Uncovering the hidden part of a large ice stream of the Laurentide Ice Sheet, northern Ontario, Canada

Quaternary Science Reviews 155 (2017) 136e158 Contents lists available at ScienceDirect Quaternary Science Reviews journal homepage: www.elsevier.co...

Download PDF

14MB Sizes 76 Downloads 70 Views

Report

Full Text

Quaternary Science Reviews 155 (2017) 136e158

Contents lists available at ScienceDirect

Quaternary Science Reviews journal homepage: www.elsevier.com/locate/quascirev

Uncovering the hidden part of a large ice stream of the Laurentide Ice Sheet, northern Ontario, Canada nard b, G. St-Jacques b J.J. Veillette a, *, M. Roy b, R.C. Paulen a, M. Me a b

Geological Survey of Canada, Ottawa, Ontario, Canada Department of Earth and Atmospheric Sciences & GEOTOP Research Center, University of Quebec at Montreal, Montreal, Quebec, Canada

a r t i c l e i n f o

a b s t r a c t

Article history: Received 19 June 2016 Received in revised form 3 November 2016 Accepted 7 November 2016 Available online 24 November 2016

This investigation was prompted by an enigmatic ice-ﬂow anomaly (Area A) on the Glacial Map of Canada which covers about 10 000 km2 in the Hearst/Kapuskasing area of northeastern Ontario. It consists of streamlined landforms and striations indicative of a major ice ﬂow toward 130 oriented at right angle to another toward 220 . Both are late glacial ﬂows but long-lasting disagreement exists regarding their relative age. The analysis of aerial photographs and satellite images in conjunction with a detailed survey of bedrock cross-striated surfaces over an area of about 30 000 km2 within and around Area A clearly indicate that the 130 ﬂow preceded the 220 ﬂow. The earlier conﬂicting interpretations within Area A are attributed mainly to the sporadic occurrence of relict striated surfaces formed by older southwestward (220 e240 ) Wisconsinan ice ﬂows that have locally escaped destruction by late glacial ﬂows, with the result that the southwestward ﬂows are older (Wisconsinan) at some sites and younger (late glacial 220 ) at others relative to the 130 ﬂow. When considered with other factors such as the maximum elevation reached by the youngest late glacial ﬂow, these ice-ﬂow relationships indicate that Area A is the outcropping southern part of a much larger ESE ice-ﬂow system, which is probably related to a large ﬂuted belt located to the north and that was identiﬁed as the Winisk Ice Stream. The distal part of the ice stream, except for Area A, escaped detection by remote sensing mapping methods because depositional and erosional features associated with it are masked by deposits laid down by the younger (220 , Cochrane) ice ﬂow and/or by postglacial marine and organic deposits (or were destroyed by the younger ice ﬂow). The only reliable indicators of the passage of the ice stream in this “buried” section are ESE relict striations crossed by SW striations. The advancing ice stream toward the ESE not only preceded the late Cochrane 220 ﬂow but probably outlasted it, as suggested by the mapping of several thousand iceberg furrows in Quebec and Ontario, also directed toward the ESE and overprinted on ﬂutes formed by the last glacial ﬂow. If this interpretation is correct, this makes the Winisk Ice stream the largest terrestrial ice stream in the Hudson Bay basin. An alternative interpretation associates this ﬂuted belt to the exposed western fringe of streamlined bedforms from a late Hudson lobe buried by younger sediments. Crown Copyright © 2016 Published by Elsevier Ltd. All rights reserved.

Keywords: Laurentide Ice Sheet Deglaciation Cross-striations Relict striations Ice stream Glacial surge Lake Ojibway Continental record of glaciations Satellite imagery Ontario

1. Introduction The reconstruction of former ice-ﬂow patterns from the systematic mapping of mesoscale (10e10,000 m) elongated bedforms, parallel (e.g. ﬂutes) or perpendicular (e.g. De Geer moraines) to the ice ﬂows which produced them provided valuable insights into the evolution of continental ice sheets (e.g. Punkari, 1982; Boulton and

* Corresponding author. E-mail addresses: E-mail address: [email protected] (J.J. Veillette).roy. [email protected] (M. Roy). http://dx.doi.org/10.1016/j.quascirev.2016.11.008 0277-3791/Crown Copyright © 2016 Published by Elsevier Ltd. All rights reserved.

Clark, 1990; Clark et al., 2000; Jansson et al., 2003; De Angelis and Kleman, 2008; Kleman et al., 2010). Despite the beneﬁts of these advances, heightened by the widespread use of relatively recent satellite imagery and digital elevation models (DEMs), remotely sensed data remain limited by technical (e.g. scale and image resolution) and human factors such as reliability of operator interpretation. Smith and Knight (2011) presented a detailed analysis of these constraints and demonstrated that the reconstruction of the last Irish Ice Sheet based on the mapping of a large number of striations added to the landform record is signiﬁcantly more accurate then that based on the mapping of bedforms alone. This

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

paper uses a combination of cross-striated and bedform datasets to reconstruct a sequence of ice ﬂows parts of which are obscured by the masking effect of younger deposits. It underlines the need for ﬁeld measurements and observations in areas where geological conditions render the application of remote sensing methods inefﬁcient. The Glacial Map of Canada (Prest et al., 1968) shows areas of intersecting streamlined landforms in several parts of the country where the relative chronology of the ice ﬂows that formed them remains poorly understood. Area A (Fig. 1) which covers approximately 10 000 km2 and includes two late glacial ﬂows showing well-deﬁned ESE (~130 ) streamlined landforms intersecting a SW (~220 ) glacial lineation surrounding it is one of these problematic areas. The initial objective of the investigation described here was to clarify the controversial origin of this apparent ice-ﬂow anomaly. Striations indicative of ice ﬂows toward the SW, found to be both older and younger than the ESE streamlined landforms and associated striations of Area A, led earlier workers to disagree on the relative age of the ice movements that formed the intersecting ﬂow features (see Section 3). Some proposed that the ESE ice ﬂow is the younger of the two late glacial ice ﬂows present in the general area and others proposed the inverse chronology. We present new evidence that indicates that the ESE ice ﬂow of Area A, although a late glacial event, is not the youngest of the region. Rather it represents the “outcropping” portion of a much larger ice ﬂow toward the ESE that extends toward the north and east of Area A where icedirectional features associated with it are masked by the sediments deposited by a younger late glacial ﬂow normal to it. Our results suggest that both the “buried” and the exposed components of this ﬂow (Area A) belong to the distal part of the Winisk Ice Stream. Although Margold et al. (2015) referred to it as the Ekwan River Ice Stream we retained the Winisk designation since the ﬂuted belt originating in the Winisk River area of upper James Bay was referred to as the Winisk Till deposited by an ice stream (Thorleifson et al., 1993). Three main themes are addressed in this paper. First it is shown that the counter-clockwise sequence of former shifting ice ﬂows characteristic of the James Bay and Abitibi regions of Ontario and Quebec, mostly absent from the landform record and expressed only in stratigraphy and relict striations (Thorleifson et al., 1993; -Loubert et al., 2013) is also present in Veillette et al., 1999; Dube the study area (Fig. 2). This was a necessary preliminary step to establish the precise ranking of the ESE ﬂow of Area A and its palimpsest extension since striations formed by the youngest ice ﬂow of the area and some former ice ﬂows are both in the same general southwest direction. The second part of the investigation focuses on the analysis of the unusual cross-striated surfaces found within certain parts of the study area. Earlier interpretations illustrating the lack of consistency in the ice-ﬂow relative chronology determined from striations are discussed and compared with the new data presented here. Lastly, evidence is presented to demonstrate that Area A and its palimpsest extension is the distal part of a much longer (~800 km) Winisk terrestrial ice stream previously thought to end in the Albany River area. This interpretation is based on the premise that the ﬂuted belt, which constitutes the Winisk Ice stream as deﬁned by Thorleifson et al. (1993), is a discrete event. The possibility that the ﬂuted belt may instead represent the outcropping part of streamlined terrain formed by a late Hudson lobe buried by younger sediments is suggested. The new results illustrate how a detailed and regional striation record is in certain cases a pre-requisite for the identiﬁcation and mapping of palimpsest ice ﬂows and dispersal trains in areas of the Canadian Shield where landforms shaped by earlier ice ﬂows were totally or partially destroyed or masked by younger deposits.

137

2. Regional setting Most of the study area (Fig. 2) lies below 300 m asl (above sea level) and shows a broad amphitheatre-like topography, open and gently inclined toward the north. It is part of the Severn and Abitibi upland divisions of the physiographic Hudson region of Canada (Bostock, 2014). Archean undifferentiated felsic plutonic, gneissic, and migmatic rocks of the Abitibi Subprovince of the Superior structural province underlie most of the study area (Percival and Card, 1985). Maﬁc, intermediate volcanic rocks, volcanogenic sediments and minor ultramaﬁc basalt of the Abitibi Subprovince occupy the central eastern portion of the study area (Abitibi greenstone belt) along with felsic and intermediate volcanic rocks (Jackson and Fyon, 1991). Silica sands, kaolin, silt and lignite of the Lower Cretaceous Formation and Paleozoic limestone, dolomite, evaporites and terrigenous clastic sediments occur just to the north of the western half of the study area (Telford and Long, 1986). Glacially transported fragments of these younger rocks are abundant in the late glacial sediments south of 50 Lat. N. Bedrock outcrops are rare. Surﬁcial geology maps cover about 40% of the study area (Fig. 2), which is characterized by extensive belts of ﬂuted terrain composed of carbonate-rich pebbly clay moulded by late glacial readvances. The ﬂuted belts cover most of the area below 300 m and the glacial lineation they created is readily identiﬁed on digital elevation models (DEMs) shown on two color maps at 1:250 000 scale (Supplementary Fig. SD1 and SD2, on-line Appendix) derived from the Shuttle Radar Topographic Mission (SRTM). The range of elevation selected was 180e315 m, the same as shown in Fig. 2, which includes all the ﬂuted terrain formed by the two late glacial ﬂows. 3. Earlier work e the controversy Low (1889, 1900) was the ﬁrst to demonstrate from crossstriations measured on outcrops along the Missinaibi River that the ESE (~130 ) ice ﬂow predates the SW (~220 ) ﬂow of Area A. Based on striations indicative of southeastward moving ice reported by Wilson (1904) to the northwest of Area A and on his own measurements near Lowther, Boissonneau (1966) recorded the same ice-ﬂow chronology as that reported by Low and proposed that the ESE ﬂow of Area A extended to the north but could not be detected from aerial photographs because its ice-directional features are masked by deposits from the youngest SW ﬂow. Richard and Hilborn (1984a and b) reported the same ice ﬂow chronology in the vicinity of Hearst as that observed by Boissonneau near Lowther. Prest (1966), however, attributed the ESE ice ﬂow of Area A to a late ice surge (Kapuskasing surge) that occurred during the ﬁnal retreat of the ice from the area. Smith (1992) based on striations measured in the Kapuskasing area also proposed that the ESE ice ﬂow of Area A was the youngest glacial event of the region and rejected Boissonneau's interpretation. Paulen (2001) also considered Area A to represent the latest Cochrane event based mainly on geomorphic features interpreted as frontal moraines. 4. Methods The analysis of cross-striated surfaces described here rests on the assumption that the striated facets preserved on the lee sides of low-angle bedrock bosses were, in part at least, sheltered from the erosive action of subsequent ice ﬂows that came in contact with the outcrops from directions others than those of the ﬂows which formed them. These sheltered striated surfaces constitute reliable indicators of the orientation and in some cases of the direction of former ice ﬂows. Most outcrops selected for analysis were bi-

138

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Fig. 1. A. Part of the Glacial Map of Canada (Prest et al., 1968) showing the anomalous glacial lineation in Area A south of Hearst and Kapuskasing. B. Area A (within red lines) as mapped by Boissonneau (1966). (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the web version of this article.)

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

faceted or tri-faceted rock bosses bearing striated surfaces intersecting along well-deﬁned planes. Cross-striated outcrops of maﬁc, ﬁne-grained metavolcanic rocks were paid particular attention because they preserve bedrock-inscribed features formed by moving ice more effectively than the dominant, coarser-grained, pale, granitoids of the study area. Whenever possible, striated surfaces were described and photographed under favourable light conditions. Slightly wet striated surfaces exposed to a low sun angle enhance the shadows of minute depressions, ﬁne striations, grooves and truncations. Fine intersecting linear features developed on some light-coloured, coarse-grained, granitoid, striated surfaces may not be visible otherwise. The cover of unconsolidated sediments was removed at several sites in search of fresh, unweathered, striated surfaces. Previously unreported striations indicative of former ice ﬂows toward the SW cross-cut by striations toward the ESE which are in turn overprinted by striations toward the SSE (orientation of the youngest late glacial ﬂow in the eastern half of the study area, see Fig. 10) were found as far north as the Pinard Moraine and as far east as the Little Abitibi and Pierre lakes area during a surﬁcial geology mapping program (Veillette et al., 2008, Fig. 2g). The unexpected presence of ESE striations at these locations hinted at a possible continuity with striations of similar orientation known to occur in the Hearst area to the west and prompted a re-examination of Area A and its presumed northern extension (Boissonneau, 1966). The relative ice-ﬂow chronology for the study area is derived primarily from measurements at 136 selected crossstriated sites, most of which were obtained during the current investigation (Figs. 3, 4, 8 and 12 and Supplementary Figs. SD1 and SD2). The stereoscopic examination of aerial photographs, surﬁcial geology maps, borehole records in unconsolidated sediments (Smith, 1992) and the interpretation of the glacial lineation from SRTM DEMs and on-line Google Earth imagery provided additional information to constraint the striation record. Conventional names that imply a recognition of the origin of the ﬂow (Ex: Cochrane I, Cochrane II, Matheson Till) were avoided as

139

much as possible and were used only to conform to the terminology used by earlier workers. From here on the two late glacial ﬂows are referred to as LG1 (toward the ESE) and LG2 (toward the SSW, S and SSE). 5. Results 5.1. Wisconsinan ice ﬂows Three broad directions of ice ﬂows (Figs. 3e6) preceding two late glacial ﬂows (Figs. 8 and 10) were identiﬁed in Ontario and Quebec. The oldest two are known only from the striation record and from stratigraphic data elsewhere in the James Bay basin (Thorleifson et al., 1993; Roy, 1998; Veillette et al., 1999; Dube Loubert et al., 2013). (1)The earliest indicates movement toward WNW (270 e300 ), and all measurements except one are located in the eastern half of the study area (Fig. 3). The reverse chronological order of formation at this site (NW ﬂow is youngest, NTS map 42G03, from Morris, 2002a; Supplementary Fig. SD1) probably results from compilation errors and is tentatively included as the westernmost site where WNW striations were recorded. The southernmost striations (285 and 275 ) attributed to this ﬂow (Paulen and McClenaghan, 1997) are in NTS maps 42A13 and 14 (Supplementary Fig. SD1). (2) The second ﬂow includes cross-striated sites showing westward and southwestward striations (<270 ) overprinted by striations from late glacial ﬂows (Fig. 4). The distinction between this ﬂow and the earliest one is arbitrary since both belong to the same counter-clockwise shift. Over 3000 striations including hundreds of cross-striated sites, were measured in the Abitibi Greenstone Belt of Quebec and Ontario (Veillette and McClenaghan, 1996) and in the James Bay and central Quebec areas to document this shift (Veillette et al., 1999).

Fig. 2. Topography of study area, northeastern Ontario, and location of surﬁcial geology maps from earlier work. A, Aa: Hughes (1956, 1959); B: Richard and Hilborn (1984a, b); C: Morris (2002a, b, c, d); D: Paulen and McClenaghan (1998a, b); E: St-Jacques (2011); F, Fa: Gao (2013a, b); Veillette et al., a,b,c,d (in press). Area A mapped by Boissonneau (1966) is shown within the solid line and its palimpsest extent within the dashed line.

140

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Fig. 3. Cross-striated sites bearing striations formed by the oldest ice ﬂow toward the WNW (1) within the study area. Black arrows show dominant direction (n ¼ 15).

(3) The last main Wisconsinan ice ﬂow was toward the SSE (165 e180 , Fig. 5) and is associated with a late phase of the sandy Matheson Till (Hughes, 1959; Paulen, 2001). Within the study area Matheson Till is nearly everywhere (below 300 m) overlain by deposits of the late glacial ﬂows and because of this is encountered almost exclusively in sections and boreholes. Where Matheson Till was not overridden by

late glacial ﬂows, as is the case in the eastern part of the study area above 300 m (NTS maps 32E04, 32E05, 42H01, 42H08; Supplementary Fig. SD1), it forms extensive streamlined terrain. The sequence of Wisconsinan ice ﬂows just described is presented schematically using overlapping arrows (Fig. 6). The oldest

Fig. 4. Cross-striated sites bearing striations formed by the shift from a WSW to a late SSW direction (WSW ¼ 1). Overlapping arrows illustrate the counter-clockwise shift in the direction of ﬂow (n ¼ 51).

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

141

Fig. 5. Streamlined landforms (large arrows) and striations toward the SSE (3) formed by the ice ﬂow that deposited a late phase of the late Wisconsinan Matheson till (n ¼ 27).

Fig. 6. Schematic illustration (overlapping arrows) of the counter-clockwise shift in ice-ﬂow from an initial WNW to a ﬁnal SSE direction that occurred over the study area prior to the onset of late glacial ﬂows.

142

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Fig. 7. Counter-clockwise shift in ice-ﬂow direction determined from relict striations within and along the coast of James Bay (adapted from Veillette, 1995).

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

WNW ice ﬂow (black arrows) was interpreted as the northwestward expansion of the Labrador Sector of the Laurentide Ice Sheet (LIS) in early Wisconsinan time from an initial ice divide in the Quebec highlands that migrated to the east of Hudson Bay where ice ﬂow gradually shifted to the SW (Veillette et al., 1999). The best examples of this counter-clockwise shift that covers a large area of the Labrador Sector of the LIS, were found on islands and along the east coast of James Bay (Fig. 7).

5.2. Late glacial (LG) ice ﬂows The Cochrane surges (Hughes, 1959; Hardy, 1977, 1982; Dredge and Cowan, 1989; Veillette et al., 1991; Paulen, 2001; Breckenridge et al., 2012; Stroup et al., 2013; Roy et al., 2011) refer to late glacial readvances in Lake Ojibway over distances of 75e100 km or more near the end of the deglaciation of northern Ontario and northwestern Quebec. The fast-ﬂowing ice overrode, grooved, incorporated and mixed the soft, ﬁne-grained Ojibway sediments with a mixture of Precambrian crystalline rock fragments and far-travelled sedimentary rocks from the Hudson Platform. The result is a carbonate-rich, pebbly, silty-clay diamicton (Cochrane Till) deposited as a veneer commonly less than 2 m thick, below 300 m. Smith (1992), however, reported thicknesses of Cochrane Till and pebbly clay in excess of 10 m in boreholes and Paulen (2001) noted a thickening of Cochrane Till near its southern limit in the Timmins area, Ontario from borehole records (McClenaghan and DiLabio, 1995). Boissonneau (1966) recognised two major Cochrane surges (see Fig. 18) in the western half of the study area, (1) an ESE (~130 ) ice ﬂow within Area A and its extension to the north, which he referred to as the Early Cochrane phase (LG1 in this paper), and (2) a SW (~225 ) ice ﬂow he referred to as the Last Cochrane phase (LG2 in this paper) that left striated surfaces overprinted on those of the ESE ﬂow. The surges were labelled Cochrane I and Cochrane II, respectively, by Prest (1970) and Richard and Hilborn (1984a and b).

143

This terminology is still commonly used today. Striations toward the ESE (LG1) were found over about 50% of the study area south of 50 N. but a distinct glacial lineation (Area A) in this direction is visible on aerial photographs and other types of imagery only south of the Hearst-Lowther-Kapuskasing axis (line joining the three localities and referred to from hereon as HLK) (Fig. 8 and Supplementary Fig. SD1). Over more than 60% of LG1 known extent the only evidence of its passage is in the form of relict striations (Fig. 8). The distribution of LG1 striations reveals a slight shift from an ESE direction in the southern part to an eastward one in the northern part. Just to the south of the Pinard Moraine, near its northeastern limit, well-preserved bi-faceted and tri-faceted bedrock bosses exposed through construction activities provide strong evidence that ice ﬂowed in a general eastward direction (~95 e120 ) after ﬂowing toward the WSW and SSW and before ﬂowing toward the SSW and SSE (Fig. 9). Other sites with similar cross-cutting relationships were found to the west of Pierre and Little Abitibi lakes (Fig. 8) (NTS map 42H07, Supplementary Fig. SD1). The SW-SSE ice ﬂow (LG2) formed a general fan-shape pattern of streamlined landforms roughly centered on Smooth Rock Falls, which appears to be dictated by the amphitheatre-like topography of the study area (Fig. 10). Streamlined landforms are characteristic of this ﬂow that was redirected by topographical obstacles and truncated Lake Ojibway glaciolacustrine and glacioﬂuvial deposits in several places (Hughes, 1959; Paulen, 2001). The LG2 surging ice shaped spectacular straight and curvilinear ﬂutes, some continuous for several kilometres, usually less than 6 m high and separated by poorly-drained peat-covered swales. LG2 is clearly the latest ice ﬂow of the study area. The ﬂutes are best seen on SRTM images illuminated from the east (Supplementary Fig. SD2). The complete sequence of former and late glacial ﬂows, is shown on Fig. 11. The intermediate age of LG1 and its “buried” extension is obvious from the superposition of arrows. Because LG1 lies roughly normal to all other ﬂows it was of pivotal importance to determine the ice-ﬂow

Fig. 8. Cross-striated sites indicative of ice ﬂowing toward the ESE (LG1), large arrows mark the approximate extent of the ice ﬂow, no distinction is made between Area A and its palimpsest extension (n ¼ 57).

144

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Fig. 9. A tri-faceted rock boss uncovered by excavation showing a striated facet eroded by a ﬁrst ice ﬂow toward 240 (shovel, upper left, points in that direction) sheltered from the erosive action of a second ﬂow (LG1, felt pen, top center of image) toward 100 (toward the photographer) and from a last ﬂow (LG2) toward 170 (toward left, felt pen bottom of image) that left striations over the summit of the outcrop.

chronology of an area that includes ice ﬂows of similar directions but of widely different ages. 6. A zone of ambiguous cross-striated surfaces Within Area A LG1 striations are oldest at some sites, youngest at others, or have an intermediate age with respect to other striations at others. These differences in relative age were sometimes observed at some sites located only a few hundred metres apart. Area A occupies most of the terrain located south of the HLK axis (Figs. 1b and 12), but in a narrow band, 15e30 km wide and about

150 km long, south and parallel to it, poorly-developed LG2 streamlined landforms are found, along with prominent LG1 streamlined landforms (Supplementary Fig. SD1). Small frontal morainic ridges formed by LG2, which are widespread in the northern part of NTS maps 42G01 and G02 (Morris, 2002c, 2002d), are absent at the mapped margin of LG2 (dotted line on Fig. 12). South of this margin SW striations were never found overprinted on ESE striations and when present are always overprinted by ESE striations which indicates they result from the counter-clockwise shift of the LIS and that LG2 did not extend that far south. It is within the elongated east-west zone, wedged between the HLK axis

Fig. 10. Generalised ﬂow directions derived from the orientation of ﬂutes for the Cochrane event (LG2) in the study area.

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

145

Fig. 11. Generalised sequence of ice ﬂows in the Hearst-Cochrane area of northeastern Ontario determined from the mapping of cross-striations and the orientation of streamlined landforms.

Fig. 12. Extent of LG1 including its “buried” palimpsest extension (shaded area) to the north and east of Area A; the dashed line shows the approximate northern and eastern limits of the ﬂow determined from relict striations and the dotted line marks the southern limit of LG2.

146

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

and the southern limit of LG2, that most ambiguous measurements on the ice-ﬂow chronology were reported from earlier work and from observations gathered during the current investigation.

6.1. The source of the conﬂicting interpretations Several cross-striated sites where earlier workers had reported conﬂicting interpretations on the ice-ﬂow chronology within and north of Area A were revisited during this study. Striations eroded by southwestward Wisconsinan ice ﬂows occur randomly throughout the study area and explain the lack of consistency of the striation record within the area occupied by LG1 and LG2. Where LG1 did not leave striations, it is not possible to distinguish striations formed by a Wisconsinan ﬂow from those formed by LG2 on the same outcrop since both show the same general southwestward direction (unless both striated facets differ by several degrees and show distinct intersecting lines). This situation cannot exist in the eastern part of the study area where striations from only one late glacial ﬂow (LG2) toward the SSE intersect Wisconsinan striations toward the SW (Fig. 12). Similarly, the constraint imposed by two ice ﬂows of similar direction but of different age is eliminated

Fig. 14. Outcrops illustrating the vanished protector effect. A. A clear truncation (dashed line) separates the LG1 facet (105 , black pen, ﬂow toward the photographer) from the LG2 facet (225 , small felt pen, ﬂow toward left of image. B. The same crosscutting relationship with similar azimuths is observed on this rock boss that shows gentler slopes than the one in A.

south of the limit of LG2 in Area A, where bi-faceted and tri-faceted outcrops show numerous Wisconsinan relict striated surfaces overprinted by LG1 striations (Fig. 13a and b; NTS maps 42G01-0203-04, Supplementary Fig. SD1). Two speciﬁc ﬁeld conditions discussed below, one related to the elevation of the sites and the other to the preservation of relict striated surfaces by layers of sediment added to the complexity of interpreting ice-ﬂow chronology from cross-striated sites within Area A. These factors produced a wide array of cross-striated conﬁgurations that, to the untrained eye, may lead to disenchantment with the striation method used to reconstruct the chronology of past ice ﬂows.

Fig. 13. A. Striations from a former ﬂow toward 220 (black tip of white felt marker, top of image, points in this direction) preserved on the sheltered facet of a granitoid outcrop shaped by LG1 that left a striated summit toward 150 (black pen points in this direction). Truncation along azimuth 220 (dashed line) separates the two striated planes, the dark tone of the rock surface is due to staining by vegetation. B. Striations toward 240 (black tip of felt marker points in this direction) preserved in the sheltered ﬂank of a large groove toward 210 (small felt marker in upper left of image and striations in central upper part of image), preserved in the lee of a granitoid outcrop with a summit striated toward 155 (black pen points in this direction) with crescentic marks indicating the direction of ice ﬂow.

6.1.1. Site elevation The southern limit and distribution of the late Cochrane surge (LG2) of northern Ontario and northwestern Quebec is controlled by topography. Richard and Hilborn (1984a, b) reported an elevation of 256 m for Cochrane till (Cochrane II) in the Hearst area and LG2 deposits or ice-ﬂow indicators were not found above 275 m within Area A during the current investigation. North of La Sarre, a few kilometers to the east of the Quebec/Ontario boundary, the late Cochrane limit is marked by a broad arcuate northeast-trending, berm-like band, up to10 km wide and about 120 km long, with a maximum elevation in the 290e295 m range. It rises about 5 m above the poorly-drained, streamlined, predominantly peatcovered Cochrane terrain, sloping gently to the north. This poorly-deﬁned feature separates the ﬂuted Cochrane terrain to the

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

north from the undisturbed Lake Ojibway deep-water facies to the south (Veillette and Thibaudeau, 2007). Comparison of Low's (1900) historic measurements taken along the Missinaibi River banks with others taken nearby at higher elevation illustrates clearly the effect of elevation on the presence (or absence) of LG2 striations. At Sharp Rock Rapids below 228 m (Low's northernmost site in NTS 42G06, Supplementary Fig. SD1), he recorded striations toward 130 crossed by striations toward 208 , and at Albany Rapids, about 15 km up river, below 243 m, measurements with the same age relationship (135 and 208 ) were obtained. Less than 20 km to the west of Low's observations, measurements acquired during this study at six different sites in NTS maps 42G05 and 42G12 show the inverse ice-ﬂow chronology (210 is older than 130 ) at ﬁve locations above 275 m; the only one that shows the same ice-ﬂow chronology as that of Low's sites is below the 275 m level. Smith (1992) concluded that LG1 was the youngest ice ﬂow within Area A based on measurements obtained at two crossstriated sites located in the vicinity of 260 m in the Kapuskasing area (NTS map 42G07, Supplementary Fig. SD1). Six other nearby cross-striated sites, however, within map 42G07 and all located below 250 m, show the inverse chronology. Field visits in the same area done during the current study suggest that her observations are correct at the sites she recorded but are obviously not representative of the whole of Area A. Differences in elevation of only a few metres between cross-striated sites are sufﬁcient to inverse the relative ice-ﬂow chronology in the distal parts of LG2. In the Missinaibi River valley and in the other river valleys of the region the late Cochrane ice (LG2) penetrated further south due to the lower elevation, was thicker and with a greater capacity to form striated surfaces than on the interﬂuves (see Fig. 10 and Supplementary Fig. SD1 for topography). 6.1.2. The vanished protector A bedrock boss or a layer of sediment (commonly till) that has sheltered striated surfaces formed by earlier ice ﬂow(s) from the destructive action of a subsequent ﬂow(s) and was later removed by natural processes or human intervention was referred to as a “vanished protector” (Veillette and Roy, 1995). Excellent examples illustrating this process in a natural setting were observed, over large (hundreds of meters) and small (centimetres) surfaces on the Archean rocks along the east coast of James Bay (Veillette and Roy, 1995) and by others along the west coast of Hudson Bay (McMartin and Henderson, 2004; Trommelen and Ross, 2011). The uneven ﬂoor of large excavations that require the removal of unconsolidated deposits to expose bedrock (open pit mines, hydro-electric dam sites, and the periphery of quarries) presents favourable conditions to observe older striated surfaces, which were “protected” in down-ice cavities from erosion by younger ice ﬂows (Veillette et al., 1999). South of the HLK axis, LG1 was found to erode bedrock to a greater depth than LG2 but north of the axis the erosive power of LG2 increased northward with increasing ice thickness (decreasing elevation). There, LG1 is consistently older than LG2 and the ambiguity in the relative age of the two ﬂows observed in the vicinity and south of the HLK axis is not present in this area (NTS maps 42G10 - G15, Supplementary Fig. SD1). Outcrops illustrating vanished protector conditions were observed in a borrow pit in the Kapuskasing area. There, pristine, unweathered, cross-striated surfaces of low -angle (<10 ) rock bosses that had remained covered by a few meters of clayey till until recently show clear truncation lines separating the LG1 and the LG2 facets with the oldest (LG1) untouched by the younger ice ﬂow (Fig. 14). The absence of LG2 erosional marks on LG1 facets is tentatively attributed to a protective layer of sediments (vanished protector)

147

deposited on the outcrop by LG1 after it striated it. Both striated facets were preserved under a cover of sediments (of different age) and exposed at the same time during the excavation of the pit (see Fig. 15c for a different example of the vanished protector). 7. Decoding the striation record in area A Ice ﬂows within Area A show striations toward: (1) the ESE only, (2) the WSW or SW only, (3) the ESE overprinted on WSW or SW striations, (4) the ESE commonly overprinted by striations toward the SW (~225 ) and in some places toward the SSE (west part of Area A), and (5) the WSW-SW, ESE and SW-SSE; here striations from a Wisconsinan ice ﬂow occur on the same outcrop with LG1 and LG2 striations. Three basic conﬁgurations of cross-striated surfaces were observed mostly in the vicinity of the HLK axis and are described using ﬁeld photographs (Fig. 15). In Conﬁguration 1, the summit of the outcrop exhibits LG2 striations that may be weakly or strongly overprinted on LG1 striations, which are also preserved in the lee side of the outcrop, sheltered from the erosive action of LG2. Conﬁguration 1 striations are common at low (<260 m) elevation in the vicinity of the HLK axis and very common north of it. These outcrops bear crosscutting relationships between the two late glacial ﬂows only and in the expected chronological order. Two wetted outcrops composed of ﬁne-grained maﬁc intrusive rocks illustrate this conﬁguration (Fig. 15a and b) and attest to the erosional vigour of LG1 compared to that of LG2. Conﬁguration 2 represents the striation record typical of the area south of the LG2 limit where only striations formed by Wisconsinan ice ﬂows and by LG1 were found (see Figs. 12 and 13). The LG1 striations occupy the summit of the outcrop and may be overprinted on WSW and/or SW striations from Wisconsinan ice ﬂows, which are also preserved in the lee side of the outcrop, sheltered from the erosive action of LG1. This conﬁguration, however, occurs sporadically within the area occupied by LG2, both within Area A and within the palimpsest extent of LG1. It was observed at a few sites in the vicinity of the HKL axis where conﬁguration 1 is widespread. An area (0.5 ha) showing several small (<10 m2) outcrops composed of ﬁne-grained, maﬁc intrusive rock located about halfway between Hearst and Kapuskasing, some of which bear Conﬁguration 1 striations and others Conﬁguration 2 striations illustrates these unusual ﬁeld conditions (Fig. 15c and d). The outcrops were masked until recently by an estimated 2e3 m cover of pebbly clay deposited by LG2. This site demonstrates the difﬁculties associated with the interpretation of cross-striated surfaces showing unexpected gaps in the striation sequence. Both 15c and d photographs indicate that LG1 truncated a 220 striated surface formed by a Wisconsinan ice ﬂow. This relationship is more obvious in 15c than in 15d because in the latter both ﬂows appear to compete for the occupation of the summit of the outcrop. The most likely process to explain the absence of LG2 striations on the summits of some outcrops at this location (below 250 m) and their presence on other outcrops nearby is the vanished protector. LG1 grooved and striated the outcrops bearing striations from a Wisconsinan ice ﬂow toward the SW and left a cover of sediments that LG2 was not competent to remove at certain locations, thus preventing bedrock surfaces bearing LG1 striations from being overprinted by LG2 striations. This protective layer was later removed by excavation. In Conﬁguration 3 the summit of the outcrop exhibits dominant LG2 striations that may be overprinted on LG1 striations in places, which are also preserved in the lee side of the outcrop sheltered from the erosive action of the younger LG2 and intersect striated facets from former ice ﬂows toward the WSW or SSW. Conﬁguration 3 shows basically the same arrangement as Conﬁguration 2 to

148

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

which are added LG2 striations. Nine outcrops bearing striations from both late glacial ﬂows and one or two striated surfaces from Wisconsinan ice ﬂows were recorded within Area A and its “buried” extension during this study and by Richard and Hilborn (1984b). Fig. 15e shows an outcrop of ﬁne-grained granitic rock in the Hearst area bearing three striated facets with striations toward 220 , 130 and 175 in this order of formation. The new data presented here clarify earlier conﬂicting interpretations on the relative ice-ﬂow chronology of the region and are summarised as follows; the ESE (LG1) glacial lineation of Area A was formed before that toward the SW (LG2), and these two ice ﬂows postdate others toward the west and SW (Wisconsinan ice ﬂows), for which evidence is only found in relict striations. The streamlined landforms of Area A, toward the ESE constitute the only “visible” part of the ﬁrst late glacial ﬂow (LG1) that extends to the north, NWt and NE of Area A. In these parts of the study area the geomorphic expression of LG1 was either destroyed or is masked by

younger deposits left by LG2 over a large area for which ﬁrm limits remain to be determined. There, striations toward the ESE crossed by striations toward the SW are the only useful tool to map the extent of LG1. 8. LG1: the southern component of a large ice-ﬂow system (ice stream?) The results of earlier studies combined with the mapping and analysis of relict striations, aerial photographs, SRTM and Google Earth images within and around the study area and relevant data gathered during a surﬁcial geology mapping program lead us to propose that LG1 is probably the ‘hidden’ unreported, distal part of the Winisk Ice Stream of Thorleifson et al. (1993). The exposed parts of the ice stream consist of a large ﬂuted belt which extends southward, uninterrupted, from a converging portion conﬁned between two topographic highs in the Winisk River area to the

Fig. 15. Conﬁguration 1 is illustrated by intersecting striations from LG1 and LG2. A. On an outcrop in the Kapuskasing area (Moonbeam), LG1 (110 , pen, ﬂow is toward the left side of photograph) shaped a well-developed stoss face with grooves and deep striations occupying most of the summit of the outcrop while LG2 (225 , compass, ﬂow away from photographer) eroded a smaller, less-developed stoss face but left faint striations on the up-ice ﬂanks of small grooves formed by LG1 on the summit of the outcrop. B. On an outcrop in the Hearst area (Hallebourg), LG1 (120 , black pen points toward 120 ) grooved and striated the whole outcrop and the larger grooves it formed were not totally erased by the later and weaker LG2 (205 , felt marker, upper left, black tip points in this direction) which also left faint striations on the up-ice ﬂanks of LG1 grooves located on the summit of the outcrop. LG2 striations are visible toward the base of the sheltered face eroded by LG1 (bottom right corner of photograph). C and D. Conﬁguration 2 is illustrated by intersecting striations from LG1 and a Wisconsinan ice ﬂow. In (C) LG1 grooves and striations toward 120 (toward photographer, black pen) occupy the summit of the outcrop and truncate (dashed line) striations toward 220 (toward left of photograph, stabilo pencil) preserved in sheltered positions; in (D), grooves and striations from LG1 toward 120 (toward upper left of image, black pen) occupy the summit of the outcrop and appear to truncate striations toward 215 (stabilo points in this direction). E. Conﬁguration 3 is illustrated by an outcrop in the Hearst area bearing 3 distinct, well-striated facets with the oldest one toward 220 (black tip of white felt pen points in this direction), an intermediate one (LG1) toward 130 (black pen points in this direction) and a youngest one (LG2) toward 175 (toward right of image, small felt marker) that occupies the upper surface of the outcrop.

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

149

Fig. 16. Map showing the northern component (ﬂuted belt) of the Winisk Ice Stream, terrain elevation, the limits of the Hudson Platform and Tyrrell Sea transgression and extension of the ice stream within the northwestern part of the study area.

150

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Albany River area (Fig. 16). The Winisk Till, a thin, massive diamict with fabric and striations on boulder pavements at the lower contact of the till parallel to the ﬂutes overlies a till deposited by southwestward ﬂowing ice and was used to delineate the extent of the ice stream (Thorleifson et al., 1993). Till veneers and striated boulder pavements were observed in association with surge deposits during this study in Quebec and Ontario and are common in terrestrial ice streams (Clark, 1991; Stokes and Clark, 2001; Lian et al., 2003; Jennings, 2006; Cofaigh and Stokes, 2008; Paulen and McClenaghan, 2015). South of the continuous ﬂuted belt ending at Albany River a 75 to 100 km-gap occurs in the ﬂuted terrain that corresponds to the lowest elevation (<152 m) encountered along the course of the ice stream. Several rivers ﬂowing into the Albany and Kenogami rivers and lying for the most part below the marine limit are located within this gap (Fig. 16). The lack of streamlined landforms in this projected portion of the ice stream is attributed to one of the following, or a combination of: (1) the masking effect of marine sediments; (2) the strong dissection of the landscape by the numerous tributaries present southwest of the Albany and Kenogami river junction; and (3) the cover of Cochrane sediments (LG2) also present to the NW of Hearst. Another important gap in the ﬂuted terrain, clearly visible on remote sensing imagery and also attributed to similar processes, occurs in the upper watershed of Attawapiskat River (Fig. 16). Flutes reappear to the west of the Hearst area at elevation above 225 m (NTS maps 42F09, 10, 15, 16; Supplementary Fig. SD1) and show a pronounced shift toward the

SE as they enter Area A due to deﬂection by higher (up to 350 m) ground to the south. Linear ridges in ﬁne-grained deposits below about 180 m in the northwestern part of the study area show a similar orientation to that of the glacial lineation of Area A and are interpreted as LG1 streamlined features masked by LG2 and/or Tyrrell Sea sediments. This buried ﬂuted terrain, expressed locally on aerial photographs viewed stereoscopically, extends sporadically southeastward to Missinaibi River and as far north as 50 300 Lat. N (Fig. 16). The curvilinear path of the ice stream mimics the asymmetrical bowlshape depression of the James Bay basin (Figs. 16 and 17). The northern part of the ice stream is underlain mostly by sedimentary rocks of the Hudson Platform except for the western ﬂank of the ﬂuted belt that lies over crystalline Precambrian rocks. South of 50 Lat. N the ice stream is almost entirely underlain by Precambrian rocks. The northern component was nearly totally submerged by Tyrrell Sea and the southern extension (within the study area) lies for the most part above the marine limit. On its western and southern ﬂanks the ice stream was restricted by upland and the elevation of the ﬂuted terrain is everywhere below 240e250 m, except in the southern part of Area A where, in forced contact with higher ground it may have reached 330 m. From this point eastward the ﬂow was deﬂected toward a lower elevation (Fig. 17). Thick (2e5 m) layers of bedded and massive sand and silty sand with paleocurrent directions toward the SE are widespread below Cochrane Till and reach 10 m or more in the upper portion of eskers everywhere east of Kapuskasing. This abundance of sand below

Fig. 17. Map showing the southern extent of the Winisk Ice Stream (shaded area), the southeastward dispersal of sand from the Mattagami Formation (stippled with decreasing shade) and the trajectories of iceberg tracks in northwestern Quebec and northeastern Ontario.

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Cochrane Till is found a few tens of kilometers within western Quebec (Veillette, 1989, 2007; Gao, 2012; Veillette et al., in press, a, b). Boreholes drilled in the Pinard Moraine revealed up to 70 m of sand (Smith, 1992). Surﬁcial geology mapping in Quebec and Ontario shows a gradual westward increase in the sand content of eskers that are the main sources of construction materials (Paradis et al., 2007; Veillette and Thibaudeau, 2007). The probable source for this abundant sand is the Lower Cretaceous unlithiﬁed deposits (silica sands, kaolin, silt and lignite) of the Mattagami Formation (Vos, 1975; Telford, 1982; Telford and Long, 1986) located a few kilometers north of the Pinard Moraine (Fig. 17). We propose that the ice stream was the main carrier of this sand eastward from the Mattagami Formation and, as such, the eastern limit of the ice stream may be further east than indicated by the distribution of relict ESE striations. Card and Sanford (1989) also identiﬁed an area with a thick drift cover extending up to 160 km east and SE of the Mattagami Formation. Scanning electron microscope analyses performed on quartz sand grains obtained from glacioﬂuvial deposits located to the SE of the Mattagami Formation suggest an ancient ﬂuvial origin for some of the grains, which were subsequently modiﬁed by glacial and glacioﬂuvial processes (Try et al., 1984; Telford et al., 1991; St-Jacques, 2011). 9. Discussion This discussion focuses on (1) the re-interpretation of the striation record of Area A, (2) the reasons LG1 was interpreted as an early Cochrane surge, (3) the problematic sequence of surges within the Cochrane episode, (4) the Winisk Ice Stream as a source of icebergs and a trigger for the Cochrane surge, and (5) a late Hudson lobe or ice streams in the James Bay basin? 9.1. The striation record of area A The conﬂicting interpretations of earlier workers on the relative ice-ﬂow chronology in Area A outlined in this study are explained by the failure to identify (1) relict striations left by older Wisconsinan ice ﬂows and (2) the impact of elevation and the vanished protector effect as the main factors causing abrupt inversions in the relative chronology of cross-striated surfaces between sites within the northern part of Area A. Both factors are intimately linked to the weak eroding competence of a thin, distal, semi-buoyant surging ice margin (LG2) controlled by topography and the level of Lake Ojibway that left distinct erosional marks within valleys and none or faint ones on the highest interﬂuves. Extrapolating cross-striated measurements obtained at speciﬁc sites within Area A to the whole of Area A by certain earlier workers led to conﬂicting interpretations on the ice-ﬂow chronology of the area. Locating, inspecting, uncovering, and measuring striated surfaces on outcrops eroded by two or more ice ﬂows are time-consuming ﬁeld activities that are unfortunately often conducted at a reconnaissance scale within surﬁcial geology mapping programs and regional studies due to time or logistical constraints. The obvious beneﬁts to be gained from detailed striation records are illustrated by several studies in speciﬁc regions of Canada (e.g. Clark, 1937; Lamarche, 1971; Lortie and Martineau, 1987; Rappol, 1993; Stea, 1994; Veillette, 1995, 1997; McMartin and Henderson, 2004; Paulen et al., 2013). These brought major corrections and additions to the ice-ﬂow chronology and regional glacial history. 9.2. Why was LG1 interpreted as an early Cochrane surge The new data raised important questions, notably the origin of LG1. Why was it presumed to be a Cochrane event (Cochrane I) by earlier workers? Have Cochrane I deposits been observed in earlier

151

studies and if so how do they differ from Cochrane II deposits? The ice-ﬂow model presented here is the coherent development of the ice-ﬂow chronology ﬁrst proposed by Boissonneau (1966). His assumption that Area A was part of a much larger ice-ﬂow system extending toward the north is validated by our data. It was based on cross-cutting relationships between the two late glacial ﬂows only, since he seemed to consider LG1 to be the only ice movement older than LG2 within Area A. Boissonneau's interpretation for the origin of Area A (Fig. 18) remained largely unexplored most likely because of the unorthodox explanation he proposed for Area A, which he interpreted as the distal part of his “Early Cochrane phase” that “wasted away” except for a large ice block that remained over it “… the ice block prevented Area A from being overridden by the late Cochrane readvance and consequently preserved the striae and drumlinoid landforms of the early Cochrane phase within Area A …. the force of the advancing late phase ice mass was almost spent when it reached the ice block and only a slight resistance of the ice block would have been sufﬁcient to stop the ice mass” (Boissonneau, 1966, p. 575). This interpretation presents Area A as a fossilised glacial landscape although there is no evidence for a large decaying ice mass at this location. This somewhat unusual attempt to explain a puzzling ice-ﬂow anomaly should not diminish his valuable contribution regarding the “buried” northern extension of his early Cochrane phase. Topography, as discussed earlier, is the reason why late Cochrane ice did not reach the southern part of Area A and why it penetrated further south to the east of it. Richard and Hilborn (1984a, b) recognised an older ice ﬂow (170 e190 ) within the northwestern part of Area A (NTS 42G13 and 42F16, Supplementary Fig. SD1) which they identiﬁed as “the main phase of the Late Wisconsinan glaciation” that was followed by a ﬁrst readvance toward 130 e140 , presumed to encroach into a late-stage Lake Ojibway (Cochrane I; Boissonneau's early Cochrane phase) and a last readvance toward 220 (Cochrane II; Boissonneau's late Cochrane phase) immediately thereafter. They reported widespread, bedded, glaciolacustrine ﬁne-grained deposits below 256 m but mainly overlying Cochrane Till, an indication of post-Cochrane advance sedimentation. Thus a minimum elevation of 256 m is inferred for the deposition of Lake Ojibway clay and Cochrane Till and no ﬁne-grained Lake Ojibway bedded sediments underlying Cochrane Till are reported. The incursion of Cochrane I ice into a late-stage Lake Ojibway was presumed but without any evidence that it overrode Lake Ojibway deposits. The carbonate-rich, poorly sorted, stony or pebbly clay diamict characteristic of late Cochrane surge deposits of northern Ontario and Quebec (Hughes, 1959; Paulen, 2001; Veillette et al., in press, a, b, c and d) was observed only in association with LG2, never with LG1. Similarly, Boissonneau (1966) did not associate a speciﬁc till with his Cochrane I ﬂow. Detailed description by Smith (1992) of several rotosonic boreholes in unconsolidated deposits drilled along a 120 km northsouth transect between the western part of the Pinard Moraine and the Kapuskasing area, all located within the extent of LG1, failed to provide information permitting to distinguish LG1 till from LG2 till. Sediments of Cochrane origin are described as one complex unit that, “… reﬂects incorporation of the persistent bed of ﬁnegrained glaciolacustrine sediment that underlies it” … and that … “Cochrane Till is only sporadically intersected in the cores, or if it had few larger clasts, was not distinguishable from the underlying massive clay that has been grouped with the Barlow-Ojibway Formation” … (Smith, 1992, p. 31). The till in the southern part of Area A mapped by Morris (2002a,b, c, d, NTS 42G01, 02, 03 and 04; Supplementary Fig. SD1) is everywhere described as a sandy, silty till with no indication of underlying ﬁne-grained deposits. Paulen tentatively interpreted large, regular and nearly symmetrical arcuate features open to the west, which rise less than

152

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Fig. 18. Location of Area A and ice front positions of the early (LG1) and late (LG2) Cochrane phases (from Boissonneau, 1966).

2e5 m above the surrounding terrain in the Oke Lake area (lake in northwest part of 42A13, Supplementary Fig. SD2), as frontal moraines marking ice-marginal positions of the latest Cochrane surge (LG2) of the area (Paulen, 2001, p. 104). The largest arcuate feature, about 4 km between the two extremities, is located to the north of the northern end of Oke Lake and wraps around the east end of a bedrock-controlled, roughly circular hill, reaching a maximum elevation around 315 m. The surface of the feature shows a nearly level surface at 302 m with variations in elevation of about 1 m over its entire length. The arcuate features were interpreted as raised shorelines by Boissonneau (1966, map S365), a likely valid interpretation given the elevation of the site, the close agreement with contour lines and the elevation of Lake Ojibway strandlines in the general area. LG2 streamlined landforms (toward 220 ) are absent in the immediate vicinity of Oke Lake because of the higher (>300 m) elevation but are widespread at lower elevation in the immediate vicinity (NTS maps 42A13 and 42G01, Supplementary Fig. SD2). If the features are truly moraines formed by the LG1 ﬂow, an unlikely interpretation given that LG1 streamlined terrain rarely exceeds 250 m, then the elevation of the site would explain why they have escaped destruction by LG2 and why Paulen (2001) interpreted them as the product of the latest Cochrane surge (Kapuskasing surge of Prest, 1970) toward the ESE (LG1). Given there is no mention in the literature of a Cochrane I till overlying ﬁne-grained Lake Ojibway deposits we believe that LG1 is not associated with a Cochrane surge. The physical and sedimentological characteristics of the deposits associated with LG1 remain to be determined. 9.3. The problematic sequence of surges within the Cochrane episode

Fig. 19. Stratigraphy of the Fauquier rock quarry section, modiﬁed from Paulen (2001).

Confusion aptly describes the chronology of Cochrane surges found in the literature. Following Antevs (1925), Hughes (1956,

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

1959) identiﬁed a single major late glacial Cochrane readvance into Lake Ojibway. Boisssonneau (1966) recognised an early and a late Cochrane phase labelled Cochrane I and II by Prest (1970) and accepted by others afterwards. Hardy (1977) identiﬁed three Cochrane surges in Quebec, Cochrane I, Cochrane II and Rupert but his Cochrane I limit corresponds to Prest's Cochrane II limit, a discrepancy also noted by Breckenridge et al. (2012). Smith (1992) recognised only one Cochrane surge overlying Lake Ojibway deposits. Veillette and co-workers recognised shifting ice ﬂows within a major Cochrane episode but did not ﬁnd stratigraphic evidence for two or more distinct surges, in spite of extensive surﬁcial geology mapping programs and regional studies carried out in Quebec (Veillette, 1989, 2007; Veillette and Thibaudeau, 2007; Veillette et al., 1991) and in Ontario (Veillette et al., in pressa, b, c and d). Paulen (2001) assigned a Cochrane age to late glacial ﬂuted terrain toward the ESE, the SW and the SE at the southernmost Cochrane limit in Ontario but attributed the formation of these ﬂows to a lobate margin. Thorleifson et al. (1993) rejected the interpretation by Prest (1969) that the ﬂuted belt north of Albany River is the western ﬂank of a large Cochrane lobe on the basis of (1) the improbability of ice ﬂow parallel to the postulated Cochrane ice margin (2) the substantial rise in elevation along the western margin of this belt, and (3) the presence of an esker that trends southward within the belt before shifting to a southwest trend parallel to ﬂutings west of the belt (Prest et al., 1968). Dyke (2004) attributed a Cochrane I origin to the Pinard Moraine. These apparent contradictions and inconsistencies are eliminated for the most part if LG1 is associated with an advancing terrestrial ice stream that ﬂowed toward the SE before the onset of the Cochrane surge. The Cochrane episode, characterized by a broad, fan-shape, southward (SW, S, SE) ﬂow with a semi-buoyant ice margin which overrode the northern fringe of glaciolacustrine deposits below 300 m elevation over a large part of the James Bay basin (Fig. 10), should thus be restricted to the very late stage of deglaciation. Several intersecting belts of ﬂuted terrain, both large and small, occur within this broad band of surging ice (Dredge and Cowan, 1989) but are largely buried in their northern parts by marine sediments and vast expanses of organic deposits. Attempts to identify the individual ﬂuted belts without proper stratigraphic control explain the confusion found in the literature. Examination of borehole logs, road cuts and some river sections failed to provide provenance and directional data in support of a lower unit (till?) deposited by ESE moving ice overlain by an upper unit (Cochrane Till) deposited by southward moving ice. Stratigraphic evidence is still lacking in support of LG1 with a possible exception, however, at the Fauquier rock quarry located about 20 km east of Kapuskasing, Paulen (2001) described a 6 m section showing from top to bottom, post-Cochrane glaciolacustrine ﬁne-grained deposits, diamicton, Cochrane Till, Matheson Till and bedrock (Fig. 19). The middle part of the section described as Cochrane Till, forms a two-layer unit with a sharp colour change at the upper contact of the lower layer, which is compact, ﬁssile, with a higher clast content; and a till fabric showing an east-west orientation. The base of the lower Cochrane layer shows a sharp planar contact with the underlying sandy Matheson Till with fabric suggesting southward ﬂow. Based on striations measured in the vicinity of the section, Paulen concluded that the sequence of ice ﬂows is (1) 220 , (2) 185 and (3) 120 . Most signiﬁcant to the current investigation is the east-west till fabric measured in the silty-sandy till deposit (below 3 m) resting on Matheson Till that suggests, along with the 120 striations measured nearby, a provenance from the west (LG1). Because the sequence of ice-ﬂow derived from cross-striated measured at Fauquier and at several other sites in the vicinity during this study

153

differs from that proposed by Paulen an alternative interpretation is suggested. Measurements obtained during this investigation, both at the quarry site and at ﬁve cross-striated sites in the surrounding area (SE part of NTS map 42G08 and SW corner of 42H05; Supplementary Fig. SD2), revealed dominant striated surfaces toward 200 e210 overprinted on others toward 110 e120 . Streamlined landforms from LG2 toward ~220 are also widespread in the low-elevation (<240 m) terrain of the Fauquier area and are clearly expressed on SRTM images (Supplementary Fig. SD2), suggesting that the last ﬂow in the area was toward the SW (220 ) and not toward the SE (120 ). The lower diamicton layer is here tentatively correlated with LG1 due to (1) the striation record, (2) the stratigraphic position of the east-west till fabric within the lower Cochrane layer, (3) the abrupt color change at the top of it, and (4) the sharp planar contact with the lowermost till. The coarse-grained lowermost diamict overlying Matheson Till/ Cochrane contact (below 3.50 m) also suggests that LG1 ice did not override ﬁne-grained glaciolacustrine deposits at this site. 9.4. The Winisk Ice Stream, a source of icebergs and a trigger for the Cochrane surges Conﬁned by topography to the north of the drainage divide, the eastward progress of the ice stream toward Lake Ojibway likely accelerated melting and mass wasting of the ice sheet. The calving front retreated to a last position somewhere in the vicinity of the Kenogami and Albany river junction until ﬁnal drainage of the lake (Fig. 17). The evidence for intense and late calving of the Winisk Ice Stream is a striking pattern of iceberg tracks indicating movement toward the ESE, that are everywhere overprinted on Cochrane ﬂutes toward the SW, S and SE over a large area of Quebec and Ontario (Fig. 17). The pattern results from the systematic mapping of over 25 000 iceberg tracks carried out in James Bay basin over the last decades through the Geological Survey of Canada surﬁcial geology mapping programs in Quebec (Veillette et al., 1991; Veillette and Paradis, 1996; Veillette, 2007; Veillette and Pomares, 2003; Veillette and Thibaudeau, 2007). Except for speciﬁc areas (Paulen and McClenaghan, 1998a and b), the mapping of iceberg tracks has not been systematic in northeastern Ontario and is currently being completed (Randour, 2015; Paradis et al., in press; Veillette et al., in press a,b, c; d). The tracks are distributed in a 125 km wide, east-west band that covers the distal margin of the Cochrane ﬂuted terrain (LG2) over a distance of about 600 km between the Matagami area in Quebec and the Hearst area in Ontario (Fig. 17). Pockets of stiff, brown, calcareous pebbly clay containing dropstones from Hudson Platform sedimentary rocks overlie deformation structures on the west ﬂanks and on the top of eskers that acted as barriers to the eastward movement of icebergs carrying debris in the Ojibway basin of northwestern Quebec (Veillette et al., 1991). Present-day sub-bottom lake deposits also provide sedimentological evidence for ice-rafted debris (IRD). Stroup et al. (2013) described a thin (25 cm) unit of clay-pebble conglomerate found in the upper portions of cores obtained from sub-bottom lake sediments in northern Ontario that contains dropstones consisting of solid rock clasts and pellets of unconsolidated sediment. Proglacial waters at the calving ice stream terminus had to be sufﬁciently deep (100 m or more?) for large icebergs to incise furrows up to 4e5 m into the soft lake bottom (Veillette and Paradis, 1996). Because the furrows (1) are still plainly visible on aerial photographs in spite of the vegetation cover, (2) are younger than the Cochrane surges, (3) have not been inﬁlled with glaciolacustrine deposits, and (4) given that the level of Lake Ojibway at the onset of the Cochrane surges reached a minimum elevation of 350 m (Veillette et al., 1991), it is proposed that this massive iceberg

154

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

discharge was the main event that preceded the ﬁnal collapse of the ice sheet and northward drainage of Lake Ojibway (Vincent and Hardy, 1979; Veillette, 1994; Barber et al., 1999). 9.5. Lobe or ice streams? Thorleifson et al. (1993) attributed the large ﬂuted belt deﬁning the Winisk Ice Stream to a late glacial readjustment in a disintegrating ice mass that occurred shortly prior to the Cochrane surges (Fig. 20). We also believe that the ﬂuted belt is indicative of fast-ﬂowing ice, which preceded the Cochrane surge, but we hesitate to restrict the ice stream to this belt for reasons outlined below. While the western boundary of the ice stream is deﬁned by topography (high ground), the eastern boundary is not. It coincides with the contact between the streamlined bedforms and younger marine and postglacial deposits located well below the marine limit (Fig. 16). No ﬁrm eastern boundary can be assigned to the ice stream since the ﬂuted terrain may extend eastward underneath a cover of younger sediments or was destroyed or modiﬁed by erosional processes, as suggested by the gap in the ﬂuted terrain and clearly expressed on satellite imagery (Fig. 16) in the upper watershed of the Attawapiskat River. A similar situation exists for the proposed southern component of the ice stream controlled to the south by higher ground but without a distinct boundary to the north, where it lies below the marine limit (Fig. 17). With no

distinct boundary toward James Bay both for its northern and southern components is this ice-ﬂow system (Winisk Ice stream) an ice stream sensu stricto? At this time we do not have the data to correctly answer this question and so retain the interpretation proposed by Thorleifson et al. (1993). Several indications, however, suggest that like Area A to the south, the northern component (ﬂuted belt) of the Winisk Ice stream is the “outcropping” part of a much larger ice-ﬂow system probably associated with a late Hudson lobe centered in the James Bay basin. This hypothesis implies that the Sutton Ridge (Fig. 16) split the lobe into a western component (ﬂuted belt) and an eastern one ﬂowing toward the depression now occupied by James Bay. Isolated striations on islands within northern James Bay and inland west of it indicate SSE late ﬂowing ice. These and numerous other striations on the eastern shore of James Bay indicative of a large converging ice ﬂow toward the center of the Bay were used to propose the James Bay Ice stream (Veillette, 1997). In the lobe hypothesis, the ﬂuted belt (SSE ﬂow) along with Area A and its palimpsest extension (ESE ﬂow) become, along with the striations and streamlined landforms of the eastern James Bay coast (SSW ﬂow), the end members of a late Hudson lobe that occupied most of the James Bay basin at elevation below 300 m. The central part of this vast depression (Fig. 20), below about 180 m west of James Bay and 220 m east of it, was ﬂooded by Tyrrell Sea and covered by extensive marine and postglacial deposits. Although the existence of the Winisk and the

Fig. 20. The Winisk Ice Stream (W) and its exposed portion (Area A) in the Hearst area; the solid red lines mark the limits of the ﬂuted belt and the dashed red line the limit of its palimpsest extension based mainly on the distribution of relict striations. Glacial Map from Prest et al. (1968). (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the web version of this article.)

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

James Bay ice streams is seriously challenged by the lobe hypothesis, we believe that further evidence is required to demonstrate that a late lobe adequately explains the ice-ﬂow data we have at this time. Consequently the Winisk and the James Bay ice streams are retained and with the Albany Bay Ice stream form the three ice streams of the James Bay basin (Fig. 21). Their relative chronology and impact on the deglaciation of the James Bay basin are brieﬂy described below. At an estimated age of 9300 yrs BP (Thorleifson and Kristjansson, 1993) the Albany Bay Ice Stream is the oldest of the three. The topographic low to the northwest of Hearst likely channelled the main trunk of the ice stream, which split into the Drowning and Kenogami river valleys (Fig. 21). It ﬂowed over the continental drainage divide into the Lake Superior basin carrying calcareous fragments and ﬁne-grained materials from the Hudson Platform that likely acted as lubricant for the fast-ﬂowing ice (Hicock, 1988; Hicock et al., 1989; Dredge, 2000; Windsborrow et al., 2010; Eyles and Putkinen, 2014). Belts of thick calcareous till in the Geraldton and Hemlo areas are ﬂanked by zones of slightly to non-calcareous thin till and bedrock (Geddes et al., 1985; Geddes and Kristjansson, 1986; Kristjansson, 1986; Thorleifson and Kristjansson, 1993). It shaped a striking lineated till plain on which glacial meltwater erosional features have been superimposed on the crystalline Precambrian substrate of the Geraldton area. About 60% of its 500 km path is located north of the Precambrian/

155

Paleozoic contact where it was presumed to be centered in the large Albany River Valley (Hicock, 1988) and it ﬂowed at right angle to the Winisk Ice stream in its northern portion. These cross-cutting patterns, dictated by topography contributed to a rapid thinning of the ice mass to the north, culminating in the Cochrane surge followed by a massive discharge of icebergs. Similar cross-cutting paleo-ice streams (corridors) were identiﬁed in the Saskatchewan Cofaigh et al., 2010). Because most of Prairies (Ross et al., 2009; O the James Bay Ice Stream was channelled in the depression now occupied by James Bay, evidence for a last converging southward ﬂow had to be sought from cross-striated surfaces found along the shorelines and on islands within the Bay (Veillette, 1997). The Archean granitoids of the eastern shore of James Bay preserved a convincing striation record of the counter-clockwise shift in ice ﬂows which turned out to be the only reliable method to demonstrate that the southward ﬂow had occurred last (see Fig. 7). Earlier workers using cross-cutting streamlined landforms on the eastern shore had concluded that the southward ﬂow was the oldest not the youngest (Lee et al., 1960; Hardy, 1976). Reliable ice-ﬂow indicators could not be found along the western shore underlain by ﬂat-lying Paleozoic sedimentary rocks covered by extensive mud ﬂats with few rock outcrops. The James Bay Ice Stream marks the onset of an abrupt shift in ice-ﬂow direction from the SW across James Bay to the SSE. It was short-lived and occurred shortly before the Cochrane surge (Hicock, 1988; Veillette, 1997).

Fig. 21. Intersecting ice stream paths of the James Bay basin. Ice ﬂow was toward the SW (large arrows) when the Albany Bay Ice Stream was formed.

156

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

10. Conclusions This investigation is based on small-scale satellite imagery analysis and on the interpretation of aerial photographs and surﬁcial geology maps supported by ice-directional ﬁeld measurements and observations. Jennings (2006) correctly pointed out that digital elevation models (DEMs), satellite imagery and all remote sensing techniques showing a continuous image of landforms at a continental scale, now allow glacial geologists mapping in small areas to place their observations within a much broader glacial context. The southern component of the Winisk Ice Stream, largely masked by younger sediment, was detected using relict striations since landform analysis could not be used. However, the continuous image of landforms clearly expressed on satellite images within Area A and in the northern component of the ice stream, made it possible to link the “buried” with the “visible” parts of the ice stream. The following methodological approach was followed to reach this interpretation. - The detailed study of cross-striated surfaces in conjunction with remote sensing methods clariﬁed the controversy surrounding the relative age of the ESE (130 ) and SW (220 ) ice ﬂows of the ice-ﬂow anomaly referred to as Area A (Boissonneau, 1966). The data presented here clearly indicate that the 130 ﬂow preceded the 220 ﬂow. - The conﬂicting interpretations on the relative age of the two ice ﬂows are attributed mainly to the sporadic occurrence of relict striated surfaces left by southwestward (220 - 240 ) Wisconsinan ice ﬂows that escaped destruction by late glacial ﬂows. The southwestward striations are older (Wisconsinan) at some sites and younger (late glacial 220 ) at others relative to the 130 ﬂow. Other factors affecting this apparent age reversal are the elevation of the cross-striated sites and the vanished protector effect. - Area A is interpreted as the only southernmost visible part of a large ice stream tentatively considered to be the southern extension of the Winisk Ice Stream (Thorleifson et al., 1993). The total length of the ice stream between its northern and southern extremities is about 800 km, twice the length of the continuous belt of streamlined landforms in its northern part, making it the longest terrestrial ice stream on the Canadian Shield. - The absence of ﬁne-grained glaciolacustrine deposits underneath the ﬂuted terrain of silty and sandy till in the southern part of Area A (Morris, 2002a;b, c, d), and the stratigraphy of the Fauquier section suggest that the ice stream did not override the Ojibway lake bed, at least in these parts of its extent. - The use of Cochrane I and Cochrane II to designate two major late glacial readvances should be abandoned since Cochrane I is the southern component of the Winisk Ice Stream (or the exposed western part of a late Hudson lobe). The Cochrane event consists of one major, short-lived, fan-shape surge comprising several criss-crossing, large and small, ﬂuted belts. - While the striation record clearly shows that the Winisk Ice Stream (LG1) preceded the Cochrane surge (LG2), a vast network of iceberg tracks showing a consistent direction of movement toward the ESE from a presumed ice stream terminus in the Albany River area cross-cut Cochrane ﬂutings over a large area of Quebec and Ontario, suggesting that the ice stream not only preceded, but also outlasted the Cochrane surge. - The late deglaciation of the James Bay basin was accelerated by the interaction of three successive major ice streams controlled by topography.

- The discovery of the southern extension of the Winisk Ice Stream that ﬂowed in an unreported direction (ESE) has obvious implications for mineral prospecting methods applicable in glaciated terrain.

11. Current and future work The distribution of ESE relict striations is at this moment the only reliable indicator available to associate LG1 with the southern extension of the Winisk Ice Stream. Stratigraphic information is required to deﬁne the eastern and northeastern limits of the ice stream or to verify the hypothesis that the Winisk and James Bay ice streams are the end members of a late Hudson lobe as discussed in Section 9.5. The stratigraphic equivalent of the Winisk Till (Thorleifson et al., 1993) should be present in the upper parts of sections along the main rivers (Missinaibi, Nagakami, Kenogami, Mattagami), which ﬂow roughly at right angle to the ice stream path (Fig. 17). The description of streamlined landforms trending toward the ESE exposed in river sections north of Hearst, Ontario, suggests an association with the southern component of the ice stream (G. Gao, Ontario Geological Survey, pers. communication, August 25, 2015). The southern extension of the Winisk Ice Stream raises questions regarding the origin of the Pinard Moraine. The east-west orientation of the moraine is the primary reason it was interpreted as a recessional deposit (Boissonneau, 1966), but this linear alignment along the path of the ice stream and the dispersal of sand toward the ESE from the nearby Mattagami Formation, the probable source of the large volume of sand in the moraine, suggest that formation of the moraine may be related to the ice stream (Fig. 17). The steeper, southern (down-ice) face of the moraine is visually evocative of an up-ice face and the northern (up-ice), gently sloping ﬂank is not typical of a grounded subaquatic continental land system (Fyfe, 1990; Teller, 2005). The movement of icebergs in the northern part of the Ojibway basin was attributed to prevailing north-westerly winds driving icebergs from the calving margins of Cochrane surges (Dionne, 1977; Veillette and Paradis, 1996). The extensive mapping of iceberg tracks described earlier provided data that are currently being analysed for a large area of eastern Ontario and western Quebec. A Winisk Ice Stream terminus calving in Lake Ojibway from a position in the vicinity of the Albany and Kenogami river junction (Fig. 17) is a likely source of icebergs driven east by katabatic winds from a much-reduced Laurentide Ice Sheet (Wolfe et al., 2004). Acknowledgments An important part of the work described here was done within a surﬁcial geology mapping project which was part of the Abitibi Targeted Geoscience Initiative program (S.J. Paradis) involving the Geological Survey of Canada (GSC), the Ontario Geological Survey re des ressources naturelles et de la Faune du (OGS) and the Ministe bec (MRNFQ). M.B. McClenaghan (GSC), R.D. Thomas (Thomas Que and Associates) and G. Gao (OGS), discussed their results and provided unpublished information for various locations within the (GSC) and Michelle general area. Ruth Boivin and Nathalie Cote Laithier (UQAM) drafted the ﬁgures. The comments of Isabelle McMartin (GSC) who revised the manuscript internally and those of Carrie Jennings and an anonymous reviewer helped to improve the manuscript. Personnel of the Cochrane District of the Ministry of Natural Resources of Ontario, provided documents and valuable information on access to speciﬁc areas by logging roads and waterways. Sincere thanks are extended to all. Natural Resources Canada, Contribution number ID 28536-ESS.

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Appendix A. Supplementary data Supplementary data related to this article can be found at http:// dx.doi.org/10.1016/j.quascirev.2016.11.008. References Antevs, E., 1925. Retreat of the Last Ice Sheet in Eastern Canada. Memoir 146. Geological Survey of Canada, Ottawa, Ontario, 142 p. Barber, D.C., Dyke, A., Hillaire-Marcel, C., Jennings, A.E., Andrews, J.T., Kervin, M.W., Bilodeau, G., McNeely, R., Southon, G., Moorhead, M.D., Gagnon, J.-M., 1999. Forcing of the cold event of 8200 years ago by catastrophic drainage of Laurentide lakes. Nature 400, 344e348. Boissonneau, A.N., 1966. Glacial history of northeastern Ontario 1. The CochraneHearst area. Can. J. Earth Sci. 3 (no. 5), 559e578. Bostock, H.S., 2014. Physiographic Regions of Canada. Geological Survey of Canada. “A” series Map 1254A, 1: 5 000 000. Boulton, G.S., Clark, C.D., 1990. A highly mobile Laurentide Ice Sheet revealed by satellite images of glacial lineations. Nature 346, 813e817. Breckenridge, A., Lowell, T.V., Stroup, J.S., Evans, G., 2012. A review and analysis of varve thickness records from glacial Lake Ojibway (Ontario and Quebec, Canada). Quat. Int. 260, 43e54. Card, K.D., Sanford, B.V., 1989. (compilers). Geological Survey of Canada, Geological Atlas. Map NW-17-G: 1 000 000, Sheet 1 of 5, The National Earth Science Series. Clark, T.H., 1937. Northward moving ice in southern Quebec. Am. J. Sci. 34, 215e220, 5th series. Clark, P.U., 1991. Striated clast pavements: products of deforming subglacial sediments? Geology 19 (5), 530e533. Clark, C.D., Knight, J.K., Gray, J.T., 2000. Geomorphological reconstruction of the Labrador sector of the Laurentide ice sheet. Quat. Sci. Rev. 19 (13), 1343e1366. Cofaigh, C., Stokes, C.R., 2008. Reconstructing ice-sheet dynamics from subglacial O sediments and landforms: introduction and overview. Earth Surf. Process. Landforms 33, 495e502. De Angelis, H., Kleman, J., 2008. Paleo-ice streams onsets: examples from the northeastern Laurentide Ice Sheet. Earth Surf. Process. Landforms 33, 560e572. Dionne, J.-C., 1977. Relict iceberg furrows on the ﬂoor of glacial Lake Ojibway, Quebec and Ontario. Marit. Sediments 13 (no 2), 79e81. Dredge, L.A., 2000. Carbonate dispersal trains, secondary till plumes, and ice streams in the west Foxe Sector, Laurentide Ice Sheet. Boreas 29, 144e156. Dredge, L.A., Cowan, W.R., 1989. Quaternary geology of the southwestern Canadian Shield. In: Fulton, R.J. (Ed.), Quaternary Geology of Canada and Greenland. Geological Survey of Canada, Geology of Canada, pp. 214e249 no.1. -Loubert, H., Roy, M., Allard, G., Lamothe, M., Veillette, J.J., 2013. Glacial and Dube nonglacial events in the eastern James Bay Lowlands, Canada. Can. J. Earth Sci. 50, 379e396. Dyke, A.S., 2004. An outline of North American deglaciation with emphasis on central and northern Canada. In: Ehlers, J., Gibbard, P.L. (Eds.), Quaternary Glaciations - Extent and Chronology, Part II. Elsevier, North America, pp. 373e424. New York. Eyles, N., Putkinen, N., 2014. Glacially-mega-lineated limestone terrain of Anticosti Island, Gulf of St-Lawrence, Canada: onset zone of the Laurentian channel ice stream. Quat. Sci. Rev. 88, 125e134. Fyfe, G.J., 1990. The effect of water depth on ice-proximal glaciolacustrine sedi€ Island, southern Finland. Boreas 19, 147e164. mentation: Salpausseka Gao, C., 2012. Surﬁcial mapping in the Detour lake eBurntbush area, project unit 08-005: a progress report. In: Summary of Field Work and Other Activities. Ontario Geological Survey. Open File Report 6280, 23e1 to 23-10. Gao, C., 2013a. Quaternary Geology, Burntbush Area, Northeastern Ontario. Ontario Geological Survey. Preliminary Map P. 3770, 1: 100 000. Gao, C., 2013b. Quaternary Geology, Detour Lake Area, Northeastern Ontario. Ontario Geological Survey. Preliminary Map, P.3769, scale 1:50 000. Geddes, R.S., Kristjansson, F.J., 1986. Quaternary geology of the Hemlo area: constraints on mineral exploration. Canadian Geology. J. Can. Inst. Min. Metall. 1, 5e8. District of Thunder Bay. Ontario Geological Survey, Miscellaneous Paper 126, 151-154. Geddes, R.S., Bajc, A.F., Kristjansson, F.J., 1985. Quaternary Geology of the Hemlo Region. Ontario Geological Survey, District of Thunder Bay, pp. 151e154. Miscellaneous Paper 126. tude ge omorphologique de la portion que be coise Hardy, L., 1976. Contribution a l’e des basses terres de la baie James. PhD thesis. Department of Geography, McGill University, Montreal, Quebec, 264p. glaciation et les e pisodes lacustres et marins sur le versant Hardy, L., 1977. La de be cois des basses terres de la baie James. Ge ogr. Physique Quat. 31, que 261e273. rieur Hardy, L., 1982. Le Wisconsinien supe a l’est de la baie James (Quebec). Le. Nat. Can. 109, 333e351. Hicock, S.R., 1988. Calcareous till facies north of Lake Superior, Ontario: implications ogr. Physique Quat. 42 (no. 2), 120e135. for Laurentide ice streaming. Ge Hicock, S.R., Kristjansson, F.J., Sharp, D.R., 1989. Carbonate till as a soft bed for pleistocene ice streams on the Canadian Shield north of Lake Superior. Can. J. Earth Sci. 22, 2249e2254. Hughes, O.L., 1956. Surﬁcial Geology of Smooth Rock, Cochrane District, Ontario (Preliminary Report and Map). Department of Mines and Technical Surveys,

157

Geological Survey of Canada, 9 p., Map 55-41, scale 1:126 720. Hughes, O.L., 1959. Surﬁcial Geology of the Smooth Rock Falls Map Area, Cochrane District, Ontario. unpublished Ph.D dissertation. University of Kansas, Lawrence, Kansas, 190 p. Jackson, S.I., Fyon, J.A., 1991. The Western Abitibi Subprovince. Ont. Geol. Surv. 4, 405e484. Part 1. Jansson, K., Stroeven, A., Kleman, J., 2003. Conﬁguration and timing of Ungava Bay ice streams, Labrador-Ungava, Canada. Boreas 32 (1), 256e262. Jennings, Carrie E., 2006. Terrestrial ice streams e a view from the lobe. Geomorphology 75, 100e124. Kleman, J., Jansson, K., De Angelis, H., Stroeven, A.P., Hatterand, C., Alm, G., Glasser, N., 2010. North American Ice Sheet build-up during the last glacial cycle, 115-21 kyr. Quat. Sci. Rev. 29 (17e18), 2036e2051. Kristjansson, F.J., 1986. Quaternary Geology of the Beardmore-geraldton Area, District of Thunder Bay. Ontario Geological Survey. Miscellaneous Paper 132, 172e175. Lamarche, R.Y., 1971. Northward moving ice in the Thetford Mines area of southern Quebec. Am. J. Sci. 271, 383e388. Lee, H.A., Eade, K.E., Heywood, W.W., 1960. Surﬁcial Geology, Sakami Lake (FortGeorge- Great Whale Area). Geological Survey of Canada, New Quebec. Map 521959, 1: 506880. Lian, O.B., Hicock, S.R., Dreimanis, A., 2003. Laurentide and Cordilleran fast ﬂow: some sedimentological evidence from Wisconsinan subglacial till and its substrate. Boreas 32, 102e113. mes de Stries Glaciaires dans les Appalaches Lortie, G., Martineau, G., 1987. Les Syste bec. Ministe re de l'Energie bec. DV 85 10, 45 p. du Que des Ressources du Que Low, A.P., 1889. Report on Exploration in James Bay, and Country East of Hudson Bay Drained by the Big, Great Whale and Clear Water Rivers, 1887e1888. Geological and Natural History Survey of Canada. Annual Report, v. 3, Part II, Report J, 62 p. Low, A. P., 1900. Geological survey of Canada, annual report, v. XIII, p. 84D. Margold, Martin, Stokes, Chris R., Clark, Chris D., 2015. Ice streams in the Laurentide Ice Sheet: identiﬁcation, characteristics and comparison to modern ice sheets. Earth Sci. Rev. 143, 117e146. McClenaghan, M.B., DiLabio, R.N.W., 1995. Overburden Database Compilation. Geological Survey of Canada, Timmins, Ontario (NTS 42A11, 12, 13, 14) Open File report 3086. McMartin, I., Henderson, P., 2004. Evidence from Keewatin (Central Nunavut) for ogr. Physique Quat. 58, 163e186. paleo-ice ice divide migration. Ge Morris, T.F., 2002a. Quaternary Geology, Roche Lake Area. Ontario Geological Survey. Map 2672, scale 1: 50 000. Morris, T.F., 2002b. Quaternary Geology, Opasatika Lake Area. Ontario Geological Survey. Map 2673, scale 1: 50 000. Morris, T.F., 2002c. Quaternary Geology, Woman Falls Area. Ontario Geological Survey. Map 2674, scale 1: 50 000. Morris, T.F., 2002d. Quaternary Geology, Wakusimi River Area. Ontario Geological Survey. Map 2675, scale 1: 50 000. Cofaigh, C., Evans, D.J.A., Smith, I.R., 2010. Large-scale reorganization and sediO mentation of terrestrial ice streams during Late Wisconsinan Laurentide Ice Sheet deglaciation. Geol. Soc. Am. Bull. 122, 743e756. Paradis, S.J., Veillette, J.J., Roy, M., St-Jacques, G., 2007. The trans-frontier QuebecOntario Abitibi surﬁcial geology map compilation project; Smooth Rock Falls e Cochrane area, NTS 42H. In: Ayers, et al. (Eds.), The Abitibi Greenstone Belt: Update of the Precambrian Geoscience Section Program, the Targeted Geoscience Initiative III Abitibi and Deep Search Projects. Ontario Geological Survey. Open File Report, 3e37-3-38. Paradis S.J., Veillette J.J., Bajc A.F., Paulen R.C., McClenaghan M.B., Surﬁcial Geology (in press). Surﬁcial Geology, Timmins, Ontario-Quebec. Geological Survey of Canda, Geoscience Map xxxx, scale 1: 250 000. Paulen, R.C., 2001. Quaternary Geology of the Timmins Area, Northeastern Ontario. M.Sc. thesis, Earth Sciences. University of Waterloo, Ontario, Canada, 448 p. (includes two surﬁcial geology maps, NTS 42A13, 42A14, 1: 50 000 scale. Paulen, R.C., McClenaghan, M.B., 1997. Surﬁcial Mapping and Quaternary Stratigraphic Studies in the Western Abitibi Greenstone Belt. Geological Survey of Canada, Timmins, Ontario. Current Research 1997-C, 191e200. Paulen, R.C., McClenaghan, M.B., 1998a. Surﬁcial Geology of the Buskegau River Area, Northeastern Ontario (NTS 42A14). Geological Survey of Canada. Open File 3619, 1: 50 000. Paulen, R.C., McClenaghan, M.B., 1998b. Surﬁcial Geology of the Manning Lake Area, Northeastern Ontario (NTS 42A13). Geological Survey of Canada. Open File 3618, 1: 50 000. Paulen, R.C., McClenaghan, M.B., 2015. Late Wisconsin ice-ﬂow history in the Buffalo Head Hills kimberlite ﬁeld, north-central Alberta. Can. J. Earth Sci. 52, 51e67. Paulen, R.C., McClenaghan, M.B., Hicken, A.K., 2013. Regional and local ice-ﬂow history in the vicinity of the Izok Lake Zn-Cu-Pb-Ag deposit, Nunavut. Can. J. Earth Sci. 50, 1209e1222. Percival, J.A., Card, K.D., 1985. Structure and evolution of Archean crust in central superior province, Canada. In: Ayer, J.D., Thurston, P.C., Card, K.D., Weber, W. (Eds.), Evolution of Archean Supracrustal Sequences. Geological Association of Canada, pp. 179e192. Special Paper 28. Prest, V.K., 1966. Glacial Geology, Albany River Sheet (NTS 42), 1: 1 000 000. Part of the Original Unpublished Manuscript with the Glacial Map of Canada (1968). Geological Survey of Canada Files, Ottawa, Canada. Prest, V.K., 1969. Retreat of Wisconsin and Recent Ice in North America. Geological Survey of Canada. Map 1257A.

158

J.J. Veillette et al. / Quaternary Science Reviews 155 (2017) 136e158

Prest, V.K., 1970. Quaternary geology of Canada; in geology and economic minerals of Canada. In: Douglas, R.J.W. (Ed.), Geological Survey of Canada Economic Geology Series, No. 1, ﬁfth ed., pp. 676e758 Ottawa. Prest, V.K., Grant, D.R., Rampton, V.N., 1968. Glacial Map of Canada. Geological Survey of Canada. Map 1253A. Punkari, M., 1982. Glacial geomorphology and dynamics in the eastern parts of the Baltic Shield interpreted using Landsat imagery. Photogrammetric J. Finl. 9 (198), 77e93. proximite de la frontie re Que bec-Ontario, 49 Randour, I., 2015. Sillons d’icebergs a partement des 50 Lat. N et 77, 78 , 79 , 82 , 83 et 84 Long W. B.Sc. thesis. De re, Universite du Que bec al, 12 p. Sciences de la Terre et de l’atmosphe a Montre Rappol, M., 1993. Ice Flow and Glacial Transport in Lower St. Lawrence. Geological Survey of Canada, Quebec. Paper 90-19, 28 p. Richard, J.A., Hilborn, L., 1984a. Quaternary Geology of the Constance Lake Area, Cochrane District. Ontario Geological Survey. Map P2695, Geological SeriesPreliminary Map, scale 1:50 000. Richard, J.A., Hilborn, L., 1984b. Quaternary Geology of the Hanlan Lake Area, Cochrane District. Ontario Geological Survey. Map P2696, Geological Series Preliminary Map, scale 1:50 000. Ross, M., Campbell, J.E., Parent, M., Adams, R.S., 2009. Paleo-ice streams and the subglacial mosaic of the North American mid-continental prairies. Boreas 38, 421e439. Roy, M., 1998. Pleistocene Stratigraphy of the Lower Nelson River Area: Implications for the Evolution of the Hudson Bay Lowland, Manitoba, Canada. M.Sc. partement des sciences de la terre, Universite du Que bec a Montre al, 220 p. De lie, J.-F., Parent, M., 2011. Insights Roy, M., Dell’Oste, F., Veillette, J.J., de Vernal, A., He on the events surrounding the ﬁnal drainage of Lake Ojibway based on James Bay stratigraphic sequences. Quat. Sci. Rev. 30, 682e692. Smith, S.L., 1992. Quaternary Stratigraphic Drilling Transect, Timmins to the Moose River Basin. Geological Survey of Canada, Ontario. Bulletin 415. 94 p. Smith, M.J., Knight, J., 2011. Paleoglaciology of the last Irish ice sheet reconstructed from striae evidence. Quat. Sci. Rev. 30, 147e160. po ^ts de surface et dynamique glaciaire du St-Jacques, G., 2011. Cartographie des de du secteur Lawagamau Lake, Ontario. M.Sc. Sciences de la Terre, Universite bec a Montre al, 139 p. Que Stea, R.R., 1994. Relict and palimpsest glacial landforms in Nova Scotia, Canada. In: Warren, W.P., Groots, D.C. (Eds.), Proceedings of the Commission on the Formation and Deformation of Glacial Deposits. A.A. Balkema, Rotterdam, pp. 141e158. Dublin, Ireland, May 1991. Stokes, C.R., Clark, C.D., 2001. Paleo-ice streams. Quat. Sci. Rev. 20 (13), 1437e1457. Stroup, J.S., Lowell, T.V., Breckenridge, A., 2013. A model for the demise of large, glacial lake Ojibway, Ontario and Quebec. J. Paleolimnol. 50, 105e121. Telford, P.G., 1982. Mesozoic stratigraphy of the Moose river basin. In: Telford, P.G., Verma, H.M. (Eds.), Mesozoic Geology and Mineral Potential of the Moose River Basin. Ontario Geological Survey. Study 21, 193 p. Telford, P.G., Long, D.G.P., 1986. Mesozoic geology of the Hudson Platform. In: Martini, I.P. (Ed.), Canadian Inland Seas, pp. 43e54. Elsivier Oceanography Series, no. 44, p. Telford, P.G., Long, D.F., Norris, G., Zippi, R., Grifﬁs, R., 1991. Mesozoic Geology and Lignite Potential of the Moose River Basin. Ontario Geological Survey. Open File Report 5777, 185 p. Teller, J.T., 2005. Subaquatic landsystems: large proglacial lakes. In: Evans, D.J.A. (Ed.), Glacial Landsystems. Hodder Arnold, London, pp. 348e370. Thorleifson, L.H., Kristjansson, F.J., 1993. Quaternary Geology and Drift Prospecting, Beardmore-geraldton Area. Geological Survey of Canada, Ontario. Memoir 435,146 p. includes one surﬁcial geology map 1: 100 000. Thorleifson, L.H., Wyatt, P.H., Warman, T.A., 1993. Quaternary stratigraphy of the severn and Winisk drainage basins, northern Ontario. Geol. Surv. Can. Bull. 442, 59 p. Trommelen, M.S., Ross, M., 2011. Far North Geomapping Initiative: Palimpsest Bedrock Macroforms and Other Complex Ice-ﬂow Indicators Near Churchill, Northern Manitoba (Part of NTS 54L16); in Report of Activities 2011, Manitoba Innovation, Energy and Mines. Manitoba Geological Survey, pp. 29e35. Try, C.F., Long, D.J.F., Winder, C.G., 1984. Sedimentology of the lower Cretaceous Mattagami Formation, Moose river basin, James Bay Lowlands, Ontario, Canada. In: Stout, D.F., Glass, D.J. (Eds.), The Mesozoic of Middle North America. Canadian Society of Petroleum Geologists, pp. 345e359. Memoir 9. Veillette, J.J., 1989. Ice movements, till sheets and glacial transport in Abitibi-

Timiskaming, Quebec and Ontario. In: DiLabio, R.N.W., Coker, W.B. (Eds.), Drift Prospecting. Geological Survey of Canada, pp. 139e154. Paper 89-20. Veillette, J.J., 1994. Evolution and paleohydrology of glacial lakes Barlow and Ojibway. Quat. Sci. Rev. 13, 945e971. Veillette, J.J., 1995. New evidence for northwestward glacial ice ﬂow, James Bay region, Quebec. In: Current Research 1995-C. Geological Survey of Canada, pp. 249e258. ^le d’un courant de glace tardif dans la de glaciation de la baie Veillette, J.J., 1997. Le ro ogr. Physique Quat. 51 (2), 141e161. James. Ge ologie des formations en surface et histoire glaciaire. Veillette, J.J., 2007. Ge re Harricana, Que bec. Map 1993A, 1:100 000. Geological Survey of Canada, Rivie Veillette, J.J., McClenaghan, M.B., 1996. Sequence of Glacial Ice Flows in Abitibitimiskaming: Implications for Mineral Exploration and Dispersal of Calcareous Rocks from the Hudson Bay Basin, Quebec and Ontario. Geological Survey of Canada. Open File 3033. Veillette, J.J., Paradis, S.J., 1996. Iceberg Furrows as Paleowind Indicators and Icebergs as Erosion and Sedimentation Agents in Lake Ojibway. Geological Survey of Canada, Quebec. Open File 3031. ologie des formations en surface et histoire Veillette, J.J., Pomares, J.-S., 2003. Ge bec. Map 1994A, 1: glaciaire. Geological Survey of Canada, Lac Matagami, Que 100 000. Veillette, J.J., Roy, M., 1995. The spectacular cross-striated outcrops of James Bay, Quebec. In: Current Research 1995-C. Geological Survey of Canada, pp. 243e248. ologie des formations en surface et histoire Veillette, J.J., Thibaudeau, P., 2007. Ge re Wawagosic, Que bec. Map 1995A, glaciaire. Geological Survey of Canada, Rivie 1: 100 000. Veillette, J.J., Paradis, S.J., Thibaudeau, P., Pomares, S.J., 1991. Distribution of Hudson Bay Erratics and the Problem of the Cochrane Limit in Abitibi. Geological Survey of Canada, Current Research, Quebec, pp. 135e142. Paper 91e1C. Veillette, J.J., Dyke, A.S., Roy, M., 1999. Ice-ﬂow evolution of the Labrador sector of the Laurentide ice sheet: a review, with new evidence from northern Quebec. Quat. Sci. Rev. 18 (8e9), 993e1019. Veillette, J.J., St-Jacques, G., Roy, M., Paradis, S.J., Boivin, R., 2008. Follow-up on the Trans-frontier Quebec-Ontario Abitibi Surﬁcial Geology Map Compilation Project: Smooth Rock Falls-Cochrane Area (NTS 42H). Summary of Field Work and Other Activities. Ontario Geological Survey. Open File Report 6226, 13-1 to 13e8. Veillette J.J., Saint-Jacques G., Paradis S.J., Surﬁcial geology, Fraserdale, Ontarioologie des formations superﬁcielles, surﬁcial geology, (in press-a), Quebec/Ge Geological Survey of Canada. http://dx.doi.org/10.4095/000000, Geoscience Map xxxx, scale 1:100 000 Veillette J.J., Saint-Jacques G., Paradis S.J., Surﬁcial geology, Frederick House River, ologie des formations superﬁcielles, surﬁcial geology, (in Ontario-Quebec/Ge press-b), Geological Survey of Canada. http://dx.doi.org/10.4095/000000, Geoscience Map xxxx, scale 1:100 000. Veillette J.J., Saint-Jacques G., Paradis S.J., Surﬁcial geology, Little Abitibi Lake, ologie des formations superﬁcielles, surﬁcial geology, (in Ontario-Quebec/Ge press-c), Geological Survey of Canada. http://dx.doi.org/10.4095/000000, Geoscience Map xxxx, scale 1:100 000. Veillette J.J., Saint-Jacques G., Paradis S.J., Surﬁcial geology, North Burntbush Lake, ologie des formations superﬁcielles, surﬁcial geology, (in Ontario-Quebec/Ge press-d), Geological Survey of Canada. http://dx.doi.org/10.4095/000000, Geoscience Map xxxx, scale 1:100 000. Vincent, J.S., Hardy, L., 1979. The evolution of glacial lakes Barlow and Ojibway, Quebec and Ontario. Geol. Surv. Can. Bull. 316, 18 p. Vos, M.A., 1975. Economic Geology of the Cretaceous Deposits. General Appraisal, Ontario Division of Mines, Moose River basin, Ontario. Open File Report 5157, 70 p. Wilson, W.J., 1904. The Nanagami River and Other Branches of the Kenogami. Canadian Department of Mines and Technical Surveys. Geological Survey of Canada. Summary Report for 1903, 109a. Windsborrow, M.C.M., Clark, C.D., Stokes, C.R., 2010. What controls the location of ice streams? Earth Sci. Rev. 103, 45e59. Wolfe, S.A., Huntley, D.J., Ollerhead, J., 2004. Relict late Wisconsinan dune ﬁelds of the northern great plains. In glacial history, paleogeography, and paleoenvirogr. Physique Quat. 58, 323e336. onments in glaciated north America. Ge

Uncovering the hidden part of a large ice stream of the Laurentide Ice Sheet, northern Ontario, Canada

Uncovering the hidden part of a large ice stream of the Laurentide Ice Sheet, northern Ontario, Canada

Recommend Documents