlined ridgesQingartaq Sub-area 6 30 198 289 235 2.6 1.5 200 – Meso- and microscale RMs forming multiswale topography

over macroscale streamlined ridges. Double plucking of bedforms.Two distinct striae sets



strong evidence to suggest an ice stream operated offshorebetween Sisimiut and Itilleq (Fig. 9). Regional ice flow velocitywas increased by two processes. The first was the influence ofhigh mountains north of Sisimiut and bordering Itilleq Fjordwhich squeezed regional ice sheet flow laterally, and thesecond, and more influential factor, was the convergent natureof Kangerluarssuk, Ikertooq and Amerdloq Fjords which actedas a natural ice stream onset zone (Figs 3 and 9), funnelling iceinto Holsteinsborg Dyb to form an ice stream (HolsteinsborgIsbrae). The onset zone was �40 km wide north to south, andthe trunk zone was topographically routed because of thegeological guidance of the Ikertooq thrust zone.The ability of fjords to act as tributaries for shelf-based ice

streams is well documented (Evans et al., 2005; Stoker and

Bradwell, 2005; Ottesen et al., 2008). In Norway, seismic andmultibeam data from fjord troughs show that glaciallystreamlined bedforms form in inner fjord zones and markzones of increased ice flow velocity. These subglacial flowsystems were convergent and confluent down-ice withadjacent fjord systems and fed composite ice streams on theshelf during the LGM (Ottesen et al., 2008). Without high-resolution offshore data the inference that Holsteinsborg Dybharboured a large regional ice stream cannot be fullysubstantiated, but natural ice flow pathways and topographicfunnelling in this area provide the necessary onset conditions,and the drawdown of ice into fjord heads across a crystallinerigid bed would have increased ice flow velocity throughenhanced strain heating, meltwater generation and sliding, thusinitiating ice stream onset (Wellner et al., 2001; Hall andGlasser, 2003; Ottesen et al., 2008).

The offshore evidence for an ice stream operating inHolsteinsborg Dyb is partially supported by the morainedistributions reported by Brett and Zarudzki (1979) and theshelf bathymetry, which shows a small trough mouth fan at theexit to the dyb (cf. Clark and Stokes, 2003). For the most part,the Outer Hellefisk moraines lie sub-parallel to the shelf edge,but to the north of Holsteinsborg Dyb on the outer to mid shelfthey clearly run sub-parallel to the upper margins of the troughfor some 100 km (Fig. 2), marking the marginal position of atopographically constrained ice stream.

The Itilleq ice stream

South of Sisimiut, the streamlined terrain across the Qingartaqstrandflat shows that Itilleq Fjord ice flowed SW and formed thesouthern margin of the Holsteinsborg Isbrae as it advancedoffshore (Fig. 6). Three factors influenced this SW flow imprint:ice thickness, topography and drawdown, and geology.Diffluent ice flow and offshore regional ice flow were theprincipal controls on flow direction. With the elevation of thestrandflat rarely exceeding 100m a.s.l., and local ice in ItilleqFjord being �1000m thick (Fig. 8), ice was able to turnsouthwestward as it exited the western end of the fjord andbecame topographically unconstrained. Secondly, Itilleq Fjordice was deflected southward by Holsteinsborg Isbrae, whichcontrolled regional ice flow direction and drawdown on to theshelf (Fig. 9). The third, more localised influence on ice flowwas the distinctive SW-trending sub-vertical joint structure andalkaline ultramafic dykes within the Archaen gneiss, whichthrough preferential erosion guided and enhanced ice flowtowards the SW, and enabled the development of streamlinedbedrock terrain in a manner similar to that described above(Fig. 6). To define this southwesterly ice flow phase acrossQingartaq as an ice stream sensu stricto, however, is difficultwithout direct evidence for increased ice flow velocities.

Using bedform morphometry to reconstruct iceflow velocity

The bedforms reported here to the east of Sisimiut and on theQingartaq coast, and their associated ELRs, demonstrate theneed for caution when using bedforms to infer increased iceflow velocities and ice stream activity. In the area between thepresent ice margin and Sismiut, ice sheet processes (lowvelocity, inter-stream areas) governed bedform development,and in this particular setting bedform erosion and elongationwere strongly influenced by the underlying bedrock structure.

Figure 7 (a) RM long axes and joint orientations from sub-area 3. Notethe predominance of NE–SW and ESE–WNW joints that have aided thedevelopment of SW-trending RM long axes. (b) Data from sub-area 6.Note the SW-trending RM long axes, striae and joint and plucked faceorientations. An initial ice advance towards the west enabled pluckingof the western side of RMs, but the later switch to a southwest flowphase enabled secondary plucking along southwest-facing bedformedges



Figure 8 (a) Aerial photograph of the Nattoralinnguaq valley showing the positions of the surface exposure age samples. These border the edge ofboth the Itilleq ice stream and the Nattoralinnguaq valley glacier. Details of each sample are given in Table 2 (see Roberts et al., 2009, for a fulldescription). (b) Cross-section of the coastal terrain between Sisimiut and Qingartaq. The maximum LGM ice surface limit is constrained bycosmogenic ages to �750–810m a.s.l. on Nattoralinnguaq, although the lower limit of blockfield is �660m a.s.l., which corresponds to trimlineobservationsmade by Kelly (1985) in the vicinity of Sisimiut. Ice sheet thickness on theQingartaq coast would have been�650–700m,while in deeperoffshore troughs, such as Holsteinsborg Dyb and Itilleq Fjord, ice thickness would have been �1000–1200m



The high ELRs of these features are the product of repeatedsubglacial abrasion of bands of contrasting lithology duringsuccessive glacial cycles. Softer lithologies become mega-grooves and, with each successive glacial cycle, basal flowlines become more focused and entrenched, leading to arealscour and small scale selective linear erosion (Sugden, 1974).These features are not analogous to soft-bedded megascalelineations (cf. Stokes and Clark, 2002) and are not a product ofincreased basal abrasion due to increased ice flow velocities(cf. Bradwell et al., 2008).

The bedrock ridges across the Qingartaq coast differ fromthose to the east of Sisimiut. The Qingartaq megascalebedforms are shorter and more fragmented. Well-definedelongated ridge and groove topography is lacking, suggestingbasal ice flow pathways have not remained fixed. On a smallerscale, the ELR and bedform density values from the meso- andmicroscale bedforms across the Qingartaq coast are similar tothose reported from Jakobshavns Isbrae, where Roberts andLong (2005) hypothesised that low ELRs<5:1 and high bedformdensity (>200 km�2) beneath ice streams were a product offixed basal bedrock perturbations and high bed roughnesswhich limited the development of stable ice flow pathwaysand, therefore, reduced bedform length. However, there arefew well-defined WBs in the Qingartaq area region whichcould be used to provide additional evidence for suppressedbed separation under thick, fast-flowing ice stream conditions(Evans, 1996), and thus this bedform evidence does notunequivocally support ice streaming, despite the regionalgeomorphic imprint suggesting that Itilleq ice formed thesouthern margin of the Holsteinsborg Dyb ice stream.

It seems more likely along the Qingartaq coast that both thelack of WBs and the low ELR values are a function of bandedbedrock structure, joint frequency and orientation, andplucking (Fig. 7). Multiple WNW–ESE joints allowed pluckedfaces to develop transverse to the first ice flow phase from eastto west. This generated short, wide bedforms as the ice cross-cut the underlying gneissic bands and ultramafic dyke swarmsthat trend NE–SW (Fig. 10(b)). These short, wide bedforms thenbecame sub-parallel to flow, as the ice turned southwestward,whereupon a second set of plucked faces trendingW/SW–E/NEwas formed. This would have shortened both the originalbedform length and width, and reduced ELR values further(Fig. 10(c)).

Ice stream influence on regional ice sheetbehaviour during the LGM and deglaciation

Little is currently known of the form or influence of ice streamsystems on GIS behaviour during the LGM. Roberts and Long(2005) and Weidick and Bennike (2007) have speculated thatJakobshavns Isbrae expanded to the shelf edge during the LGM,but the Quaternary history of these shelf-based ice streamsystems remains to be explored. The recent estimates of icethickness along this coast (840–1000þ m; Rinterknecht et al.,2009; Roberts et al., 2009; Fig. 8) do, however, concur withrecent modelled ice sheet profiles for the west coast ofGreenland (Simpson et al., 2009), which reconstruct LGM icethickness over the Nattoralinnguaq and Sisimiut areas to�900m a.s.l., and demonstrate that the ice sheet attained atleast a mid-shelf position, and probably the Outer Hellefiskmoraines during the LGM (Roberts et al., 2009).

The Outer Hellefisk moraines to the north and south ofHolsteinsborg Dyb originated as part of the wider ice sheetLGM advance to the shelf edge, and Holsteinsborg Isbraewould have formed a coherent flow corridor within the iceT

able

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Internal

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Cb2676

6627.777

5325.913

856

1.408�106

3.655�104

11.82

0.273

119.8

3.2

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Cb2677

6627.783

5325.856

841

1.658�106

4.621�104

11.73

0.272

142.9

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6627.816

5326.385

755

6.047�105

2.117�104

10.80

0.264

55.40

1.97

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6627.809

5326.116

813

1.160�106

3.230�104

11.44

0.270

101.4

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6627.849

5326.516

734

3.179�105

1.118�104

10.58

0.262

29.53

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6627.793

5326.656

574

1.947�105

6.995�103

9.08

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21.01

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5327.219

498

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6.303�103

8.38

0.241

20.39

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8.07

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26.78

0.78

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6627.202

5326.102

359

1.387�105

5.040�103

7.35

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18.43

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6626.990

5326.613

230

1.066�105

3.360�103

6.35

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16.30

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6627.838

5329.317

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5.857�103

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2.720�103

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sheet during this phase, contributing to the formation of theterminal moraines and trough mouth fan at the western end ofthe trough (Fig. 2). However, the moraine situated sub-parallelto the north edge of the trough (Fig. 2) also clearly delimits alateral margin of the ice stream being topographicallyconstrained by the trough following the formation of the OuterHellefisk moraines and, thus, the ice stream operated inisolation at some point during ice retreat, or alternatively,during a later readvance. It has been speculated that the InnerHellefisk moraines may be of Younger Dryas age (Rinterknechtet al., 2009; Simpson et al., 2009), and so it is feasible that theice stream formed part of the wider ice sheet margin that re-expanded across the shelf during the Younger Dryas.During overall deglaciation, ice sheet thinning to the east of

Qingartaq began by ca. 21.0 ka, with the ice surface loweringfrom �570m in the Nattoralinnguaq valley (Roberts et al.,2009). This was followed by rapid thinning down to �150ma.s.l. by 13.6 ka, although the Itilleq ice stream may haveoccupied the coast later than this, with cosmogenic agesindicating the presence of ice until ca. 10 ka at elevations up to240–340m a.s.l. (Roberts et al. 2009) (Table 2 and Fig. 8). Thisoverall pattern of thinning coincides with increased airtemperatures over Greenland running up to the BøllingInterstadial (GIS1e) at ca. 14.5 cal. ka BP (Dansgaard et al.,1993) and points to surface ablation partially driving earlydeglaciation (Fig. 11).Funder and Hansen (1996) also advocated sea-level-driven

calving and thinning across the western Greenland shelfbetween ca. 16 and 11 cal. ka BP, although Simpson et al.

(2009) suggest that sea-level forcing of ice retreat was onlypronounced between 12 and 11 cal. ka BP when sea-level risepeaked at �100m a.s.l. (Fig. 11). A sea-level rise of suchmagnitude by 11 cal. ka BP (marine limit �100–140m a.s.l.;Ten Brink andWeidick, 1975; Roberts et al., 2009) would haverapidly destabilised the shallow inter-stream areas of the shelf,where ice thickness was less than 400m by this time (Fig. 8 andTable 2). In contrast, the ice streams may not have destabilisedas quickly due to increased ice thickness and ice discharge. Thecosmogenic ages adjacent to Itilleq Fjord indicate ice thicknesswould have been of the order ofþ800m prior to 10 ka, and it isnot until after this time that ice thickness in the troughmay havebeen less than 600m when, with water depths approachingþ500m, the ice stream may have become buoyant and rapidlycalved.The De Geer moraines marking the retreat of the Nattor-

alinnguaq valley glacier support a late-stage calving mechan-ism along the coast, with the southern margin of the Itilleq icestream succumbing to grounding line retreat up to the localmarine limit (i.e. ca. 11 cal. ka BP; Fig. 11) (Roberts et al.,2009). Post 10 cal. ka BP, increasing ocean temperatures willalso have accelerated ice retreat rates from the coast, with thewarm West Greenland Current invading coastal waters as itflowed northward towards Disko Bugt and northern Baffin Bay(Funder and Weidick, 1991; Lloyd et al., 2005; cf. Hollandet al., 2008).This pattern of ice stream retreat mirrors the deglaciation

pattern of Jakobshavns Isbrae, which remained on the shelfuntil ca. 10 cal. ka BP, with early thinning being driven

Figure 9 Reconstructed ice flow pathways during the LGM. Holsteinsborg Isbrae was fed from the onset zone formed by Kangerdluarssuk, Ikertooqand Amerdloq Fjords. These fjords funnelled ice into Holsteinsborg Dyb, where the trough confined flow and formed the trunk zone of an isbrae. Wehypothesise that other large fjords could have also formed distinct flow pathways within the ice sheet, which was composed of ice stream and inter-stream zones (satellite image source: Google Earth – � 2009 Google Image IBCAO; � 2009 Europa Technologies; Image � 2009 Terra Nova). Thisfigure is available in colour online at www.interscience.wiley.com/journal/jqs



predominantly by surface ablation, and not rapid marinedrawdown, as ice thickness and high rates of ice dischargenegated ice sheet lift-off (Long and Roberts, 2003). By ca.10 cal. ka BP the Disko Bugt ice sheet margin, including

Jakobshavns Isbrae, had retreated rapidly to the present easternshores of Disko Bugt, driven by increased surface ablation postthe Younger Dryas, calving triggered by bathymetric over-deepening (Long and Roberts, 2003, Long et al., 2003) andocean temperature forcing as the West Greenland Currententered Disko Bugt (Lloyd et al., 2005).

The behaviour of the topographically controlled ice streamsmay therefore have been fundamental in controlling andslowing the retreat of the wider ice sheet from shelf areas duringdeglaciation. Our topographic, bedform and ice thickness datasupport the existence of topographically constrained icestreams flowing onto and reaching the edge of the WestGreenland shelf during ice advance. We suggest that theadditional thickness and ice discharge routed through these icestreams was of sufficient magnitude to maintain grounded,quasi-stable ice margin conditions throughout early deglacia-tion, despite accelerated ice surface thinning and rising sealevels. Only by 10 cal. ka BP had the ice streams reached thepresent coast, finally succumbing to retreat off the inner shelf asa result of prolonged increases in air and ocean temperaturesand sea-level rise.

Conclusion

During successive glacial cycles in West Greenland, gravita-tional spreading, bedrock structure, areal scour and topo-graphic over-deepening guided ice from the present ice marginto the coast along an ENE–WSW flow trajectory. Megascalestreamlined bedrock ridges and grooves result from lithologicaland structural contrasts within the gneissic bedrock, with theabrasive capacity of the ice sheet enhanced by an impermeablerigid bed, strain heating, basal meltwater lubrication andsediment availability. Such bedforms are not the result of icestreaming.

At the coast, the convergent pattern of Kangerluarssuk,Ikertooq and Amerdloq Fjords funnelled ice into HolsteinsborgDyb to form an ice stream. To the south of this, Itilleq Fjord iceflowed southwestward across the coast, its trajectory controlledby diffluent flow due to a loss of topographic constraint anddeflection by Holsteinsborg Isbrae to its immediate north.

Bedrock bedform evidence partially supports ice streamingto the south of Itilleq Fjord, with dense populations(>200 km�2) of meso- and microscale bedforms forming overstreamlined hills. However, a lack of well-defined whalebackbedforms does not point to suppressed bed separationconditions under thick, fast ice flow conditions, and low ELRsprovide only equivocal evidence for ice streaming. Changes inice flow direction, bedrock structure and joint orientation, andplucking processes were the primary controls on bedformgeometry, rather than flow velocities.

Exposure ages limit the Itilleq ice stream surface to �750ma.s.l. during the build-up to the LGM. Taking bathymetric datainto account, ice in Holsteinsborg Dyb and Itilleq Fjord wouldhave been �1000–1200m thick, suggesting the ice sheet wascapable of attaining at least a mid-shelf position, and mostprobably the Outer Hellefisk moraines during the LGM. Thisassertion is supported by newmodelled profiles of the ice sheet,and the morphology of the shelf moraines peripheral toHolsteinsborg Dyb, which show ice terminating at the westernend of the trough during advance and, at times, an ice streamoperating independently.

The timing of outer shelf ice retreat remains poorlyconstrained, but the available evidence supports a possiblefour-phase model for deglaciation: (i) initial sea-level rise and

Figure 10 A schematic model for the development of streamlinedbedrock bedforms in areas of banded, contrasting lithologies. (a) NE–SW-trending bedrock ridges prior to ice overrun. (b) Roche moutonneeleeside plucked faces develop on the down-ice side of bedforms andalong joints running tranverse to ice flow in areas where ice flow istranverse or oblique to bedrock structure. Where lithological bands arenarrow this produces short, wide bedforms with low ELRs. (c) If ice flowswitches and becomes sub-parallel to lithological banding, newplucked faces form down-ice, while the original plucked faces becomepartially abraded in flank-side positions. This reduces both length andwidth (in the opposite direction to flow stage 1), resulting in shorter,smaller bedforms. Bedrock structure and changing ice flow directionthus control bedform evolution



air-temperature-driven thinning on the shelf in the run-up to theBølling Interstadial (ca. 16–14.5 cal. ka BP); (ii) a possibleYounger Dryas (12.8–11.5 cal. ka BP) ice sheet re-expansion onthe shelf driven by air temperature cooling; (iii) between 12 and10 cal. ka BP both increased air temperatures post the YoungerDryas, and peak sea-level rise up to the marine limit causedaccelerated thinning and marginal retreat through calving. By10 cal. ka BP the ice sheet margin had reached the presentcoast, where further retreat was enhanced by increasing oceantemperatures; (iv) between 10 and 7 cal. ka BP the ice sheetretreated a further 200 km eastward due to increased earlyHolocene air temperatures.

Acknowledgements This work was supported by Durham Universityand NERC CIAF grant no: 9035-0407. We would like to thank WillTodd, LeanneWake and SarahWoodroffe for their help in the field, andBrice Rea and Chris Stokes for their input and comments. The Designand Imaging Unit, Department of Geography, Durham University, isthanked for the production of figures.

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Ice stream influence on West Greenland Ice Sheet dynamics during the Last Glacial Maximum