Late Wisconsinan - Early Holocene deglaciation of Québec-Labrador Serge Occhietti ¹, É. Govare ¹, R. Klassen ², M. Parent ³ and J.-S.Vincent ² ¹ Département de géographie and Geotop, UQAM CP 8888 Centre-ville, Montréal, QC Canada H3C 3P8, E-mail: [email protected] ² Geological Survey of Canada, 601 booth street, Ottawa, Ont. Canada K1A 0E8 E-mail: [email protected] and [email protected] ³ Geological Survey of Canada, 2535 boul. Laurier, C.P. 7500, Sainte-Foy, QC Canada G1V 4C7 E-mail: [email protected]

Abstract During the Wisconsinan, the area of Québec-Labrador was completely covered by the Laurentide Ice Sheet (LIS), except for nunataks in the Torngat Mountains. For this reason, pre-Upper Pleistocene events are only documented in scattered stratigraphic sections, and by erosional glacial marks. After the Last Glacial Maximum (LGM), the Labradorean Sector of the LIS evolved from a single, predominant dispersal centre with subsidiary ice divides, into peripheral ice domes and masses which remained connected or not to the central dome. The central dome was not a simple and stable dome-shaped ice mass, but an evolving ice mass. Thinning of the LIS through ablation, and mechanical drawdown along its margins as the result of diachronic ice streams in the St. Lawrence Corridor and Hudson Strait are the main features of Late-Glacial ice flow dynamics. In the areas south of the St. Lawrence Corridor, ice masses over the Appalachian uplands evolved from a glacier complex confluent with the LIS into separate local ice caps. During a part of the warm Bølling-Allerød phase, a series of ice front positions mark a fast retreat of the ice front in the Appalachians of southern Québec, between about 11,900 (or as recent as 11,600) and 11,600 (or 11,300) conventional 14 C years B.P. (equivalent of ages of terrestrial fossils). The ice mass over the Canadian Shield, north of the St. Lawrence Corridor, dissipated slowly, between about 11,000 (or several centuries later) and 6,500 B.P. The deglaciation pattern includes the differentiation of an ice mass over the Hudson Bay, early deglaciation of the Labrador Highlands, a major change of ice flow in the Ungava Bay, and a very roughly concentric ice retreat pattern in the south-west, south and south-east margins of the remnant main ice mass. In the northern Ungava and Labrador peninsulas, major glacial lakes in low-lying areas were dammed between the ice front and the tilted deglaciated land. Lowlands depressed by glacioisostasy were momentarily invaded by marine waters, mostly between 13 ka and 7 ka. The last glacial ice masses were

located in the Labrador Trough and Nunavik and finally disappeared c. 6.5 ka.

Overview of the glacial history of Québec-Labrador Observational evidence suggest a very erosional Illinoian (s.l.) glaciation (Saalian s.l .) over all the territory of Québec-Labrador. The evidence for pre -Wisconsinan events is limited, including rare exposures in the St. Lawrence Valley, tills and stratified deposits encountered in boreholes in the St. Lawrence estuary (Occhietti et al., 1995) and the Appalachians of southern Québec (Shilts & Smith, 1986), and inter-till stratified units at one site in the Labrador Trough (Klassen et al., 1988). In the middle St. Lawrence estuary, unpublished seismostratigraphic data provide evidence for tunnel valleys incised into preIllinoian deposits and bedrock. During the last glaciation (the Weichselian Glacial which is the equivalent to the cold and cool phases of the Sangamonian and the Wisconsinan Glacial in Canadian terminology, Fulton, 1984), all the territory was glaciated during the Last Glacial Maximum (LGM) or earlier, except for small areas in the Torngat Mountains of northern Labrador and Québec. For Wisconsinan glacial history, three main areas can be distinguished: the Appalachian area, the St. Lawrence Corridor (including the St. Lawrence Valley, Estuary and Gulf), and the larger area over the Canadian Shield (Laurentians to the south, Abitibi to the west, New Québec to the north and Labrador to the east). Although each of the areas has a specific glacial style, the stratigraphic record of stadial and interstadial events in the St. Lawrence Corridor is critical because it has been affected by both the Canadian Shield ice dome and the Appalachian glacier complex. Early ice sheet inception over the Laurentian highlands (including the Manicouagan Plateau) is evidenced by glacial striations indicating divergent ice flow to the northwest and the south-east (Parent et al., 1995; Veillette et al., 1999). Although the age of this early glacial phase is not

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Fig. 1. Compilation of Late Wisconsinan - Early Holocene moraines and eskers in Québec-Labrador, mostly from the Atlas du Québec et de ses régions (Québec area) and from a compilation supplied by A. Moore (Labrador area).

Québec known, we infer that ice accumulated as early as Marine Isotope Substage 5d. There is strong evidence for extensive ice over the Canadian Shield during Marine Isotope Substage 5b. In the St. Lawrence Valley, early glacioisostatic subsidence followed by glacial lake inundation and glaciation is indicated by raised fluvial (Lotbiniere Sand), varved lake sediments (Deschaillons Varves) and till (Levrard Till) (Lamothe, 1989; Occhietti, 1982). A coeval (?) Appalachian ice cap over southern Québec deposited a local till (Chaudière Till, MacDonald & Shilts, 1972; Lamothe et al., 1992). Glacial cirques in highlands north of the middle St. Lawrence Estuary (Charlevoix area; Rondot, unpublished data; Govare, 1995) formed during repeated inception phases during the Pleistocene. From stratigraphic data in the St. Lawrence corridor and in the James Bay lowlands (Andrews et al., 1983), ice masses persisted in the Laurentians and New Québec during Marine Isotope Substage 5a. From Marine Isotope Stage (MIS) 4 to the end of MIS 2, most of New Québec remained glaciated, with limited deglaciation episodes in the Appalachians (Lake Gayhurst episode), marine episode in the estuary and gulf of St. Lawrence c. 35 ka (Gratton et al., 1984; Dionne & Occhietti, 1995), glacial lake in the upper St. Lawrence Valley (c f. Occhietti, 1989) and marine episodes in the James Bay lowlands (Andrews et al., 1983). The New Québec ice masses formed an irregular dome whose limits changed through time. The dome extended over the Labrador Trough, was connected to the Labrador highland glaciers, extended to the Hudson Bay, the St. Lawrence Valley, Estuary and Gulf, and the Great Lakes basins and lowlands. The Laurentide ice also flowed als o over the Appalachians uplands and highlands of Québec and New England. The New Québec dome was one of the major domes of the Laurentide Ice Sheet (LIS), and all the related autochtonous and allochtonous ice masses are referred to as the Labradorean Sector of the LIS (Prest, 1984; Fulton, 1989). The Late Wisconsinan (Late Weichselian)-Holocene deglaciation of Québec-Labrador: overview The digital map (see enclosed CD) comprises a sheet of the digital Atlas du Québec et de ses régions, prepared by É. Govare and S. Occhietti, and the digital map of glacial features of Labrador, compiled by R. Klassen and supplied by A. Moore (Geological Survey of Canada, Ottawa). It includes a systematic inventory of the eskers and moraines (Fig. 1) reported on the 143 topographic maps of Québec at the 1: 250,000 scale, in papers, public reports and established by original research (in bibliography). The digital contour of the moraines from the Appalachians of southern Québec (Fig. 2) was already prepared for a paper (Occhietti et al., 2001a). The outlines of the main moraines (Fig. 1) of the Canadian Shield area (Saint-Narcisse, Harricana, Sakami, North Shore, etc.) are drawn from

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published papers and special volumes. The Mars-Batiscan Moraine is outlined from unpublished data (Govare, 1995; Bolduc, 1995). This moraine is now extended to the west by 100 km (F. Robert, 2001, and Simard, 2003) (Fig. 2). The ice front positions of the Laurentide Ice Sheet on the Laurentian Plateau during the first centuries of the Holocene are mapped on a 400 km-long transect analysed by Simard (2003). The digital outlines of post-glacial glacioisostatic marine invasions and lakes (Fig. 3), extracted from the general maps of Dyke & Prest (1987), were kindly supplied by A. Moore. The deglaciation of Québec can be subdivided into three major phases: 1) the St. Lawrence Ice Stream phase and its consequences (from about 15,000 to 12,200 y. B.P.), 2) the Younger Dryas episodes (11,000 to 10, 000 y. B.P.), and 3) the early Holocene deglaciation. In south-eastern Québec, the margin of the Laurentide Ice Sheet covered a physiographically diverse region: the Appalachian Uplands in southern Québec and Gaspé Peninsula, the St. Lawrence Corridor, and the deeply-indented southern margin of the Laurentian Highlands. This landscape promoted deglacial styles that differed substantially from region to region. In northern Québec and in Labrador, a concentric thermo latitudinal retreat pattern (retreat rate depending on the distance from the centres of ice mass or from the ice divides over New Québec, and on the insolation) predominated, nevertheless with shifting centres of ice mass (Occhietti, 1983; Boulton & Clark, 1990; Parent et al., 1995, 1998; Clark et al., 2000). Glacial landform assemblages have been interpreted as broadly concentric zones comprising: Zone 1) the outermost-characterised by extensive end moraines, hummocky moraine, ice thrust masses, and several generations of ice flow lineaments; Zone 2) characterised by long eskers and contemporaneous ice flow lineaments, and Zone 3) the innermost, characterised by extensive ribbed moraine associated with drumlins and flutings (Dyke & Prest, 1987). These assemblages have been interpreted as recording deposition in the marginal zones of a retreating sheet, and to become younger toward the geographical centre of glaciation. From the striation evidence, and new mapping of the landform record, which includes evidence of glacial overprinting (e.g. Kleman et al., 1995; Veillette et al., 1999), the record requires reinterpretation in terms of its relative age, context in the ice sheet, and ice flow dynamics (c f. Syverson, 1995). Further, they provide a new basis for ice sheet modeling and palaeoclimatic reconstructions.

The deglaciation in southern Québec during Late Wisconsinan: from the St. Lawrence ice stream to the early phase of Younger Dryas The mode of deglaciation in the southern Québec part of the Laurentide Ice Sheet was controlled by a series of semiindependent climatic and non-climatic factors (Occhietti et al., 2001a). The global warming after the Last Glacial Maximum continued with some minor climatic coolings

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Fig. 2. Moraines and eskers in southern Québec.

from 17,000 y. B.P. to the end of the Allerød c. 11,000 y. B.P. (see GRIP curve, Johnsen et al., 1992; Dansgaard et al., 1993; Grootes et al., 1993). The generally negative mass balance conditions for the LIS resulted in the gradual transition from a multidome-shaped ice sheet, such as the Antarctic ice sheet (Lliboutry, 1965), to a much flatter plateau-like ice sheet, mostly in the marginal areas, muchlike modern-day Greenland ice sheet (Lliboutry, 1965). This thinning of the ice sheet had several consequences in its marginal areas including: 1) increasing topographic control on ice-flow patterns, 2) development of ice streams, as in the Greenland and Antarctic ice sheets, 3) migration of ice divides and ice mass centers and 4) several secondary glacio-dynamic readjustments. It must be mentioned that the chronology of the deglaciation episodes is based on unconventional 14 C ages from marine shells (d 18 O = 0 ‰). These ages are younger than the conventional ages by 400 years (d 13 C - 25%). Nevertheless, they are apparently older than ages established from terrestrial material and the New England

varve chronology by at least 350 years (cf. Pending problems, below).

The St. Lawrence Ice Stream Below a certain ice thickness and in conjunction with accelerated calving in the Gulf of St. Lawrence, a major north-east-trending ice stream formed within the ice sheet (Occhietti et al., 1996, 1997; Parent & Occhietti, 1999). Diachronously, between about 15,000 and 12,000 14 C years B.P., the head of the St. Lawrence Ice Stream migrated 1000 km along the axis of the St. Lawrence Corridor deep into the Laurentide Ice Sheet. This major feature of the LIS was characterised by flow rates that were at least one order of magnitude higher than in adjacent ice masses, similar to present flow rates in Greenland (Lliboutry, 1965). This accelerated ablation regime (ice stream and iceberg calving) caused substantial thinning within the catchment area of the ice stream, particularly on the north-west flank of the

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Figure 3. Outlines of the post-glacial marine invasions and lacustrine inundations in Québec-Labrador (from a compilation by V. Prest and A. Dyke).

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Fig. 4. Tentative model of the ice-front retreat of the Laurentide Ice Sheet, in Québec-Labrador, between 13 and 6.5 ka.

Appalachian Uplands (Shilts, 1981; Genes et al., 1981; Lowell, 1985) and along the southern margin of the Laurentian Highlands (Fournier, 1998). Instead of the expected concentric marginal ablation pattern in the south-

eastern margin of the LIS (concentric to the New Québec Dome), the St. Lawrence Ice Stream favored the progressive isolation of Appalachian ice masses in New Brunswick and northern Maine, in the Gaspé Peninsula and.

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Fig. 5: Estimated limits of the glaciers during the Saint-Narcisse Moraine episode, c. 10,850 y. B.P. (marine shells), during the main phase of Younger Dryas Chron.

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in the Notre Dame Mountains of southern Québec (Stea et al., 1998). Instability in the ice masses generated by the St Lawrence Ice Stream strengthened the role of regional topographic control on regional glacial dynamics and increased the sensitivity of ice margins to climatic fluctuations.

Deglaciation in the Appalachian sector: overview Mostly allochthonous ice masses over the Appalachians were progressively isolated from the New Québec Dome, at first as a result of an ice-flow reversal and then by a marine embayment which cut them off from the main ice sheet (Fig. 4). This two-fold sequence of events occurred in rapid succession between 14,000 and 12,000 yrs B.P., from the Gaspé Peninsula to the Bois -Francs region (Allard & Tremblay, 1981; Lebuis & David, 1977; David & Lebuis, 1985; Locat, 1977; Shilts, 1981; Genes et al., 1981; Lowell, 1985; Lowell et al., 1986 and 1990; Newman et al., 1985; Rappol, 1993; Martineau & Corbeil, 1983; Chauvin et al., 1985; Dionne, 1977). This separation, initially caused by the dynamics of the St. Lawrence Ice Stream, led to the formation of independent ice domes in the Gaspé Peninsula, in New Brunswick, in northern New England and adjacent Québec (Lowell et al., 1986; Stea et al., 1998). The latter dome became increasingly autonomous as a result of recession through coastal Maine and ice flow towards the St. Lawrence at its northern margin. The eastern part of this dome remained connected with the New Brunswick one (Rappol, 1989) while its western part was characterised by rapid ablation and by flow convergence toward the Hudson-Champlain Valley (Connally & Sirkin, 1973; Connally, 1982; Hughes et al., 1985). This late differentiation of the ice mass over the Appalachians was the result of topographic, glacio-dynamic and climatic factors. At the height of the Bølling warm interval (ca 12,700 to 12,000 conventional y. B.P.), the overall budget must have been strongly negative. The proposed mode of deglaciation in the Appalachians is that of a low mountain ice. The high ridges of the White Mountains and of the Notre Dame Mountains progressively separated ice masses on their north-western flank in Québec from those on their south-eastern flank in Maine (Shilts, 1981; Genes et al., 1981; Lowell, 1985). Four sectors can be recognised on the northern flank of the mountains (Fig. 2): 1) southwestern ice masses which were still connected with the main LIS (see Thompson et al., 1999 for deglaciation in New Hampshire, south of the Québec border), 2) the Bois Francs residual ice cap, which was characterised by widespread downwasting and stagnation (Parent & Occhietti, 1988; LaSalle & Chapde-laine, 1990), 3) the Lower St. Lawrence River area which was affected by late glacial flow toward the St. Lawrence and then by downwasting, and 4) the Gaspé Peninsula (Gaspésie) which acted like an isolated ice cap and downwasted lately (Richard et al., 1997).

The western arm of Goldthwait Sea: early deglaciation along the southern coast of the lower Estuary of St. Lawrence During deglaciation, thinning of the ice along the Appalachian piedmont favoured early deglaciation and incursion of the western arm of the Goldthwait Sea between the main ice sheet and the piedmont. This mode of glacial retreat does not require the development of a calving bay across the axis of the corridor; on the contrary, it implies calving parallel to the axis and to the shore. The major water gap of the Chaudière Valley adds further local complications in the overall deglaciation pattern. The upstream extent from the Gaspé Peninsula to the opposite side of the Saguenay mouth occurred before or by 13,000 y. B.P. At that time, the other areas of Québec were still glaciated (Fig. 4).

Dynamics of the Appalachian ice masses during the ice flow phase of Chignecto between 13,000 and 12,500 y. B.P.: deglaciation of the Baie des Chaleurs area In the Atlantic Provinces, the Ice Flow Phase 4 or Chignecto Phase records a glacio-dynamic event which occurred between 13,200 and 12,500 y. B.P. and is correlated to the Port Huron Event in the Great Lakes area (Stea et al., 1998). At that stage, the dome over the southeastern and central part of the Gulf of St. Lawrence separated into several interconnected, radially flowing ice masses. At the end of this phase, the Baie des Chaleurs was deglaciated, although ice remained over New Brunswick and the Gaspé Peninsula.

The major northward Appalachian ice flow reversal A late glacial reversal in ice flow is well documented by striations and rat-tails in the eastern part of the southern Appalachians of Québec (Lamarche, 1971, 1974; Lortie & Martineau, 1987; Rappol, 1993). This reversal was a powerful event as it caused a northward ice flow beyond the St. Lawrence River (Lanoie, 1995; Fournier, 1998; Paradis & Bolduc, 1999), and related glacial striations are observed almost as far south as the Québec/Maine boundary (Shilts, 1981). Shilts (1981) suggested an ice divide, the Quebec Ice Divide, which is the equivalent to the late phase of a moving North Maine Ice Divide defined later by Lowell & Kite (1986) and Lowell et al. (1986). The Appalachian reversal is considered as a major event and seems to have been a reequilibration event of the Appalachian ice masses. The ice over southern Québec and northern Maine flowed towards thinner ice in the middle and upper estuary. This event could be pre-Bølling in age. A tentative 13,800 – 12,900 y. B.P. (marine shells) age bracket is proposed and the event could be more or less coeval of the Anticosti Island Phase or the Chignecto Phase. Heavy snow falls on the Maine - New Hampshire-

Québec southern Québec ice mass might have amplified the strength of the reversal.

Reequilibration of the Maine-New Hampshire and southern Québec ice mass and strong ice stream in the middle estuary From striations, Bla is (1989) and Shilts (1997) identified north-eastward ice flow in the upper middle Chaudière Valley subsequent to northward ice flow. The change in ice flow direction is related to drawdown toward the St. Lawrence middle estuary documented by north-north-east to north-east striations in the Bois Francs, the Appalachian piedmont, an area north of the St. Lawrence River, and on islands and the southern coast of the middle estuary (Dionne, 1972). Inland of western and central Charlevoix, glacial striae record eastward ice flow convergent towards the estuary (Lanoie, 1995; Fournier, 1998). The ice flow patterns indicate unstable ice and rapid thinning of Appalachian ice over southern Québec. The following deglaciation events were influenced by ice thinning, emergence of nunataks in the Appalachians and differencial ice-surface lowering, as described by Syverson (1995) in historic deglaciation of Burroughs Glacier in Alaska.

Rapid disintegration of the southern Québec Appalachian ice masses during a part of the Bølling-Allerød warm episode At the start of the final deglaciation, Appalachian ice in southern Québec occupied three settings. In the west, it was connected to the Hudson-Lake Champlain lobe of the LIS. In the central part, it terminated at the Saint-Maurice lobe, and was fed by the LIS. The eastern part, from Bois Francs to Bas du Fleuve, was initially connected to the LIS, but thinned rapidly, ablating between 12,550 and 12,150 14 C y B.P.(marine shell ages) (Parent & Occhietti,1999) during a part of the Bølling-Allerød warm episode (cf. Pending questions, for a reappraisal of these ages). Recessional moraines comprising the Frontier, Dixville-Ditchfield, Sutton-Cherry River-East-Angus-Megantic, Mont Ham and Saint-Ludger, and Ulverton-Tingwick morainic belts in the southern Appalachians of Québec (Fig. 2) were described by MacDonald (1966, 1967, 1968, 1969), Shilts (1970, 1981), Prichonnet (1984) and Parent (1987) (see Parent & Occhietti,1999). These ice-front features (Fig. 2) are transverse to the structural trend of the Appalachians. The ice margin outlined by the moraines is lobate and topographically-influenced, with lobes in valleys and reentrants on adjacent uplands and ice surface slopes gentle. The morainic belts do not occur in north-western Maine, where ice was probably stagnant. Local ice readvances are identified on the Appalachian in the Bas du Fleuve (Rappol, 1993), in the Rimouski area (Rappol, 1993; Hétu,1998) where the Neigette Readvance is dated at 12,400 y. B.P. (marine shell age). Later, a local

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readvance or reactivation of the Laurentide ice occurred in the Chaudière Valley, named the Beauce Event (Occhietti et al., 2001a) and formerly related to the Highland Front Moraine System (Gadd et al., 1972a, b). The local recessional Saint-Sylvestre Moraine constructed later is tentatively related to the Ulverton-Tingwick morainic belt. The age of the last morainic belt on the south-western Appalachian piedmont, the Ulverton-Tingwick Moraine, is established from the number of glacial varves at the Rivière Landry section (Parent, 1987; Parent & Occhietti, 1999). The ice front receded from the moraine more than 120 years prior to the opening of the St. Lawrence Valley to marine waters, c. 12,000 y. B.P. (marine shells). Glacial Lake Candona (Parent & Occhietti, 1988; or Lake St. Lawrence of Rodrigues, 1992), which resulted from the coalescence of glacial lakes Memphremagog, Vermont and Iroquois, inundated the deglaciated lowlands of the St. Lawrence Valley and valleys in the Appalachians until 12,000 y. B.P. (marine shells) (Parent & Occhietti, 1988).

The ultimate deglaciation along the Appalachain piedmont and Champlain Sea transgression Along the north-eastern Appalachian piedmont, the ultimate episode of Laurentide ice retreat is the Saint-Raphael Moraine (Fig. 2). After the retreat from this position, the western arm of Goldthwait Sea extended rapidly between the Appalachian piedmont and the retreating LIS margin upstream up to the Bois Francs piedmont where the SaintMaurice lobe front still abutted against the piedmont. Some residual ice may have persisted on the Bois Francs plateau and in the northern side of the middle estuary. Then, the opening of the St. Lawrence Valley to the marine waters occurred by thinning of the LIS ice front along the piedmont and by overflow of the lacustrine waters of Lake Candona through a spillway inset into the ice and parallel to the piedmont. As recorded in the Rivière Landry section (Parent, 1988) and in the western basin of Champlain Sea (Rodrigues, 1992), an early phase of mixing fresh and marine waters is recorded at the base of the Champlain Sea sediments. The age of the marine incursion is close to 12,000 y. B.P. (marine shells) (c f. Pending questions). Deglaciation in the St. Lawrence Valley and middle Estuary The style of deglaciation in the St. Lawrence Valley is poorly known because many deglacial features were either buried or reworked as the result of marine inundation. The northward pattern of glacial retreat implies that glacial isostatic rebound began earlier along the Appalachian piedmont than along the southern edge of the Laurentians. Glacial thinning generated by the St. Lawrence Ice Stream also caused early crustal unloading along the axis of the St. Lawrence Corridor and a delayed response from east to west (Dionne, 1977; Locat, 1977; Lebuis & David, 1977; Parent, 1987; Dionne, 1988; Dionne & Coll, 1995).

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Deglaciation along the southern margin of the Laurentian Highlands The southern margin of the Laurentian Highlands includes the Mont Tremblant Highlands, the Saint-Maurice Valley, the Parc des Laurentides Highlands and the Saguenay fjord. This physiographic context favoured faster ice-marginal retreat in areas down-glacier (south) of the highlands while arcuate lobes were maintained ni the two main valleys (Occhietti, 1980; Govare, 1995). The St. Lawrence Ice Stream interfered with this topographic context because flow convergence toward the estuary from the eastern flank of the Parc des Laurentides Highlands (Charlevoix) resulted in accelerated thinning. On the north shore of the middle and upper estuary, ice lobes persisted in major valleys and readvanced later (Dionne & Occhietti, 1996; Govare, 1995; Occhietti, 1980), modifiying significantly the mode of deglaciation of the St. Lawrence Valley. The Younger Dryas Episode The Younger Dryas is a worldwide cold phase which lasted from about 10,900 to 10,100 y. B.P. In Québec, the SaintNarcisse Moraine has traditionally been correlated with this episode (LaSalle & Elson, 1975; Occhietti, 1980). More recently, in the Québec City area, fossiliferous diamictons, with Balanus hameri fragments, have also been correlated to a Younger Dryas glacial readvance (the Saint-Nicolas readvance of LaSalle & Shilts, 1996). North of the SaintNarcisse Moraine, a new outline of ice front features, named the Batiscan moraine (Bolduc, 195), was compared to one of the Salpausselkä moraines of Finland. This ice front feature was previously noted, close to the Mars River, in the Charlevoix area (Govare, 1995). Recent research (F. Robert, 2001) shows that several outlines of glacial features, younger than the Saint-Narcisse Moraine, can be followed as far 100 km west of the Batiscan area, at least until the Saint-Maurice River. The first author thinks that the Saint-Nicolas readvance is an ice flow in the Champlain Sea. The climatic significance of this local episode is not demonstrated. It could record a first minor cold pulse of Younger Dryas age, after c. 11,000 y. B.P. The SaintNarcisse Moraine episode, between about 10,800 and 10,500 y. B.P., is related to the first major cold episode of the Younger Dryas. The younger morainic systems, such as the Mars -Batiscan-Saint-Maurice and unnamed intermediate moraines result probably from the late cold episodes of Younger Dryas, between 10,500 and 10,100 y. B.P. The Saint-Narcisse Moraine is an almost continuous feature, extending more than 750 km from the Ottawa River area (Occhietti, unpublished) to the Saguenay fjord (Dionne & Occhietti, 1996) (Figs 1, 2 and 5). This length is equivalent to the distance between Bergen, on the Atlantic coast of Norway, and Stockholm, on the Baltic coast of Sweden. Some undated moraines are correlated with the Saint-Narcisse Moraine: the outer Baie-Trinité Moraine on the north shore of the St. Lawrence Estuary (Vincent,

1989), the Bradore Moraine close to the Québec-Labrador limit, and the Belles Amours Moraine in Labrador. The last three moraines are not a part of the following discussion. The outline of the Saint-Narcisse Moraine can be subdivided in several lobes (Gatineau, Saint-Maurice, central Charlevoix, Saguenay) and reentrants (Mont tremblant, Parc des Laurentides). A topographic influence is implied from this general setting: the gentle slope of the frontal zone of the LIS favoured the ice convergence toward the relative depressions of the Laurentian Highlands and locally to the post-glacial sea. This general outline is related both to the climatic cause and to the topographic context. The western limit, at a lower latitude than the eastern limit (45°45' N as compared to 48°10' N), and at a greater distance from the New Québec dome, is characterised by discontinuous features with dominant meltwater deposits. In the Saint-Maurice River area, the moraine consists of large bodies of glacial, glaciomarine and proglacial deposits. In the eastern area, in Charlevoix, the moraine is in fact a series of thirty small concentric frontal ridges (Rondot, 1974) which indicate a more active ice than in the western part. These local ice dynamics are partly related to the proximity of the ice centre and mainly to a steeper slope of the LIS margin which is the consequence of the fast ablation generated earlier by the St. Lawrence Ice Stream. The Saint-Narcisse Moraine indicates a global readvance or at least a stabilisation of the LIS and that the flow lines of the LIS margin were maintained for several centuries (Fig. 5). The late Younger Dryas episode is indicated by the Mars-Batiscan-Saint-Maurice moraine and other parallel ice front features. These features emphasise the different styles of ice retreat, with dominant meltwater features (eskers, ice margin trains) in the western part and ice front rims in the eastern part.

Deglaciation of the Canadian Shield area (Québec And Labrador) during the Early Holocene The deglaciation of the Canadian Shield in Québec and Labrador is described by Vincent (1989). The general deglaciation style reflects thermolatitudinal retreat with topographic control, and shifting of ice centers and divides (Figs 4, 5 and 6). Glacial flowline reconstructions indicate large, complex and evolving ice masses located over Québec-Labrador, referred to as the New Québec Dome in this text, and as the Hudson Bay ice which became dynamically independent. In addition to ablation by surface melting, ice loss to the Labradorean Sector was the result of mechanical, physical processes along its margins, including: 1) an ice stream in the Hudson Strait which induced convergence of ice flow towards Ungava Bay and Hudson Strait, 2) accelerated ice flow in the south, towards the Great Lakes, and to the west, towards an extended Lake Agassiz

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Figure 6: Estimated limits of the glaciers (New Québec Dome) during the Sakami Moraine episode, c. 8,000 y. B.P. (marine shells), prior to the opening of Hudson Bay to Tyrrell Sea.

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Mountains) which were progressively deglaciated during late Wisconsinan and Early Holocene. In southern Labrador, a concentric retreat pattern predominated, due to the termination of the ice sheet margin on land, unaffected by ice streaming and calving. All these factors, including a negative budget of climatic origin, reequilibration of the ice masses (rapid ice flow and thinning in response to the physical loss of ice at the margins), accelerated deglaciation of the Labrador Highlands and the development of ice streams explain the shape and dynamics of the Labradorean Sector at the eve of the Holocene and the subsequent mode of ice sheet decay. By 10,000 y. B.P., the LIS was still a major ice mass of about 6.106 km2 (Occhietti, 1983); it took at least a further 3500 years to dissipate completely. Taking into account the latent heat of ice (80 cal./g to transform ice at 0°C to water), the LIS probably decreased the rate of global climatic warming during Early Holocene. The loss of heat energy due to phase change was reinforced by the albedo (reflective heat loss from the ice surface). The LIS was in disequilibrium with the interglacial insolation conditions, and generated a major effect on the world’s climatic system.

Ice retreat in the southern margin of the New Québec Dome: the Laurentians North of the Younger Dryas moraines (Saint-Narcisse and the Mars-Batiscan-Saint-Maurice morainic systems), the marginal retreat history of the ice sheet is well known in the south-western part (the Lac Témiscamingue-Ottawa River area), and poorly known in its central part, between the Ottawa and Manicouagan Rivers. Along the northern shore of the St. Lawrence lower Estuary (the middle and lower Côte Nord) and in southern Labrador, controversy remains on the age of regionally extensive morainic systems (North Shore, Bradore, Belles Amours, Paradise and Little Drunken moraines). Although they outline marginal positions of Labrador-Québec ice, their age and correlation are unknown.

Deglaciation in the western part of the Laurentians and Lac Témiscamingue area West of Ottawa River, the southern margin of the Hudson Bay ice mass over the James Bay lowlands and northern Ontario and Manitoba, was dissipating rapidly, with successive readvances or surges in Lake Superior (Marquette readvance, circa 10,000 y. B.P., Lowell et al., 1999), in Lake Agassiz (several secondary lobes and the Marquette readvance, Dredge & Cowan, 1989, c. 10 ka), and by the Cochrane non climatic ice surges in Lake Ojibway, circa 8400 y. B.P. In the upper Ottawa River and Lac Témis camingue area, uncoupling along the suture between the Hudson Bay ice and the New Québec ice is evidenced by

converging ice flow indicators (Veillette, 1986), by the diachronous Harricana interlobate Moraine (Vincent & Hardy, 1979), and by a continuous glacial lake inundation which followed the ice front retreat and partly resulted from the glacioisostatic lowering at the front of the ice sheet. During deglaciation, southwest flowing New Québec ice and southeast flowing Hudson ice converged towards the Harricana Interlobate Moraine (Hardy (1976), which was originally described by Wilson (1938) as "a moraine between two ice-sheets". This origin and interpretation of this moraine is disputed (Brennand & Shaw, 1996; and Brennand et al., 1996). According to these authors, the Harricana features are a mega-esker, deposited prior to the deglaciation and during a short episode. From the west southwest-east northeast orientation of ice front during the deglaciation on the Laurentian Plateau in Québec (Simard, 2003), on the eastern side of the Ottawa River, and the lack of visible interlobate features between the Lake McConnell Moraine and the southern limit of the Harricana Moraine, it seems that the two pre-existing models do not apply in the southern Temiscaming area. The ‘Harricana Moraine’ is probably a composite feature, interlobate converging deposits and tunnelled-meltwater deposits. The southwestern margin of the New Québec Dome retreated initially in a north-northeasterly direction (Veillette, 1997a; Simard, 2003). The southern region on the west side of the Harricana Moraine became ice free first (Veillette, 1983b, 1986, 1988). The Laverlochère Moraine (Veillette, 1983a,b, 1986, 1988) was built on the borders of an ice lobe in the northern Lac Témiscamingue trough. West of the Harricana Interlobate Moraine and north of Lac Témiscamingue, the Roulier Moraine (Vincent & Hardy, 1977, 1979) was formed as a result of a halt or as a reequilibration moraine. Glacial lakes abutted the ice margin in low-lying western and northern areas (Vincent & Hardy, 1977, 1979 Veillette, 1983b, 1988). The earliest glacial lake phase is related to the north-eastern extension of the PostAlgonquin glacial lake from the Great Lakes basin (Harrison, 1972) (Fig. 3). Possibly during the Sheguiandah and certainly during the Korah Phase, the lake was still dammed by ice blocking drainage down Ottawa River in the Mattawa area, and extended later north-east of Témiscaming along a reentrant in the ice front along the Harricana Interlobate Moraine (Vincent & Hardy, 1977, 1979; Veillette, 1988; Figs 1 and 4). When ice withdrew from the Ottawa River valley, water levels dropped, and glacial Lake Barlow (Wilson, 1918) occupied the Lac Témiscamingue basin. Maximum glacial Lake Barlow levels in the area east of Lac Témiscamingue were at about 300 m and rise to the northeast to about 380 m in the vicinity of the present Hudson Bay watershed where the Harricana Moraine crosses it (Veillette, 1988). The north-eastern tilt of glacial Lake Barlow water planes confirms that ice in New Québec was thicker than in Hudson Bay (Hillaire -Marcel et al., 1980). The chronology of deglaciation in the Matawa and Temiskaming areas has been recently reassessed. Anderson

Québec et al. (2001) resampled basal lake deposits. Terrestrial plant debris yielded an AMS age of 9450 ± 50 y. B.P. (CAMS46195) which shows that original dates from total organic matter sampled at the same depth are too old by 2000 years (hardwater effect). This result is confirmed by the recent reassessment of the position of the western extent of the St-Narcisse Moraine in a more southerly location that pre-viously interpreted (Robert, 2001; Occhietti in Bhiry et al., 2001). The nearly west-east ice-front outline of the LIS on the Laurentian Plateau between about 10,000 – 9700 y. B.P. is another field evidence of a late deglaciation in the area. Laurentian Highlands between Lac Témiscamingue and the Manicouagan River The characteristic deglacial deposits of this poorly-studied area are short segments of end moraines, isolated icecontact deposits, eskers and outwash trains (Parry, 1963; Hardy, 1970; Denis, 1974; Lamothe, 1977; Pagé, 1977; Tremblay, 1977; Occhietti, 1980). In this moderate relief area higher summits first became ice free. The pattern of deglaciation is mostly the result of thinning of a wide marginal zone of the ice sheet. In the western area, the ice front outline changes from a nearly west-east outline observed by Simard (2003) to a curved outline related to the southwest limit of the New Québec Dome. From west to east , the rate of ice retreat decreases, the more distant ice margin from the ice center is thinner and melts faster. In the Lac Saint-Jean area,, Tremblay (1971), Dionne (1973), and LaSalle & Tremblay (1978) have shown that the ice moved in a southerly direction except along the Saguenay fjord, where ice flow was generally south-easterly along the axis of this depression (Dionne & Occhietti, 1996). As the ice retreated from the uplands south of Lac Saint-Jean, small glacial lakes were dammed and De Geer moraines were built. A late ice lobe occupied the Lac Saint-Jean depression and ice-contact materials were deposited at its receding margin, among them is the Metabetchouane moraine (LaSalle & Tremblay, 1978). As the ice receded north-westerly up the Saguenay Valley and into the Lac Saint-Jean basin, marine waters from the Gulf of St. Lawrence extended over lower-lying deglaciated land. This arm of the Goldthwait Sea was referred to as the Laflamme Sea by Laverdière and Mailloux (1956). Marine limit lies at 167 m at the mouth of Rivière Saguenay and at about 167 m and 198 m south and north of Lac Saint-Jean, respectively (LaSalle & Tremblay, 1978). The oldest ages (marine shells) for marine invasion and deglaciation are 10 400 ± 150 BP (1-5922) at the mouth of Rivière Saguenay and 10 250 ± 350 BP (Gif-424) south of Lac Saint-Jean. They seem too old by several centuries. Following deglaciation of the Lac Saint-Jean basin, the ice retreated northward leaving behind numerous eskers and fluted landforms.

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Québec North Shore The Quaternary history of the coastal fringe north of the Gulf of St. Lawrence between the mouth of Saguenay River and the Québec/Labrador border is relatively well known but few data are available for farther inland. During the LGM, south flowing ice of the New Québec Dome covered the Québec North Shore and extended in the Gulf of St. Lawrence (cf. Grant, 1989), without apparently overrunning the eastern tip of Anticosti Island (Gratton et al., 1984). Ice retreated generally north-westwards during deglaciation. The eastern extremity of the North Shore and the headland in the Baie-Trinité region were probably the first areas to become ice free. Moraines were built in the BaieTrinité area between 13.5 and 9 ka (Dredge, 1976b, 1983b). The outer Baie-Trinité Moraine is correlated with the SaintNarcisse Moraine. The marine submergence (Fig. 3) in the Estuary and Gulf of St. Lawrence, east of Québec City, was named Goldthwait Sea by Elson (1969) and studied by Dionne (1977), Hillaire-Marcel (1979), and Dubois (1980). Marine limit varies from 150 m in the Québec/Labrador border area (Boutray & Hillaire -Marcel, 1977) to 130-145 m along the North Shore (Dubois et al., 1984) and near the mouth of Rivière Saguenay (Dionne & Occhietti, 1996). The age for deglaciation of the Québec North Shore comprised between 10,900 ± 140 BP (GSC- 2825) (site in Labrador near the Québec border) and (10,230 ± 180 BP, Gif-3770; Rivière Romaine), 9,140 ± 200 BP.(GSC-1337, Rivière Moisie), and 9,970 ± 130 BP (QU-574 (Rivière Manicouagan). The upstream part of the North Shore was apparently deglaciated as early as 12,470 ± 80 y. B.P. (Beta-170455, from marine shells, Les Escoumins), and reglaciated during Younger Dryas. As ice retreated farther north onto the Canadian Shield, the Québec North Shore Moraine (Dubois & Dionne, 1985), more than 800 km long, was emplaced between Rivière Manicouagan and south of Lake Melville in Labrador (Dubois, 1979, 1980; Fulton & Hodgson, 1979). The moraine includes segments in the Rivière Manicouagan area (Sauvé & LaSalle, 1968), the Lac Daigle Moraine of Dredge (1976b, 1983b), the Manitou-Matamec Moraine of Dubois (1976, 1977, 1979, 1980), the Little Drunken Moraine of Fulton & Hodgson (1979), and the Aguanus-Kenamiou Moraine of Dionne & Dubois (1980). The moraine is considered by Dubois & Dionne (1985) to represent a halt of the ice sheet during a cooler climatic phase. The age of the moraine, still uncertain, is apparently of 9.5 to 9.7 ka (Dubois & Dionne,1985). The Goldthwait Sea was in contact with the moraine locally in the Rivière Moisie and Rivière Manicouagan areas. The ice margin retreated northward and north-westward from the Québec North Shore Moraine towards central New Québec leaving eskers and fluted landforms. According to Dubois (1980), the Goldthwait Sea in the middle North

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Shore area attained its inland limit (128-131 m) at about 9.5 ka. Deglaciation of the eastern Labradorean Sector: central and southern Labrador Labradorean Sector ice flow was outward, towards the central and southern Labrador coasts. During deglaciation it was generally north-eastwards in central Labrador, eastward in the Lake Melville area and south-eastwards in south-eastern Labrador. Flow was topographically controlled by the Mealy Mountains, around which flow trends part, and by Lake Melville, where a calving bay may have existed (Fulton & Hodgson, 1979). Major end moraines were built during the retreat phase. If Laurentide Ice covered all of south-eastern Labrador and joined with the Newfoundland Ice Cap, then the Bradore, Belles Amours and Paradise moraines (D.R. Grant, in Vincent, 1989, King, 1985) are probably Late Wisconsinan retreat features. Many other moraines were built during Holocene. The longest of these is the Sebaskachu-Little Drunken Moraine System (Blake, 1956; Fulton & Hodgson, 1979), which could be the extension of the Québec-North Shore System of Dubois & Dionne (1985). Later retreat towards central New Québec was marked largely by construction of numerous eskers and fluted landforms. Glaciofluvial landforms generally trend in the last principal directions of ice flow defined by glacially streamlined landforms. There are, however, notable exceptions that occur principally in central and western Labrador. In the central Smallwood Reservoir, a large esker system with prominent tributaries entering from the northwest crosscuts ice flow trends, suggesting major change in subglacial hydrology during the late deglaciation. The change may have been in response to evolution in the glacial lakes of northern Labrador or to subglacial discharge toward Lake Melville. Although free drainage to the Atlantic prevented major glacial lake development in western Labrador, glacial lakes of uncertain extent were dammed against the watershed divide; their full extent and evolution are not well known. Atlantic Ocean waters submerged glacio-isostatically depressed coastal areas of central and southern Labrador during deglaciation. Marine limit has been traced by R.J. Fulton (Geological Survey of Canada) (Figs. 4 and 5). Along the coast, north of Straits of Belle Isle it may have reached 150 m, between Sandwich Bay and Lake Melville 113 m (Rogerson, 1977), and north-east of Lake Melville 85 m (Hodgson & Fulton, 1972). In the Lake Melville area marine limit increases in elevation inland from about 75 m on the outer coast to 150 m west of the lake (Fitzhugh, 1973). The maximum age and elevation of marine inundation decreases northwards along the Labrador Coast, and the 0 m hinge line is located at the northernmost tip of the Labrador Peninsula (Løken,1962a, b). On the basis of radiocarbon ages of shells, the area north of the Straits of Belle Isle was ice free by at least

10,900 ± 140 BP (GSC2825), the coastal area east of Kanairiktok River by 10,275 ± 225 BP (GX-6345), central Lake Melville by 7,970 ± 90 BP (TO-200), and the west end of Lake Melville by 7,600 +- 100 BP (GSC-2970). The age determinations of 9,640 ± 170 BP (GSC-3067), 10,550 ± 290 BP (S1-3139) and 10,240 ± 1,240 BP (S1-1737) on lake sediments provide minimum ages for the deglaciation of the upper St. Paul River area, the southern Mealy Mountains and the eastern end of Lake Melville, respectively. They also provide a minimum age for the construction of the Paradise Moraine. Farther inland a date of 6,460 ± 200 BP (GSC1592) on peat provides a minimum age for the deglaciation of upper Churchill River area and a 6,500 ± 100 BP (GSC3241) date for lake sediments, a minimum age for the deglaciation of central Labrador. At Hopedale and Nain, on the central Labrador coast, glacier ice remained much later, until c. 7600 ± 200 and 8500 ± 200 yr ago, respectively (Clark & Fitzhugh, 1990, Awadallah & Batterson, 1990).

Deglaciation on the western side of the New Québec Dome Québec Clay Belt and James Bay area East of the Harricana Interlobate Moraine ice flow was first towards the south-south-west and swung progressively to the west (Hardy, 1976). In the La Grande Rivière area, superposed drumlins and striae (Lee et al., 1960) show that New Québec Ice was free to move in a more westerly direction following separation of the ice masses (Vincent, 1977). West of the Harricana Interlobate Moraine, Veillette (1986) has shown Late-Glacial south-easterly flow. As New Québec and Hudson ices retreated, glacial Lake Ojibway was dammed between the ice front and the drainage divide to the south (Coleman, 1909; Vincent & Hardy, 1977, 1979). The maximum lake limit rises northward from about 380 m to more than 450 m. The maximum depth of Lake Ojibway was more than 500 m on the east coast of James Bay. Lake waters flooded the lowlands east of James Bay as far as the Sakami Moraine (Hardy, 1976) and as far north as Kuujjuarapik (Hilla ireMarcel, 1976). Glacial Lake Obijway became separated from glacial Lake Barlow with the emergence of a sill (Vincent & Hardy, 1977, 1979). This sill was then the lowest point on the drainage divide, which was displaced south by isostatic tilting of the crust. Large channels on the divide and wide and deeply-incised Rivière Kinojévis valley were cut by overflowing glacial lake waters. Extensive fields of De Geer moraines were built in glacial Lake Ojibway. The varve chronology established for lakes Barlow and Ojibway (Antevs, 1925; Hughes, 1965; Hardy, 1976) indicates 2110 years of lacustrine sedimentation to the time Lake Ojibway drained into Tyrrell Sea. Using this, Hardy (1976) calculated the rate of ice retreat which increased from about 320 m/yr south-east of James Bay to 900 m/yr in the La Grande Rivière area.

Québec Late glacial ice surges occurred into glacial Lake Objiway. Hardy (1976) has indicated three Cochrane surges in the lowlands south-east of James Bay which came from Hudson ice. The maximum Cochrane I and II surges occurred 300 years and 75 years before glacial Lake Ojibway drained. These surges presumably formed at the southern terminus of a short-lived ice stream that formed in the central James Bay region (Parent et al., 1995; Veillette, 1997). When New Québec Ice had retreated to the approximate position of the Sakami Moraine, marine waters from Hudson Strait penetrated Hudson and James bays and flooded the isostatically -depressed lowlands. Opening to the sea led to sudden drainage of glacial Lake Ojibway and to formation of the Sakami Moraine (Hardy, 1976), a major feature extending inland over a distance of 630 km from the coast of Hudson Bay to southern Lac Mistassini (Figs. 3 and 6). Hillaire -Marcel et al. (1981) considered the Sakami Moraine as a re-equilibration moraine which results from the stabilisation of the ice front when the glacier grounded after drainage of Lake Ojibway. This drainage led also to a slight readvance in the Lac Mistassini area (Bouchard, 1980), previously recognised as the Waconichi ice advance by DiLabio (1981). The sudden drainage of glacial Lake Ojibway and the submergence by the Tyrrell Sea are recorded on hills in the Lake Ojibway basin by upper and lower wave washing limits (Norman, 1939; Hardy, 1976). The Tyrrell Sea followed the retreating ice front after construction of the Sakami Moraine. Marine limit is at about 198 m in the southern part of its basin (Hardy, 1976) and rises northwards up to 315 m in the Kuujjuarapik area (Hillaire -Marcel, 1976). In La Grande Rivière area, marine limit decreases eastwards (inland) from 270 m in the area of Sakami Moraine to 246 m farther up river (Vincent, 1977). Extensive swarms of De Geer moraines, many of which overlie drumlins and eskers, were built east of the Sakami Moraine. Rates of ice retreat averaged 217 m/a (Vincent, 1977). In the Lac Mistassini area, the relatively shallow glacial Lake Mattawaskin (Bouchard,1986) followed the retreating ice front (Fig. 3). Rates of ice recession in this lake basin were estimated at 220 to 260 m/yr (Bouchard, 1980). An approximate age of c. 8.0 ka from marine shells (Hardy, 1976) and from concretions (Hillaire -Marcel, 1976) is assigned to the Sakami Moraine and the Tyrrell Sea invasion. This age corresponds closely with the age proposed for the drainage of glacial lakes Agassiz and Ojibway and for the incursion of Tyrrell Sea on the west side of Hudson Bay (Dyke & Dredge, 1989; Dredge & Cowan, 1989). Using this age and the varve chronology, the area west of Lac Témiscamingue was deglaciated 10,000 B.P. and the height of land, at the Québec-Ontario border, 9.2 ka ago; the Cochrane I reached its maximum 8.3 ka ago and Cochrane II 8 ka ago. East of the Sakami Moraine, based on the De Geer moraine chronology (Vincent, 1977), the Tyrrell Sea reached its eastern limit in the La Grande Rivière area about 7.5 ka ago. Farther inland the oldest

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radiocarbon age of 6,600 ± 100 BP (B-9516) was obtained on basal gyttja (Richard, 1995). Area east of Hudson Bay Ice from central New Québec and central Ungava Peninsula (Nunavik) flowed generally westwards and southwestwards into Hudson Bay (Hillaire-Marcel, 1979; Gray & Lauriol, 1985; Bouchard & Marcotte, 1986; Parent et al., 1997; Veillette, 1997; Veillette & Roy, 1995) as far as the Belcher Islands (Jackson, 1960). When Hudson ice finally separated from New Québec ice in Hudson Bay, only the area west of the Sakami Moraine in the Kuujjuarapik area, and perhaps the extreme north-western part of Ungava Peninsula, was ice free (Fig. 3). Location of the ice-stillstand features in the Nastapoca Hills at ages ranging between 7850 and 7610 BP (Lajeunesse, 2000) suggests that the position of the New Québec ice-front was much further to the east from what was reported in previous models. By 7700 B.P. the ice front extended in an arc-like shape to Manitounouk Strait and Petite-Rivière -de-la -Baleine, where contemporaneous submarine ice-contact sediments and landforms are identified (Zevenhuizen, 1996). From there, the iceterminus slightly curved northwards to the Nastapoka Hills near the Nastapoka River mouth, where moraines and emerged ice-contact submarine fans occur (Lajeunesse & Allard, 2003). North of the Nastapoka River area, the position of the ice front is unknown. The 7430 ± 180 age from Ottawa Islands (GSC-706; Andrews and Falconer, 1969) was provided from shells found on the surface of deltaic deposits, 17 m below marine limit, and indicates a rather late deglaciation of western Ungava. The ice-front followed probably the actual coastal area of Hudson Bay up to Inukjuak, and from there, continued in the shallow water zone (< 80 m) that extends from Inukjuak to the Ottawa Islands. North of Ottawa Islands, in the northern section of eastern Hudson Bay, the ice-front curved north-eastwards to Cape Smith where oldest age of 8040 ± 110 BP; UQ761) (Gray & Lauriol, 1985, redated at c. 6800 B.P., Lauriol & Gray, 1987).) was obtained. From the Cap Smith area, the ice-front continued north-westwards to Mansell Island and then curved to the Ivujik area. As the ice margin retreated to the east, the Tyrrell Sea covered the newly deglaciated areas and swarms of De Geer moraines were built at the ice front. The marine limit decreases northwards from about 315 m to possibly as low as 105 m east of Povungnituk. From there it rises northerly to about 170 m near Hudson Strait (Gray & Lauriol, 1985). Marine limit also declines inland from 248 m to 196 m along Rivière Nastapoca (Allard & Seguin, 1985), and from 158 m on the Ottawa Islands (Andrews & Falconer, 1969) to 105 m east of Povungnituk (Gray & Lauriol, 1985). Rates of uplift, as measured in the Lac Guillaume -Delisle (Richmond Gulf) area, were 9.6-10 m/century at the time of deglaciation (Hillaire-Marcel, 1976; Allard & Seguin, 1985). According to Hillaire -Marcel (1976), by 7000 B.P.

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the rate of uplift was 6.5 m/century and it decreased linearly to the present rate of 1.1 m/century. East of marine limit, ice continued its retreat towards central Ungava Peninsula in the north or towards central New Québec farther south. Beyond the Tyrrell Sea limit, a shallow glacial lake was formed in the Lac A I'Eau Claire area between an uplifted sill and the ice front (Allard & Seguin, 1985). At the head of Rivière aux Mélèzes, glacial Lake Minto formed between the Hudson Bay-Ungava Bay drainage divide and the receding ice front (Lauriol, 1982; Lauriol & Gray, 1983). Deglaciation in the northern areas of Québec-Labrador Ungava Peninsula (Nunavik) Both northerly flow into southern Ungava Bay, from central New Québec, and flow towards Hudson Bay, western Hudson Strait, and western Ungava Bay, from a central north-south ice divide on Ungava Peninsula are recorded (Gray & Lauriol, 1985; Bouchard & Marcotte, 1986). Ice flowing westwards and northwards from this ice flow centre, called the Payne center by Bouchard & Marcotte (1986), apparently coalesced with ice moving northeastwards in Hudson Bay and eastward in Hudson Strait (flow in offshore area from Andrews & Falconer, 1969; Shilts, 1980; Laymon, 1984; Gray & Lauriol, 1985). In Hudson Strait, major submarine moraines lie subparallel to the north-eastern part of the Ungava Coast (MacLean et al., 1992). Although their age is not known, they appear to define the northern limit of late glacial ice on the Wakeham Bay-Baie Héricart region c. 8,500 BP, or a little earlier, stratigraphically overlying and deforming stratified glaciomarine sequences that floor the central part of the Strait (Maclean et al., 1992). The north-western and northern extremities of the Ungava Peninsula were deglaciated first (Fig. 4). Marine waters submerged the Ungava coast along Hudson Strait where marine submergence generally decreases from west to east (167 m at Cape Wolstenholme, Matthews, 1967; 138 m near Diana Bay, Gray et al., 1980), and from north to south (170 m on Charles Island and 120 m at the head of Deception Bay, Gray & Lauriol, 1985). Rates of uplift at about 8000 B.P. were estimated at 7.9 m/century (Matthews, 1967). The timing of deglaciation of the southern Hudson Strait shore is subject of controversy. Most researchers agree that some areas were ice free at least 7,970 ± 250 BP (GSC- 672), but some, on the basis of three ‘older’ radiocarbon age determinations, postulate earlier deglaciation. A 10,450 ± 250 BP (1-488; Matthews, 1966, 1967) age may in fact be erroneous since shells collected by B. Lauriol (University of Ottawa) from the immediate vicinity of the Deception Bay site, originally sampled by Matthews, were dated at 7,130 ± 100 BP (GSC3947). Notwithstanding this, two other age determinations, the oldest of which is 9,800 ± 220 BP (Beta-11121), have been obtained from the Deception Bay area by dating in

situ Portlandia arctica and Nuculana minuta shells collected in glacial marine sediments overlying till (Gray & Lauriol, 1985; Lauriol & Gray, 1987). This conflicts with evidence on Meta Incognita Peninsula of Baffin Island which requires that ice extended to the mouth of Hudson Strait until 8.6 ka or later (Andrews, 1989). Miller et al. (1988) have suggested that a late readvance of ice from Labrador across Hudson Strait, between 9 ka and 8.2 ka, could account for the late presence of ice on southern Baffin Island. Upon retreat of the ice front towards the interior of Ungava Peninsula, glacial lakes were dammed between the ice front and higher ground on the Hudson Bay-Hudson Strait and Hudson Bay-Ungava Bay drainage divides (Prest et al., 1968; Prest, 1970). Of these lakes, the best documented is glacial Lake Nantais (Lauriol & Gray, 1987). Standing water bodies extending far inland in the lower parts of the valleys of Rivière aux Mé1èzes (Gray & Lauriol, 1985), Rivière Caniapiscau (Drummond, 1965) and of the George, Shepherd and Koroc Rivers require significant Late-Glacial ice remaining in Ungava Bay when interior of Québec-Labrador was largely ice–free. Ice in Ungava Bay retreated westwards towards the central Ungava Peninsula and southwards towards central New Québec. Eastern Hudson Strait was deglaciated by 9.1 ka on the basis of a radiocarbon age on shells collected from a seabed core (9120 ± 480 BP, GSC-2946). The D'Iberville Sea (Laverdière & Bernard, 1969) followed the retreating ice front. On the west coast of Ungava Bay marine limit rises from 138 m near Diana Bay to 195 m in upper Rivière aux Mé1èzes drainage basin (Gray & Laurio l, 1985). On Akpatok Island, 75 km offshore, marine limit is much lower (58-74 m; Løken, 1978). According to radiocarbon ages by Allard et al. (1989) on the south-eastern coast of the Ungava Bay and other ages on the western coast, all Ungava Bay coast became ice free more or less simultaneously, by about 7 ka. North-eastern Québec and northern Labrador During the LGM, Labradorean Sector ice flow was east north-easterly from central New Québec and easterly from Ungava Bay (Ives, 1957, 1958; Løken, 1962a) towards the Labrador Shelf over north-eastern New Québec and northern Labrador (Saglek Glaciation in the Torngat Mountains; Andrews, 1963a). North of Fraser River, coastal summits and much of the northern Torngat Mountains and adjacent coastal forelands were not overtopped by ice (Ives, 1978). Clark (1984) illustrated that ice from west of the Labrador watershed crossed the Torngat Mountains as outlet glaciers extending to fiord mouths, leaving large nunataks. In part, the evidence is based on differentia l weathering zones delimited by lateral moraines ( the Koroksoak and higher weathering zones). The probable upper limit of Late Wisconsinan ice is recorded in several local studies, and the regional limits of ice cover are not well constrained (Andrews, 1963;

Québec Clark,1984; Evans, 1984, and Evans & Rogerson,1986). Generally the upper glacial trimline of Saglek Glaciation declines eastward from the watershed divide to the sea, and decreases northwards in overall elevation. Between Fraser River and Okak Bay, ice reached maximum elevations of about 700 m a.s.l. (Andrews, 1963). Mount Thoresby and Man O'War Peak are the most southerly nunataks (Andrews, 1963; Johnson, 1969). The glacial limit in the Saglek Fiord area is 615 m (Smith, 1969). In the Ryans Bay region, Clark (1984) stated that the ice passing through the Torngat Mountains did not reach elevations of more than 800 m on the drainage divide. In summit areas lying above the Saglek Glaciation level, cirque glaciers or small local ice caps existed independent of the LIS (Ives, 1960a; Løken, 1962a; Evans, 1984; Clark, 1984; Evans & Rogerson, 1986; Bell et al., 1987). The extent of grounded glacier ice on the Labrador Sea continental shelf is the subject of controversy. Initial studies by Clark (1984) and the studies of Evans (1984), Evans & Rogerson (1986), and Rogerson & Bell (1986) indicated that the Late Wis consinan ice only extended to fjord mouths. In opposition, Josenhans et al. (1986), and Clark & Josenhans (1986) stated that grounded glacier ice extended well offshore to near the shelf edge. The interpretative differences may be reconciled by low shear strength of subglacial deposits on the shelf. The deglacial history of the rugged coastal areas of Labrador is complex, characterised by local readvances , and has been discussed for various locations by Ives (1958, 1960a), Løken (1962b, 1964), Tomlinson (1963), Andrews (1963), Johnson (1969), Clark (1984), Evans (1984) and Evans & Rogerson (1986). At the limit reached by the Late Wisconsinan ice, extensive systems of lateral moraines and kame complexes were built (Saglek Moraines of Ives, 1976). Several end and lateral moraines mark positions of halts or local readvances of the ice front during retreat from the Saglek Moraines towards an ice mass located west of the Labrador Sea/Ungava Bay watershed. Notable examples are the Tasiuyak Moraines in the Fraser River-Okak Bay area (Andrews, 1963), and the well-correlated Noodleook, Two Loon and Kangalaksiorvik ( = Sheppard) moraines of Løken (1962b, 1964) on the Torngat Peninsula. Andrews (1977) suggested that the Kangalaksiorvik Moraines may be equivalent to the moraines of Cockburn age elsewhere in Arctic Canada (Andrews & Ives, 1978) which are dated about 8,000 y. BP. In their northernmost 100 km, the Kangalaksiorvik (Sheppard) moraines contain abundant indicator erratics derived from the Labrador Trough. They demonstrate net eastwards to north-eastwards ice flow over 100’s of kilometres across Ungava Bay and onto the northernmost Labrador Peninsula. They are interpreted to reflect a major, late glacial readvance after c. 9000 yr BP. (Løken, 1964) that extended to c. 280 m a.s.l. on the northern Peninsula. In contrast, there is no evidence for onshore ice flow in south-eastern Ungava Bay (Allard et al., 1989). Ice-flow indicators between the south-western Torngat Mountains and lower Rivière George (Mathew, 1961), and

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along the east coast of Ungava Bay clearly indicate a LateGlacial reversal of ice flow into Ungava Bay which is interpreted as the result of drawdown toward a calving bay. The area affected by draw-down extended rapidly southwards to the Schefferville area of central LabradorQuébec (Ives, in Vincent, 1989; Parent & Paradis, 1998; Veillette et al., 1999). As a result of glacial isostatic depression, both ice-free and newly deglaciated, low-lying coastal areas were submerged. Generally marine limit progressively decreases in elevation northwards, from 93 m south of Okak Bay (Andrews, 1963) to 16 m on Killinek Island (Løken, 1964). The oldest Holocene age so far obtained on the north coast is 9820 ± 70 BP (TO-305). The presence of tilted shorelines in northernmost Labrador truncated by a 15 m high horizontal shoreline is considered by Løken (1962b) as recording an early Holocene transgression and also suggests that thick continental ice did not overlie the northern tip of Labrador. During deglaciation of the Torngat Mountains, east of the drainage divide, numerous small ephemeral glacial lakes were ponded in tributary valleys, blocked by ice tongues. As the ice margin receded westward from the divide, larger and longer-lived glacial lakes were created. Lakes in tributary valleys of Rivière Alluviaq drained into a fjord south of Iron Strand (Ives, 1957), whereas other lakes, in the Rivière Koroc basin, drained towards Saglek Bay (Ives, 1958), and later into Nachvak Fiord via Palmer River with continued deglaciation and westward marginal retreat. The largest lakes were glacial lakes Naskaupi and McLean (Ives, 1960a, b; Matthew, 1961; Barnett & Peterson, 1964; Barnett, 1964, 1967; Peterson, 1965; Fig. 3), which extended over large areas in the upper Rivière George and Rivière à la Baleine drainage basins. These lakes were dammed between the watershed divide on the east, the south-westwards retreating main body of New Québec ice on the west, and Late-Glacial ice in Ungava Bay (Prest, 1970, 1984). Glacial Lake Naskaupi cut a series of well defined strandlines, some of which are incised into bedrock. Fine-grained glacial lake deposits, however, are virtually non-existant. Glacial Lake McLean, in upper Rivière à la Baleine basin, was separate from Lake Naskaupi but drained into it by a channel west of Lac de la Hutte Sauvage. Both lakes finally drained into the D'Iberville Sea when ice in Ungava Bay had retreated sufficiently to allow free northwards drainage. When the ice receded from the Saglek Moraines is unknown. From marine shell dates, the minimum age for deglaciation of the coastal areas is of about 9 ka. The deglacial history of the central Labrador coast has been used to infer that the glacial lakes were fully established only after 7,600 ± 200 y. ago, and that Late-Glacial ice in Ungava Bay collapsed about 7,000 y. B.P. (Clark & Fitzhugh, 1990). Abandoned moraines and lichen-kill areas adjacent to present cirque glaciers and glacier-free cirques provide a record of Neoglacial expansion of glaciers in the

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Torngat Mountains (McCoy, 1983; Clark, 1984; Evans, 1984, Evans & Rogerson, 1986).

The latest deglaciated areas in Nunavik, Central Ne w Québec and western Labrador Since Low (1896) first recognised central Labrador-Ungava Peninsula as one of the final centres of ice disintegration, controversy has surrounded the situation of residual ice centres of the Labradorean Sector. As portrayed in Wilson et al. (1958), Ives (1960a), Prest et al. (1968) and Prest (1969), net patterns of glacially-streamlined landforms and eskers indicate that New Québec Ice flowed radially from a horseshoe-shaped area extending from north-west of Lac Delorme in the west, to northern Smallwood Reservoir in the east. In the area enclosed by the horseshoe, ice flow was convergent northwards towards Ungava Bay, whereas outside ice flow was broadly radial outwards. Whether the ice divide, determined from the landforms, was a stable feature of Late Wisconsinan ice, or whether its location fluctuated considerably, or whether it existed at all (e.g. Kirby, 1961a, b) has only recently been addressed. Indicator erratics derived from bedrock of the Labrador Trough have been used to constrain net glacier flowpaths and key aspects of glacial history. For example, Hughes (1964) and Richard et al. (1982) suggest that the ice divide must at one time have been situated well to the north-east of its final position because indicator erratics were transported from areas east of the assumed ice divide. Ives (1960) identified evidence for transport from areas west of it. In common with Shield terranes elsewhere in Canada (e. g. Veillette et al., 1999, Parent, et al., 1996; Klassen & Thompson, 1993), the erosional record of striations has been proved to provide a comprehensive basis for reconstructing much of Wisconsinan glacial history, not only the Late-Glacial. Along with the evidence for net distances and directions of ice flow provided by indicator erratics, striations show a complex glacial history for LabradorQuébec (Bouchard & Marcotte, 1986; Bouchard & Martineau, 1985; Kirby, 1961a, b; Ives, 1960; Henderson, 1959; Hughes, 1964; Klassen & Thomson, 1993; Parent et al., 1999; Veillette, 1986; Veillette & Roy, 1995; Veillette et al., 1999), and indicate that the depositional record of landforms does not define a simple integrated record of ice flow in the marginal areas of a decaying ice sheet (e.g. Dyke & Prest, 1987; Kleman et al., 1995). From the relative ages shown by striations, the ice divide determined from the glacial landform record is the product of multiple, distinct ice flow events. It does not reflect a single glacial configuration of the Labradorean Sector but is the integrated product of multiple glacial events (Klassen & Thomson, 1993; Veillette et al., 1999). Field investigations in the Schefferville area of the Labrador Trough by Perrault (1955), Grayson (1956), Henderson (1959), Ives (1959, 1960a, b, c, 1968, 1979), Kirby (1961a, b) and Derbyshire (1962), brought conclusive evidence for the presence of small ice remnants in the

low-lying basins of Howells River (Kivivic ice divide; just west of Schefferville) and Swampy Bay River valleys (north-north-west of Schefferville). Other authors, basing their arguments on glacial ice flow indicators and landforms (Low, 1896; Hughes, 1964; Laverdière, 1967; Richard et al., 1982) and on the intersection of projected strandline tilt directions of glacial lakes as an indicator of the position of the maximum ice thickness (Ives, 1960b; Harrison, 1963; Barnett, 1964; Barnett & Peterson, 1964), proposed that the final ice masses disintegrated in the Schefferville area, near the southern base of the ‘horseshoe shaped’ ice divide. Much controversy has ensued between the different authors (Ives, 1968; Bryson et al., 1969; Laverdière, 1969a, b; Laverdière & Guimont, 1982), but it is likely that there were numerous retreat centres both along the final location of the divide and in adjacent low-lying basins where discrete ice masses finally melted. Dates from numerous lake sediment cores have been used to date the final disappearance of ice. On or near the final position of the ice divide, basal dates of 6,320 + 180 BP (GSC3094; Richard et al., 1982) in the Lac Delorme area, and 6,200 ± 100 BP (GSC-3644; King, 1985) in the Lac Stakel area provide the best minimal estimates. Apart perhaps from small remnant ice masses in depressions, it is probably safe to assume that glacier ice in the QuébecLabrador interior had completely melted by 6.5 ka. Pending questions on the deglaciation of QuébecLabrador There are two types of problems which limit the complete model of deglaciation in Québec-Labrador, problems related to chronology and the need for geological field evidence. The chronology of deglaciation in southern Québec is established from marine shells of Champlain and Goldthwait Seas and seems too old by as much as 750 years when compared to the New England varve chronology and by about 800 years when compared to the onset of organic deposition in lakes and ponds in the deglaciated Appalachian uplands. The key age is the opening of the Champlain Sea, dated at 12,000 y. B.P. on marine shells (Parent & Occhietti, 1988, 1999) "until the question of local marine 14 C reservoir corrections has been investigated more thoroughly". The mean ocean reservoir effect on this age from marine shells is already corrected to d 18 O = 0 ‰ but the hard water and local reservoir effects are not corrected. Nevertheless, using a saltwater wedge model, Rodrigues (1992) places the beginning of the marine invasion in the deeper parts of Champlain Sea and the coeval highest marine levels between 11,400 and 11,600 y. B.P. or based on the stenohaline Balanus hameri association (Rodrigues, 1988), between 11,400 and 11,100 y. B.P., in western Champlain Sea. On the basis of pollen zones in Candona Lake-Champlain Sea deposits, Anderson (1988) estimated an age of 11,500 y. B.P. for the marine invasion, an age compatible with the early phase of the post-Iroquois falling

Québec Lake phase and the beginning of the Early Lake Ontario low-level phase. From these data, a variable hard water (inorganic dissolved carbon) effect resulting from both the influx of old CO2 previously bounded in glacial ice and from the dissolution of carbonates from bedrock and glacial sediments (Hillaire-Marcel, 1981), would overdate the marine shells by at least 500 years. Two ages of the ultimate phase of Champlain Sea, from shells and wood collected in a non-carbonated sandy tidalite of the SaintNicolas site, close to Québec City, differ by 340 years (9,810 ±,70 BPwith d13 C = 0 ‰ vs 9,470 ± 40 BPwith d13 C =-25 ‰) (Beta-143297 and Beta-143298, Occhietti et al., 2001b). This would be the minimal value of the hard water effect in the marine basin. An acceptable age of the beginning of Champlain Sea transgression in the central St. Lawrence Valley would therefore be 11,650 y. B.P. The New England varve chronology is calibrated with remanent declination and inclination and several 14 C ages, mainly from the glacial Lake Hitchcock varves and from 14 C ages from vegetal material found in varves and lake sediments (Ridge et al., 1999). In this chronology, the interpreted age of the transition from lacustrine to marine sediment of Champlain Sea is c. 11,350 B.P., from a site of Vermont close to the border of Québec (Ridge et al., 1999; Ridge, personal communication). This age would mean a higher hard water effect in Champlain Sea. The minimum ice retreat rate of the upper Connecticut Valley is estimated at 230 m/yr. These two data suggest that the 200 m/yr estimated rate of ice retreat in the Appalachians of southern Québec by Parent & Occhietti (1999) is underestimated. A retreat rate of 250 m/yr seems more consistent with the data from adjacent regions. For comparison, during the past two centuries the ice front retreated by 100 km in Glacier Bay, Alaska (Syverson, 1995). The complete deglaciation of the Appalachians of southern Québec, from the Frontier Moraine to the Ulverton-Tingwick Moraine, over a distance of 85 km, would have lasted 300 to 340 years. This rapid rate is in agreement with the type of recessional moraines which are more likely ice front positions than continuous rims of till and with the strongly lobate outline of the ice front positions. Sometime during the Bølling-Allerød warm phase, the margin of the LIS was thinning and rapidly melting on the Appalachian uplands, between about 11.911.6 and 11.6-11.3 ka. The Champlain Sea invasion happened one century after construction of the UlvertonTingwick Moraine. The second chronological problem is the question of the age of the moraines on the North Shore of the St. Lawrence lower Estuary and Gulf and in southern - south-eastern Labrador. Radiocarbon evidence for the deglacial history of south-eastern Labrador is summarised and discussed by King (1985). Evidence for tundra conditions in southeastern Labrador during the glacial maximum is represented by analyses of fine-grained organic carbon in marine cores (Vilks & Mudie, 1978). The potential for contamination by old carbon makes the radiocarbon dates suspect. From dated lake sediment dates, King places the Late Wisconsinan ice margin east of the Paradise Moraine, with the

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maximum limits possibly at the Bradore Moraine and Belles Amours Moraines. The interpretation of extensive ice cover in south-eastern Labrador is supported by Grant (1992) who places the Late Wisconsinan ice limit offshore, in the Strait of Belle Isle, consistent with the reconstruction of Dyke & Prest (1987). According to Grant (1992), the Bradore and Belles Amours moraines formed about 12.6 ka through regional readvance of the ice margin. The Paradise Moraine is thought to have formed c. 10 ka, possibly as the result of regional climatic cooling during glacial retreat. The Little Drunken Moraine and the Kenamu Moraine are represented as segments of the Québec North Shore Moraine, and they are estimated to have formed about 9 ka BP; the Sebaskachu Moraine to the north of Lake Melville is inferred to be younger, forming about 8 ka BP. The Little Drunken Moraine is associated with glacial draw-down into Lake Melville; no major terminal moraines occur west of it, in the Labrador interior. There are large areas of Québec-Labrador that have never been mapped or simply explored, mostly because of the lack of mineral interest and the inaccessibility (c f. the general map of Quaternary deposits of Canada, Fulton, 1995). This is the case of most of the Laurentians between the Saint-Narcisse Moraine and the Sakami Moraine. Current work in the middle reaches of the Saint-Maurice River shows that there is a high potential for the discovery of new significant evidence about the LIS front retreat. In almost every new large excavation, something new is discovered. The two interglacial deposits at Wabush (reported only in an abstract) are a prime example. There is a problem that is not only unresolved but scarcely looked at that concerns the age, origin and timing of the glacial lakes in the interior, notably Naskaupi, McLean and Koroc, but also the existence of others in the Caniapiscau River basin (glacial Lake Minto, etc.). The problem is compounded by the INSTAAR models on the southern Baffin/western Hudson Strait that focus on the geological record of south-eastern Baffin Island and are difficult to reconcile with the Québec-Labrador record. The problem is also linked to the glacial model applied to the Labrador highlands, specially to the Torngat mountains. The nunatak theorie favors the absence of ice over the mountain tops. Gangloff (1983) shows that tors and felsenmeers are not necessarily indicative of unglaciated areas. Another problem, currently in the process of resolution, concerns the acquisition of ice-flow data, and the construction of a comprehensive ice flow record for much of the last glaciation. The striation evidence has proven to be of inestimable value, and in the last twenty years it has unlocked more than Late-Glacial history, representing much of the last glaciation. It is a key for resolving the configuration of the ice sheet and ice flow dynamics, and one that should prove invaluable for ice-sheet reconstructions and climatic modeling. The existence of an early centre for ice sheet accumulation and dispersal near the St Lawrence Estuary (considered by Occhietti, 1982 and others) has never been adequately accounted for through modeling.

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Conclusion During the Wisconsinan, the area of Québec-Labrador was completely covered by the LIS, except may be for nunataks in the Torngat Mountains of north-eastern Québec-northern Labrador. For this reason, pre-Upper Pleistocene events are only represented in scattered stratigraphic sections, mostly in southern Québec, and by erosional glacial marks. Early Wisconsinan ice flow patterns indicate the LIS could have been initiated in highland regions north of the St. Lawrence Corridor (Valley and Estuary). After the LGM, the Labradorean Sector of the LIS evolved from a single, predominant dispersal centre with subsidiary ice divides, into peripheral ice domes and masses which remained connected or not to the central dome. The central dome was not a simple and stable dome-shaped ice mass, but an evolving ice mass. The general model of deglaciation is still evolving quickly as new field evidence becomes available. Thinning of the LIS through ablation and through mechanical drawdown along its margins as the result of diachronic ice streams in the St. Lawrence Corridor and Hudson Strait are the main features of Late-Glacial ice flow dynamics. In the areas south of the St. Lawrence Corridor, ice masses over the Appalachian uplands evolved from a glacier complex confluent with the LIS into separate local ice caps. During a part of the warm Bølling-Allerød phase, a series of ice front positions mark a fast retreat of the ice front in the Appalachians of southern Québec, between 12,400 and 12,000 y. B.P. (about 11,900-11,500 conventional 14 C y. B.P.). The ice mass over the Canadian Shield, north of the St. Lawrence Corridor, dissipated slowly, between about 11,000 and 6,500 y. B.P. The deglaciation pattern includes the differentiation of an ice mass over the Hudson Bay, early deglaciation of the Labrador Highlands, a major change of ice flow from the southern part of the Ungava Bay towards the Hudson Strait and a very roughly concentric ice retreat pattern in the south-west, south and south-east margins of the remnant main ice mass. A series of concentric moraines (SaintNarcisse, Mars-Batiscan-Saint-Maurice, and intermediate moraines), built between about 10,800-10,000 y. B.P., are related to the Younger Dryas Chron. During the early Holocene, the Harricana interlobate Moraine records the separation of ice in Hudson Bay from the ice mass over Québec and western Labrador. The Sakami Moraine, built c. 8 ka, is related to the drainage of Lake Ojibway toward the Hudson Strait and grounding of the ice front. In the northern Ungava and Labrador peninsulas, major glacial lakes in low-lying areas were dammed between the ice front and the tilted deglaciated land. Lowlands depressed by glacioisostasy were momentarily invaded by marine waters. Fossils from these post-glacial seas (Goldthwait, Champlain, Laflamme, Tyrrell, and Iberville seas) give approximate ages of the deglaciation episodes, between 13 ka and 7 ka. The last glacial ice masses were situated in the Labrador Trough and Nunavik and finally disappeared c. 6.5 ka.

Acknowledgements The compilation of maps and documents, and the preparation of the digital maps of glacial features of Québec were technically and financially supported by the Département de Géographie, Université du Québec à Montréal (UQAM), and the Groupe Atlas du Québec et de ses régions (financed by the Ministère des Régions, Québec). The Glacial Landforms Map of Labrador and parts of adjacent Québec (compilation of R. Klassen) was supplied in digital format by A. Moore, Geological Survey of Canada (Ottawa). Research on the field and by air photography was supported by the Natural Sciences and Engineering Council of Canada (S. Occhietti). Pierre Roy, from the Cartothèque, and Francine Robert, from the MSc programme of Geography, both at UQAM, are personally thanked for their patient assistance.

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