Talus slope characterization in Tasiapik Valley (subarctic
Québec): Evidence of past and present slope processes
Samuel Veilleux, Najat Bhiry, Armelle Decaulne
To cite this version:
Samuel Veilleux, Najat Bhiry, Armelle Decaulne. Talus slope characterization in Tasiapik Valley
(subarctic Québec): Evidence of past and present slope processes. Geomorphology, Elsevier, 2020,
349, pp.106911. 10.1016/j.geomorph.2019.106911. hal-03170489
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Geomorphology 349 (2020) 106911
Contents lists available at ScienceDirect
Geomorphology
journal homepage: www.elsevier.com/locate/geomorph
Talus slope characterization in Tasiapik Valley (subarctic Québec):
Evidence of past and present slope processes
Samuel Veilleux a,b,∗ , Najat Bhiry a,b , Armelle Decaulne c
a
b
c
Département de géographie, Université Laval, Québec, Canada
Centre d’études nordiques, Université Laval, Québec, Canada
CNRS, Laboratoire LETG, Université de Nantes, LabEx DRIIHM, France
a r t i c l e
i n f o
Article history:
Received 6 June 2019
Received in revised form 14 October 2019
Accepted 14 October 2019
Available online 18 October 2019
Keywords:
Morphometry
Slope dynamics
Snow avalanches
Periglacial
Nunavik
a b s t r a c t
Topographic, granulometric, morphometric, petrographic and vegetation surveys were conducted on the
slopes of Tasiapik Valley, near Umiujaq (Nunavik), to document mass wasting processes and their geomorphological impact. Talus slopes, widespread at the foot of the steep rockwalls of Tasiapik Valley, are
an important landscape feature in the area. The lithology of the slope deposits attest their local origin,
namely the result of rockfalls coming from the adjacent wall. Locally, poor vegetation covering the clasts
exhibits recently fallen debris; elsewhere, dense shrub cover has colonized the slopes demonstrating
the low activity nowadays. On-going periglacial processes have led to extensive dismantling of the rockface, enabling for debris supply. Following the last deglaciation, paraglacial processes have potentially
favoured slope instabilities. The use of automatic cameras during the winter 2017–2018 resulted in the
observation of many snow-avalanche events; however few rockfall events have been observed. Spring
snow avalanches have carried rock debris to the talus at the foot of the slope; snow also enabled debris
redistribution on the slopes.
© 2019 Elsevier B.V. All rights reserved.
1. Introduction
Northern landscapes have undergone many changes since
their deglaciation. In particular, paraglacial conditions (e.g. glacioisostatic rebound and rockface dismantlement) induced talus slope
formation by supplying debris through the pressure release on rock
fractures and freeze-thaw processes (Ballantyne and Benn, 1994;
Matsuoka and Sakai, 1999; Ballantyne, 2002; Matsuoka, 2008).
Nunavik is part of the low-Arctic region of eastern Canada and its
landscape consists of low hills, basins and plateaus. The few studies that have been conducted in this vast region have demonstrated
the occurrence of slope processes on slopes less than 100 m high
(Belzile, 1984; Bégin and Filion, 1985; St-Cyr, 1986; Marion et al.,
1995; Germain and Martin, 2012; Germain, 2016). Recent studies
(Decaulne et al., 2018; Bhiry et al., 2019) conducted at Wiyâshâkimî
Lake in Tursujuq National Park (Nunavik) showed that talus slope
formation started after deglaciation at about 4600 BP, and that slope
processes are still active today. Some of the villages in Nunavik
(Salluit, Kangiqsujuaq and Kangiqsualujjuaq) are located within
∗ Corresponding author at: 2405, rue de la Terrasse, Université Laval Québec,
Québec G1V 0A6, Canada.
E-mail address: samuel.veilleux.4@ulaval.ca (S. Veilleux).
https://doi.org/10.1016/j.geomorph.2019.106911
0169-555X/© 2019 Elsevier B.V. All rights reserved.
glacial valleys with prominent slopes, while other villages (Umiujaq) are situated near high cuesta relief (∼230 m). Accordingly,
it is crucial to document slope dynamics and to evaluate associated risks on the local population, visitors and infrastructures. For
instance, in Kangiqsualujjuaq (northeastern Nunavik), a dreadful
snow avalanche hit the gymnasium of Satuumavik school during
the 1999 New Year’s Eve celebrations, causing the death of 9 people
and injuring 25 (Bérubé, 2000; Lied and Domaas, 2000; Germain,
2016). However, no extensive research has been conducted on
slope processes in the Umiujaq area (including snow avalanches,
landslide and rockfalls), their triggering factors, their occurrence
and their runout distance. General conditions are conducive for
bedrock dismantling and mass wasting, even with limited slope
heights, but additional knowledge about slope processes is still
required.
The main objective of the study is to document landforms organisation built by gravitational processes in Tasiapik Valley and their
contribution to talus development based on geomorphological surveys. This study discusses slope evolution during the Holocene,
from the retreat of the Laurentide Ice Sheet in the area to the
present-day, highlighting the potential risk at the valley bottom.
2
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
Fig. 1. Location of Tasiapik Valley within the Umiujaq area (A); regional geology and quaternary sediments in Tasiapik Valley (B); distribution of the talus slopes and
investigated slopes (C). Sources of background images: MRNF (A, B), UMI orthomosaic (2010) (C).
1.1. Regional setting
Tasiapik Valley (56◦ 33′ N, 76◦ 28′ W) is located 5 km east of the
Inuit village of Umiujaq, on the east coast of Hudson Bay in Nunavik,
Québec (Fig. 1a). It is approximately 4.5 km long and 1.5 km wide,
following a northwest-southeast orientation. At the southeastern
end of the valley lies Tasiujaq Lake (formerly named GuillaumeDelisle Lake or Richmond Gulf), a 691 km2 brackish water body
connected with the Hudson Bay by a narrow cataclinal channel
called Le Goulet (ARK, 2007). The lake is part of Tursujuq National
Park, created in 2013.
The regional geology is characterized by a Paleoproterozoic
volcano-sedimentary sequence lying unconformably on the Precambrian shield (Fig. 1b). The volcano-sedimentary sequence
includes limestone, quartz arenite, dolomite and sandstone
strata (Qingaaluk Formation) underlying a thick (∼15 m) basalt
layer (Nastapoka Group) dipping westward (Stockwell et al.,
1979; Chandler and Schwarz, 1980; Chandler, 1988; Eaton and
Derbyshire, 2010). This asymmetrical monoclinal relief (cuesta)
consists of a gentle western slope and a steep eastern slope and
extends over 650 km along the east coast of Hudson Bay (Dionne,
1976; Guimont and Laverdière, 1980). Tasiapik Valley lies at the
frontslope of the cuesta on its southwestern side, whereas the
northeastern side consists of a residual butte called Umiujaaluk
Hill.
Quaternary deposits on the east coast of Hudson Bay are the
result of a succession of sedimentary environments following the
retreat of the Laurentide Ice Sheet at about 8200 cal. BP (HillaireMarcel, 1976; Allard and Séguin, 1985; Lavoie et al., 2012) (Fig. 1b).
Lowlands below 271 m a.s.l, the altitudinal limit of the postglacial
Tyrrell Sea in Tasiujaq Lake area (Fraser et al., 2005; Lavoie et al.,
2012), are covered by deep-water and shallow-water marine sediments, and littoral deposits (raised beaches) that were formed
during stages of rapid glacio-isostatic uplift (Hillaire-Marcel, 1976).
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
A glaciomarine fan complex lies in the upstream part of the valley.
It consists of fluvioglacial material that was deposited during a stillstand of the ice margin around 8000 cal. BP (Lajeunesse and Allard,
2003b).
The study area has a cold subarctic climate and it is located
in the discontinuous permafrost zone (Allard and Lemay, 2012).
Mean annual air temperature recorded between 2013 and 2017
varies between -5.6 and -4.2 ◦ C, with maxima of 23 ◦ C and minima
of −36 ◦ C (Fortier, 2017). Mean annual precipitation is approximately 500 mm, with 40% falling as snow (Ménard et al., 1998).
The Umiujaq area is located at the edge of the shrub and forest
tundra zones; low shrubs, ericaceous plants and lichens cover the
upstream part of Tasiapik Valley, while dense forest cover occupies
the downstream part (Payette, 1983). Shrub cover has expanded
significantly (shrubification) during the 20th century (Ménard et al.,
1998; Provencher-Nolet et al., 2014; Pelletier et al., 2018).
The SW side of Tasiapik Valley has a near-vertical rockwall. It is
approximately 50 m high in the upstream part of the valley, but it
increases to 230 m near Tasiujaq Lake in the downstream part. Slope
deposits lie at the base of the escarpment, connecting the rockwall
to the valley floor, but have also accumulated on basaltic rocky
outcrops in the uppermost part of the rockwall (Fig. 1c). The NE side
features a step-like topography, with slope deposits either located
at the base of the slope or perched on basaltic and sedimentary
rocky outcrops. A gravel road connecting Umiujaq to Tasiujaq Lake
follows the cuesta frontslope on the SW side.
1.2. Methods
For this study, 18 talus slopes were investigated, on the SW
and NE sides of Tasiapik Valley (Fig. 1c). Data were collected
over four field campaigns during the summers of 2016 (August),
2017 (August) and 2018 (June and August). Several slope deposits
were identified by satellite imagery prior to initiating fieldwork.
Two sets of orthophotos from Québec’s Ministère des Ressources
naturelles et de la Faune (MRNF) were used, one dating from 2004
(scale 1/10,000, 25 cm resolution) and the other from 2010 (scale
1/10,000, 15 cm resolution).
Topographic surveys were conducted along 18 longitudinal
transects using a Leica DGPS (Differential Global Positioning System). Waypoints were recorded from the apex of the slope deposits
to their base, perpendicular to the rock face. Data were processed
in ArcGIS and Excel to produce topographic profiles, revealing the
microtopographic features such as inflection and texture in accurate details. The estimation of the stage of evolution is carried out
from the Ho/Hi ratio, where Ho corresponds to the height of the
talus slope and Hi to the total height of the slope including the rockwall (Francou, 1988; Sellier, 1992). The ratio gives an overview of
the exhaustion of the remaining rockwall (debris source) in concomitance with the talus slope formation on a longer timescale,
namely since the last deglaciation. For example, a ratio approaching 1 indicates an advanced slope development stage due to the
low height of the residual wall compared to the height of the talus
slope.
Topographic data and satellite imagery were used to measure
rockfall runout distances (horizontal travel distance) calculated
from the source area to the farthest slope debris. In addition, the
reach angle, calculated from the source-area to the farthest slope
debris, and the shadow angle, calculated from the apex of the talus
to the farthest slope debris, were documented to provide information about the extent of slope processes in the area.
On 12 of the 18 longitudinal transects, granulometric and petrographic surveys were conducted by sampling 25 randomly selected
rock fragments at intervals of 10–15 m along the transects. Debris
were measured along three axes: length (a-axis), width (b-axis) and
thickness (c-axis). Measurements were compiled in Excel and then
3
analyzed to produce descriptive statistics. Morphometric indices
were also calculated from these measures (Pérez, 1989, Hétu and
Gray, 2000). The flattening index (Fi) is calculated as follows:
Fi =
a+b
2c
(1)
where a corresponds to the length, b to the width and c to the thickness of the fragment (Cailleux, 1947). A high Fi value indicates that
the debris has a flatter shape. The elongation index (Li) is calculated
as follows:
a
Li =
(2)
b
where a and b correspond to the length and width of the fragment
(Schneiderhöhn, 1954). A high Li value indicates that the debris
tends to be elongated. Finally, the sphericity index (Si) is calculated
as follows:
Si =
bc 13
a2
(3)
where a, b and c correspond to the length, width and thickness of
the fragment (Krumbein, 1941). A value Si approaching 1 indicates
that the debris has a more massive shape, spherical in the case of a
rounded fragment and cubic for an angular fragment. These indices
document the falling behavior of clasts, since spherical debris are
prone to rolling, while elongated flat debris are more likely to slide.
Petrographic surveys provide lithological data for the measured
fragments. Their origin, either local (associated with the local slope
development) or exogenous (generally from glacial transport and
deposition), is closely related to their lithology, thus their general
shape, and their position on the slope. The edges of the debris were
characterized, with a view to determining their origin: an angular fragment has undergone very little erosion, indicating a short
transportation distance/local source, while debris transported by
glaciers or reworked in the Tyrrell Sea has a pronounced rounded
shape.
Vegetation cover was described at each sampling station in
order to assess recent and current process activity. Hierarchical values were attributed to each station based on the type of vegetation
and the estimated percentage of coverage on the debris, providing
relative age-estimates:
1) Fresh debris: no lichen species observed on the clast;
2) Recent debris: some lichen species observed on the clast;
3) Medium-aged debris: several lichen species partially cover the
clast;
4) Old-aged debris: several species of lichens and mosses partially
cover the clast;
5) Very old-aged debris: several species of lichens and mosses
totally cover the clast; potentially also covered with low shrubs.
Vegetation classification values and Ho/Hi ratio values were
used to estimate the stage of slope development. The addition of
these two values gives an overview of the slope evolution from both
short term (vegetation) and long term (Ho/Hi) perspectives. Values
of 1 to 5 were assigned to each longitudinal profile according to
their Ho/Hi ratio, following the Jenks natural breaks classification
method (Jenks, 1967); a value approaching 1 indicates a low Ho/Hi
ratio, thus a younger development stage. The vegetation values (i.e.
the lowest - freshest - value per profile, ranked 1 to 5 according to
the relative age estimate described above) were added to provide
an overall development score. In addition, the age of shrubs at the
bottom of talus along the SW-07, SW-08 and SW-09 profiles was
determined using dendrochronology on 11 black spruce samples
(Picea mariana).
Finally, in order to monitor slope movements on a shorter time
scale, three Reconyx PC800 automatic time-lapse cameras were
4
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
Fig. 2. Location of the cameras along the SW side of Tasiapik Valley. Frame view of each camera is shown on the left. Source of background image: UMI orthomosaic (2010).
installed on the SW side of the valley in August 2017. One of the
cameras (TAS 1) is located on the cuesta edge, above the rockwall
and talus along the SW-07 and SW-08 profiles (Fig. 2). The latter is
covered by a second camera (TAS 2) that is located 300 m away from
the rockwall. A third camera (TAS 3) is located further upstream
near the base of the talus along the SW-06 profile. Approximately
14,000 photos were taken over a one-year period from August 2017
to August 2018 in the valley. Photos were taken during daytime at
one-hour intervals until June 2018, then at 15 or 30 min intervals
(depending on the location) until August 2018.
2. Results
2.1. Topography of slope deposits
The SW profiles show a steeper slope gradient, with a mean
angle of 25.3◦ and a median angle of 26.3◦ , while the NE profiles
have a mean angle of 21.2◦ and a median angle of 17.5◦ (Table 1;
Fig. 3). A slight or distinct concavity of the talus slope is apparent
on six of the nine SW profiles, while the SW-01 and SW-06 profiles
are virtually linear. The SW-09 profile shows a more complex shape
(linear proximal part and chaotic distal part). On the NE side, there
is no distinct concave profile, yet seven of the nine profiles show
either a slightly concave or a linear shape, while the NE-04 and NE07 profiles respectively show a complex and a convex shape. Four
profiles (SW-07, SW-08, SW-09 and NE-09) exhibit a strong basal
concavity (Fig. 3).
2.2. Relative dating of slopes
The mean Ho/Hi index on the SW (0.36) and NE sides (0.25)
indicates that the remaining rockwall is generally higher than talus
slopes (index < 0.5) (Table 2). However, the step-like topography
on the NE side, compared to the near-vertical rockwall on the SW
side, could mean that the NE side has reached a more advanced
stage of development. Talus slopes near the southern margin of
both sides (along the SW-07, SW-08, SW-09 and NE-09 profiles)
are located under high vertical rockwalls with Ho/Hi index below
0.2, thus indicating a younger development stage. The SW-04 and
SW-05 profiles indicate an older stage than the other SW profiles,
with respective index values of 0.56 and 0.62.
5
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
Table 1
Topographic parameters of the investigated slopes.
Profiles
Slope angle (◦ )
SW-01
SW-02
SW-03
SW-04
SW-05
SW-06
SW-07
SW-08
SW-09
NE-01
NE-02
NE-03
NE-04
NE-05
NE-06
NE-07
NE-08
NE-09
18.4
28.1
28.3
19.3
24.9
27.1
26.3
30.0
25.2
14.2
12.3
17.5
16.1
17.4
31.8
27.1
23.2
31.6
Mean angle (◦ )
Median angle (◦ )
25.3
26.3
21.2
17.5
Inflection
linear
slightly concave
slightly concave
concave
concave
linear
concave
concave
complex
linear
slightly concave
slightly concave
complex
linear
linear
convex
linear
slightly concave
Table 2
Calculation of Ho/Hi index.
Profiles
Talus base elev. (m)
Talus apex elev. (m)
Ho value
Rockwall elev. (m)
Hi value
Ho/Hi index
SW-01
SW-02
SW-03
SW-04
SW-05
SW-06
SW-07
SW-08
SW-09
NE-01
NE-02
NE-03
NE-04
NE-05
NE-06
NE-07
NE-08
NE-09
171.3
147.3
146.3
151.4
148.0
127.4
46.5
47.6
49.6
156.1
157.2
164.6
163.0
162.1
225.6
131.7
134.6
161.2
187.6
178.3
176.8
183.2
187.9
159.2
82.0
67.2
75.9
169.9
167.3
182.8
184.3
178.1
260.8
167.5
184.0
188.8
16.3
31.0
30.5
31.8
39.9
31.8
35.5
19.6
26.3
13.8
10.1
18.2
21.3
16.0
35.2
35.8
49.4
27.6
222
216
208
208
212
212
214
222
272
202
252
254
270
282
282
292
302
312
50.7
68.7
31.7
56.6
64.0
84.6
167.5
174.4
222.4
45.9
94.8
89.4
107.0
119.9
54.4
160.3
167.4
150.8
0.321
0.451
0.494
0.562
0.623
0.376
0.212
0.112
0.118
0.301
0.107
0.204
0.199
0.133
0.624
0.223
0.295
0.183
Examination of the vegetation covering the surficial debris on
the slope deposits revealed the presence and position of few fresh
deposits. Most of the debris had a clear vegetation cover on the SW
side (Fig. 4). In the apical parts of the talus slopes, various lichens
and/or mosses are abundant, while a discontinuous thin strip of
herbaceous plants and low shrubs is located at the edge of the rockwall undisturbed by present-day slope activity. Distal parts also
feature abundant lichens and mosses on most of the clasts along
with thick mosses covering the slope deposits and low shrubs; this
trend is especially evident on the SW-07 and SW-09 profiles. A similar trend was observed on the NE profiles, as most sampling stations
in the apical part of the talus slopes show medium to old-age status, while sampling stations in the distal parts indicate older-age
status. However, debris along the SW-08, NE-06 and NE-09 profiles
appear to be more recent, with little overall coverage and the presence of few lichens. Some fresh debris were scattered along most
of the profiles (Table 3).
By adding up the Ho/Hi index values and the vegetation classification values, we can estimate the developmental stage of the
slopes. As shown on Fig. 5, the SW-08 and NE-09 profiles seem to be
at the youngest developmental stage among all the talus slopes. The
Ho/Hi index and the vegetation classification values are consistent
for some profiles, showing a concomitance for both parameters.
For example, the talus slope along NE-09 profile has a low Ho/Hi
index (0.183) and there is very poor lichen cover on the debris.
However, the two parameters proved to be contradictory for some
talus slopes, particularly for the SW-09 profiles, due to the overly
high rockwall (increased debris supply potential) and the presence
of well-developed vegetation (limited debris supply on the talus
slope). Both of these findings indicate that the debris supply is
sporadic.
2.3. Source and morphometry of slope deposits
Three classes of debris were identified along the investigated
profiles: 1) basalt, 2) sedimentary rocks (comprising dolomite,
limestone, quartz arenite and sandstone), and 3) granitic gneiss
(Fig. 6).
On the SW side, sedimentary rock debris comprised 64.5% of
the sampled clasts, whereas basalt debris accounted for 35.4% and
gneiss for 0.07%. Sedimentary rock debris represented a larger proportion on the NE side, accounting for 89.4% of all clasts, while basalt
and gneiss accounted for 6.4% and 4.2%. Assuming the top basalt
layer is ∼15 m thick throughout the valley; those values coincide
with the large proportion of sedimentary rock strata available for
debris supply on the exposed rockwall. Sedimentary rock strata
account for 90% (∼140 m) of the rockwall (∼155 m) above the SW07, SW-08 and SW-09 profiles. In the upstream part of the SW
side, sedimentary rock strata account for 57% (∼20 m) of the rockwall (∼35 m). On the NE side, the top basalt layer has considerably
receded on most of the investigated slopes, revealing rocky outcrops composed of sedimentary rock strata. However, the rockwall
6
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
Fig. 3. Longitudinal cross-section of the investigated slopes. Source of background image: UMI orthomosaic (2010).
above the NE-06 profile is mainly composed of basalt (68% of total
height).
By comparing the respective proportions of each lithology at
sampling stations along the profiles, many of the slope deposits
show consistent ratios of sedimentary rocks and/or basalt debris
from the apex to the base of the slope. On the SW side, the proportion of basalt debris varies between 80% and 84% throughout
the sampling stations on the SW-02 profile and between 14% and
26% on the SW-07 profile. On the NE side, the proportion of sedimentary rock debris varies between 80% and 88% on the NE-02
profile, while the NE-08 and NE-09 profiles show no difference
as sedimentary rock debris compose 100% of the talus. However,
some of the slope deposits show an increasing proportion of basalt
material toward the foot of the talus. For example, along the SW08 profile, the basalt debris percentage increases from 17% at the
apex to 37% at the bottom of the slope; on the SW-09 profile, it
increases significantly from 12% to 100%. Finally, on the NE-06 profile, the basalt debris are only located at the bottom of the talus,
whereas the sedimentary rock debris comprises the entire apical
part.
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
7
Fig. 4. Vegetation stages along the talus slopes.
Fig. 5. Slope development stage according to three parameters: 1) vegetation; 2) Ho/Hi index and 3) a combination of the two parameters. Source of background image: UMI
orthomosaic (2010).
Clasts were measured to document the size and morphometry of the slope deposits, and their distribution along the deposits
(Fig. 7). On the SW side, the debris (a-axis) range in size from 30
to 154 cm on average (Table 3). However, standard deviation values show important disparities between the measured clasts, as the
largest clasts range from 200 to 1900 cm. In addition, mean flatness
index values range from 2.36 to 4.22. Lengthening and sphericity
indices do not vary much, ranging from 1.48 to 1.86 and from 0.54 to
0.65. On the NE side, sizes range from 83 to 178 cm on average, while
standard deviation values show important disparities between the
8
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
Table 3
Debris size and morphometric parameters along the longitudinal profiles.
Profiles
SW-01
SW-02
SW-03
SW-06
SW-07
SW-08
median
mean
standard dev.
maximum
minimum
range
median
mean
standard dev.
maximum
minimum
range
median
mean
standard dev.
maximum
minimum
range
median
mean
standard dev.
maximum
minimum
range
median
mean
standard dev.
maximum
minimum
range
median
mean
standard dev.
maximum
minimum
range
a-axis (cm)
b-axis (cm)
c-axis (cm)
b/a
c/b
Flatness
Lengthening
Sphericity
45.00
65.54
58.57
450.00
11.00
439.00
60.00
78.18
80.30
600.00
8.00
592.00
84.00
154.48
262.49
1900.00
8.00
1892.00
50.00
77.14
78.61
470.00
5.00
465.00
23.50
30.08
24.17
206.00
6.00
200.00
36.00
48.16
42.23
280.00
4.00
276.00
32.00
45.74
44.99
350.00
7.00
343.00
37.00
48.93
48.04
350.00
5.00
345.00
50.00
95.20
141.39
1000.00
7.00
993.00
36.00
54.89
55.69
310.00
5.00
305.00
14.00
17.27
11.83
100.00
2.00
98.00
20.00
27.31
26.74
223.00
3.00
220.00
19.00
27.54
26.57
180.00
3.00
177.00
21.00
30.99
35.48
290.00
1.00
289.00
23.00
60.77
108.40
800.00
3.00
797.00
17.00
29.70
34.26
190.00
3.00
187.00
6.00
8.24
7.91
70.00
1.00
69.00
11.00
14.29
12.83
100.00
1.00
99.00
0.70
0.69
0.16
0.98
0.37
0.61
0.63
0.64
0.17
1.15
0.28
0.88
0.67
0.67
0.18
1.00
0.30
0.70
0.72
0.72
0.17
1.04
0.26
0.78
0.63
0.62
0.18
1.00
0.15
0.85
0.59
0.60
0.18
1.00
0.10
0.90
0.63
0.64
0.22
1.00
0.11
0.89
0.64
0.63
0.21
1.04
0.05
0.99
0.57
0.58
0.24
1.11
0.10
1.01
0.58
0.56
0.24
1.00
0.15
0.85
0.50
0.89
1.33
10.00
0.04
9.96
0.56
0.57
0.22
1.00
0.09
0.91
1.96
2.36
1.24
10.42
1.10
9.31
2.10
2.55
2.44
27.50
1.17
26.33
2.21
2.84
1.88
13.75
1.07
12.68
2.21
2.69
1.36
7.50
1.06
6.44
3.00
4.22
3.78
29.50
1.10
28.40
2.57
2.96
1.53
11.17
1.07
10.10
1.43
1.53
0.40
2.73
1.03
1.70
1.58
1.68
0.51
3.64
0.87
2.77
1.50
1.62
0.49
3.32
1.00
2.32
1.38
1.48
0.45
3.86
0.96
2.90
1.58
1.81
0.76
6.80
1.00
5.80
1.69
1.86
0.80
10.42
1.00
9.42
0.65
0.65
0.12
0.93
0.38
0.55
0.62
0.62
0.11
0.87
0.26
0.61
0.60
0.62
0.13
0.91
0.32
0.59
0.61
0.64
0.13
0.93
0.33
0.60
0.53
0.54
0.13
0.94
0.21
0.73
0.56
0.57
0.12
0.92
0.19
0.73
measured clasts. The largest clasts range in size between 350 and
950 cm, while the mean flatness index values range from 2.29 to
3.58. Lengthening and sphericity indices show similar values as the
profiles on the SW side, ranging from 1.40 to 1.82, and from 0.57
to 0.69. By comparing the calculated index values along all profiles,
the clasts tend to have a flatter and more elongated shape in the
downstream part of the valley on the SW side (SW-07, SW-08 and
SW-09), and along the NE-06 and NE-09 profiles. This morphology
is more associated with sedimentary rock debris due to the dismantling of the thin sedimentary layers which represents the majority
(84.5%) of the sampled debris along these profiles (given the larger
proportion of sedimentary rock within the rockwall). For most of
the profiles, the mean and median a-axis values tend to increase
toward the base of the slopes. As for the morphometry indices, 50%
of the profiles show an increasing trend for flatness and sphericity
values toward the base of the slope, while the lengthening index of
clasts increases from the apex to the base for 75% of the profiles.
Debris from every sampling station on the SW side were analyzed and described as being either angular or subangular, which
indicates their local slope-related provenance, whereas the NE side
shows a greater diversity of debris shape. For the NE-02 and NE-04
profiles, rounded/sub-rounded debris are abundant at the base and
angular debris are found near the apex of the talus slopes, meaning that slope-related debris have accumulated in the proximal part
and mixed with rounded heavily reworked debris toward the distal
part.
2.4. Debris runout
Several scattered clasts are located in the distal parts of the
slopes. These are mostly located on the SW side and some of them
were deposited only a few meters from the road. Their angular
shape and lithology (mostly basalt) indicate their local sloperelated provenance, by contrast with rounded fluvioglacial/glacial
debris.
According to Corominas et al. (2003), the maximum reach angle
for small-scale (1-10 m3 ) rockfalls on an unobstructed path is 48◦ .
Therefore, material falling from the uppermost source area in the
rockwall (basalt layer) would theoretically be transported no further than the talus at the bottom of the slope, where the terrain
levels out and the shrub vegetation is often denser (Fig. 8). This
perfectly matches the very large boulder accumulation on the SW09 profile. However, the reach angle for the farthest basalt debris,
located beyond the talus slopes, ranges from 24◦ to 40◦ . These
lower reach angles show a greater horizontal displacement of fallen
debris, resulting from either a large-scale rockfall (10-100 m3 for a
40◦ reach angle; 100-1000 m3 for a 33◦ reach angle; >1000 m3 for a
26◦ reach angle) (Corominas et al., 2003) or an external process. For
example, there are a dozen large basalt boulders (a-axis > 100 cm)
located in the distal part of the SW-07 profile at a 30◦ angle. However, the volume is not sufficient for such a displacement to result
from a rockfall. The shadow angle ranges from 10◦ to 25◦ , with the
lowest angles measured on the SW-07 (10◦ ) and SW-09 (17◦ ) profiles. These angles are lower than the minimum rockfall shadow
angle ranging from 22◦ to 30◦ (Rapp, 1960; Govi, 1977; Lied, 1977;
Hungr and Evans, 1988; Evans and Hungr, 1993). Therefore, it can
be assumed that slope debris were deposited beyond the rockfall
runout zone by another process. The hypothetical travel distance
for the farthest debris ranges from 105 to 318 m, with the longest
distances measured in the downstream part of the SW side.
In June 2018, numerous dirty snow-avalanche deposits were
observed on the SW side, but their terminus rarely exceeded the
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
9
Fig. 6. Proportion of the different lithologies along the talus slope.
talus slope, thus supporting the assumption that snow-avalanche
runout is generally limited to the foot of the talus. However,
longer runout was observed for clean snow avalanches occurring
in February through April. Extreme snow avalanche runout cannot
be documented at this stage in the valley.
Tree-ring counting on 11 black spruce (Picea mariana) samples,
dating back to 1900–1930, and located at the foot of the SW-07 and
SW-08 profiles, showed clear signs of eccentric growth, such as the
formation of reaction wood during their lifespan. As these reaction
wood periods are not significantly concomitant from one tree to
another, wind or snow cover could hardly be the main controlling
factors for these deformations. Instead, slope processes, such as
snow avalanches, are likely causes. However, the limited number
of samples does not support a precise chronology of events.
2.5. Short-term slope dynamics
Analysis of the photographs from the three automatic cameras
has documented the slope dynamics from summer 2017 to summer
2018, with the record of one full winter. Snow avalanches were the
main gravitational processes observed, occurring from November
2017 to June 2018; snow-avalanche events occurred more frequently from the end of May 2018. During this period, most of
the snow-avalanche deposits exhibited a dirty appearance, because
rock debris were incorporated in the dense and humid snow. This
characteristic implies a flowing motion with frequent contact with
the regolith (Fig. 9a). These observations were validated in the field
in June 2018, when about 20 wet snow avalanches were observed
on the SW side of the valley, from June 11 to June 15.
Along the SW-07 and SW-08 profiles (TAS2), some larger rock
debris (∼1 m a-axis) were transported and deposited downslope
by snow avalanches. Other debris were observed falling onto the
snow-covered talus until early July 2018, concomitant with snowavalanche events. Two significant rockfalls occurred one-hour apart
on June 30 2018. Several debris could be observed sliding on the
snow covered talus after their fall; they travelled ∼30 m before settling at mid-slope (Fig. 9b). During the same period, only a few
small snow avalanches occurred in the area covered by camera
TAS3, along the SW-06 profile, and no rockfall were observed. Out
of the seasonal presence of snow, no movement was observed on
the slopes from the analysis of the photographs.
Based on the observations during 2017–2018, slope movements
occur more frequently in the spring. Prior to June 2018, no rocky
deposits on the snow cover had been observed following snowavalanches, which occurred sporadically from December 2017 to
April 2018.
3. Discussion
3.1. Talus slope formation
Results from the morphometric and petrographic surveys suggest an accumulation of rock debris on the SW and NE sides
resulting from successive discrete rockfalls that formed into scree
slopes. The vast majority of the sampled debris is from a local
source, namely the lithologies exposed on the rockwall, and they
exhibit subangular to angular shapes.
Vegetation cover on the talus suggests that the most recent
debris supply occurred along the SW-08, NE-06 and NE-09 pro-
10
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
Fig. 7. Distribution of the debris size (a-axis) along the talus slopes. Close-up of the largest sampled debris along SW-09 profile (A) and NE-09 profile (B). Source of images:
UMI orthomosaic (2010).
files, as the debris were characterized by low coverage and few
or no species of lichens. The debris likely originated from recent
notches formed in the rockwall above the SW-08 (Fig. 10a) and
NE-09 profiles (Fig. 10b). Along the other profiles (SW-01, SW09, NE-02, NE-07), the freshest debris are found in the proximal
part of the talus, presumably supplied by small-scale rockfalls. The
proximal part of talus slopes generally consists of smaller debris,
whereas larger debris and well-developed vegetation are found on
the distal part of most of the slopes. Thus, small-scale rockfalls
are unlikely to reach the base of the talus, where the accumulation is associated with large-scale rockfalls that have occurred in
the past. Nevertheless, debris-size-sorting on the talus slopes, with
increasing debris size toward the base, could either be attributable
to rockfalls (Kirkby and Statham, 1975; Statham, 1976; Church
et al., 1979) or to talus reworking by snow avalanches (Rapp, 1960;
Luckman, 1978, 1988; Jomelli and Francou, 2000; Decaulne and
Saemundsson, 2006). Basalt debris are larger and exhibit a more
massive shape than sedimentary rock debris. They also have a
higher fall height, resulting in greater travel distances toward the
distal part of talus slopes, as observed along the SW-06, SW-08,
SW-09, NE-02 and NE-06 profiles.
Nowadays, expanding shrub cover in Tasiapik Valley
(Provencher-Nolet et al., 2014; Pelletier, 2015; Pelletier et al.,
2018) could suggest a recent decline in debris supply, as rockfalls would not be frequent enough to limit shrub expansion
on some slopes. On the NE side, a dense shrub cover (Fig. 10c)
separates perched and basal talus slopes, while shrubs established
themselves at the very base of the talus slope along the SW-07,
SW-08 and SW-09 profiles (Fig. 10d). In both cases, shrubs have
developed over highly weathered rock debris or rocky outcrops.
Such a pattern of vegetation colonization highlights areas that lie
beyond the reach of the most recent slope dynamics.
3.2. Slope debris redistribution
Scree slopes generally have a mean angle between 25◦ and 30◦ ,
as reported by Sauchyn (1986) and Francou and Manté (1990).
Steeper scree slopes (>30◦ ) have been studied in Québec, namely in
Schefferville (Andrews, 1961), in Gaspésie (Andrews, 1961; Hétu,
1995; Hétu and Gray, 2000; Germain and Hétu, 2016) and on the
central islands of Wiyâshakimî Lake (Decaulne et al., 2018). Scree
slopes also have a segmented profile with a steeper gradient in the
proximal part and a strong basal concavity in the distal part; this
slope geometry was observed on several talus slopes (Fig. 3). Given
these findings, debris redistribution must be an ongoing process
(Church et al., 1979; Francou and Manté, 1990) that is likely the
result of snow-avalanche rework. Their impact results in a concave
inflection of the talus slopes and an increase in debris size towards
the base, and can be observed on most talus slopes (Rapp, 1960;
Luckman, 1977, 1978; Luckman, 1988; Jomelli and Francou, 2000;
Decaulne and Saemundsson, 2006). The general concavity of talus
slopes suggests that surficial debris redistribution is more important than debris supply from the rockwall. Instead, smaller debris
tend to be trapped in the numerous surficial cavities that are found
on the talus (Statham, 1976; De Blasio and Saeter, 2010).
Photographic monitoring of a portion of the SW side found
evidence of discrete rockfall events in June and July 2018, but it
has mainly highlighted numerous wet and dirty snow avalanches
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
11
Fig. 8. Cross-section of the slopes (talus/rockwall) showing the reach and shadow angle for the farthest slope debris, and the hypothetical travel distance from the source
area. Source of background image: MRFP.
Fig. 9. Snow-avalanche event exhibiting a dirty deposit, as the result of incorporated debris (A); discrete rockfall event and debris sliding onto the snow-covered talus (B).
that occurred in the spring (April-June 2018). The latter have a
greater erosive capacity due to higher friction at their base. They
can also dislodge rock material from the bare rockwall, thus supplying the talus slope with new debris (Gardner, 1983a; Jomelli,
1999; McClung and Schaerer, 2006). During the same period, the
snow-covered talus enabled recently fallen debris to slide down
from the apex to mid-slope (∼30 m). Sedimentary rock debris are
prone to efficient sliding because of their flattened shape (the mean
flatness index for sedimentary rock debris (3.04) is higher than
basalt debris (2.49) on the investigated slopes), as reported by
12
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
Fig. 10. Recent notches formed in the rockwall along SW-07 (A) and NE-09 profiles (B); dense shrub cover at the base of the slope near SW-07 and SW-08 profiles (C) and
mid-slope near NE-05 and NE-06 profiles (D).
Pérez (1989) at Lassen Peak, California, and Hétu (1995) in Gaspésie,
Québec.
Results of reach angle measurements have shown that scattered
slope debris located far downslope originated from large-scale
rockfall events along most of the talus slopes, as the angle is greater
than 26◦ (Corominas et al., 2003). Furthermore, low shadow angles,
especially along the SW-07 and SW-09 profiles, provide evidence
that their deposition is not entirely due to the rockfall itself, and that
the transition from the talus to the valley floor requires an external transport agent (Domaas, 1994). The lowest measured shadow
angle was 17◦ in Norway (Domaas, 1994), whereas it generally
ranges between 22◦ and 30◦ (Rapp, 1960; Govi, 1977; Lied, 1977;
Hungr and Evans, 1988; Evans and Hungr, 1993). Debris falling onto
an ice-covered talus slope are unlikely to be deposited at such a
great distance, considering that Evans and Hungr (1993) reported
a 24◦ shadow angle for a boulder falling onto a smooth glacier. In
addition, as mentioned above, the current position of the farthest
lying boulders does not support the hypothesis of snow-avalancherelated transport. The abrupt slope transition (steep rockwall to
valley floor) is not conducive to extreme snow-avalanche runout
distances (Bakkehoi et al., 1983; McClung and Schaerer, 2006).
From the base of the talus, the slope angle decreases to ∼4◦ , which
means that a boulder that is located beyond this zone could not
have been transported by a snow avalanche, being located outside
of the runout zone.
Therefore, the deposition of debris would most likely be related
to the deglaciation/postglacial marine episode, starting at around
8200 cal. BP in the area (Hillaire-Marcel, 1976; Lavoie et al., 2012).
In particular, the presence of highly altered basalt boulders (Fig. 11)
overlying littoral marine sediments would suggest that their depo-
sition occurred during the regression of the Tyrrell Sea. The great
distance from the surrounding slopes could then be attributable
to sea-ice-related processes, such as ice rafting or ice pushing.
Finally, similar processes could have distributed glacially transported rounded boulders against the talus slopes in the lower part
of the NE slope, where the vast majority of the subrounded and/or
gneiss debris were observed.
3.3. Rockwall erosion
Different dismantling mechanisms affect the various volcanosedimentary lithologies exposed on the cuesta frontslope. The top
basalt layer shows a distinct columnar polygonal jointing, resulting
in the detachment and subsequent fall of large (>5 m a-axis) monoliths (Fig. 12a). The edge of the basalt layer on the SW side reveals a
sawtooth shape (Fig. 12b), also documented by Belzile (1984) at
the Manitounuk Peninsula, 100 km south of Umiujaq (55◦ 42′ N,
77◦ 07′ W). This pattern highlights the numerous monolith falls that
occurred in the past. Ongoing periglacial processes such as gelifraction and frost heave caused extensive basalt dismantling (Fig. 12c).
Michaud and Dionne (1987) observed similar periglacial weathering in the basalt bedrock about 45 km south of Umiujaq (56◦ 09′ N,
76◦ 36′ W), resulting in block field development. Exhaustion of the
underlying sedimentary rock layers could also cause the basalt layer
to overhang the slope, eventually leading to rockfalls (Fig. 12d).
The underlying sedimentary rock layers are prone to frost
shattering due to their thin-bedded sub-horizontal structure and
abundant fractures exposed on the rockwall, which means that
these layers are also subject to rockfalls (Frayssines and Hantz,
2006; Mateos et al., 2012; Letortu, 2013; D’Amato et al., 2016).
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
13
Fig. 11. Cluster of basalt boulders located ∼250 m away from the slope, overlying littoral sediments.
Fig. 12. Erosion of the basalt layer at the top of the cuesta on the SW side, with tall monolith on the verge of falling (A), resulting in a sawtooth shape exposed on the rockwall
(B); frost heave and extensive jointing of the basalt bedrock (C); basalt and sandstone layers overhanging quartz arenite layers on the NE side (D). Photos: Decaulne, 2017.
Notches have recently developed in the sedimentary rock layers, as shown on Fig. 10a–b, and have formed prominent debris
cones. Subsequently, these notches will induce more discrete rockfalls from the overlying layers. Locally, rocky outcrops reveal small
escarpments (∼1-5 m high) that supply debris for nearby talus
through short fall height and distances, especially on the NE side.
Freeze-thaw weathering is known to be a major factor in rockfall triggering, especially on deglaciated rockwalls (Matsuoka and
Sakai, 1999; Ballantyne, 2002; Matsuoka, 2008). The air temperature data record for the Umiujaq region (CEN, 2018), from
September 2002 to June 2018, suggests that freeze-thaw cycles
are on average more frequent during the spring (27 - March to
June) than in autumn (19 – September to December), but with
important interseasonal differences. For instance, only 9 cycles
occurred in autumn 2007, followed by 34 cycles during the next
spring. However, according to Hales and Roering (2007), the frequency of freeze-thaw cycles would not be a determining factor
in rockfall triggering. Instead, they propose a temperature threshold (−3 ◦ C to −8 ◦ C) whereby thin water layers can infiltrate and
cause segregated ice to form in the fractures, leading to the subsequent dismantling of the bedrock. The air temperature data from
October 2002 to May 2018 (total of 5688 days) suggests that there
were 642 days that had a mean temperature between -3 and
−8 ◦ C. They are more frequent in November (31.8%), April (20.9%),
December (15.7%) and May (13.9%). At a depth of 100 cm in the
basalt bedrock, the number of days within this threshold temperature range were more frequent in April (39.5%), December (35.2%)
and January (11.9%) between 2001 and 2006 (Allard et al., 2016),
which shows a potential delay from autumn to winter compared
to air temperature. It should be noted that slope orientation and
14
S. Veilleux, N. Bhiry and A. Decaulne / Geomorphology 349 (2020) 106911
altitude could influence the temperature at the surface of the rockwall, and that significant differences could be expected between
the SW (NE facing) and NE sides of the valley. Nevertheless, the
available freeze-thaw cycle data and temperature threshold data
concur with the seasonal trend (autumn/spring) for potential rockfall activity (Gardner, 1983b; Matsuoka and Sakai, 1999; Hales and
Roering, 2007; Matsuoka, 2008).
On a longer time scale, paraglacial adjustment is another
rockfall-triggering factor that should be considered on deglaciated
slopes (Ballantyne and Benn, 1994; Ballantyne, 2002; KellererPirklbauer et al., 2010; Cossart et al., 2013; Grämiger et al., 2017).
Deglaciation occurred at about 8200 cal. BP in the study area
and induced rapid rates of isostatic uplift (Hillaire-Marcel, 1976;
Lajeunesse and Allard, 2003a; Lavoie et al., 2012). This in turn
caused an adjustment of the rock substratum and led to stressrelease on the rock slopes resulting from the ice retreat, also known
as ‘debuttressing’ (Ballantyne, 2002). The rockwall became susceptible to slope failures as a result of these changes.
Slope development results (Ho/Hi index) suggested that the SW
side of the valley is at a younger stage of development, considering the height of the remaining rockwall compared to the step-like
topography of the NE side. However, well-developed vegetation
on the debris observed throughout the SW profiles (except for the
SW-08 profiles) raises questions about the spatial distribution and
frequency of the slope processes. The exhaustion model, in which
slope failures occurring after ice retreat gradually decline until
complete exhaustion of the rockwall (Cruden and Hu, 1993), would
not fit in this case since local characteristics control the triggering of
slope processes (i.e. periglacial processes). A bimodal/multimodal
model (Krautblatter and Leith, 2015) would thus be more representative considering that the region has undergone periods of
climate fluctuations since its deglaciation (i.e. Little Ice Age). However, the occurrence of extreme meteorological events, such as
warm summers (Gruber et al., 2004) and heavy rainfalls (Rapp
and Strömquist, 1976; Delonca et al., 2014) should also be considered, as these events can trigger rockfalls and contribute to
the slope development. Furthermore, a winter/spring period with
many snow avalanches could result in increased erosion on the
slope, thus favoring rockfalls, and cleaning available debris accumulated in the rockwall chutes.
4. Conclusion
In this study, we provided evidence of slope activity resulting in
talus formation in Tasiapik Valley. The results of the topographic,
granulometric, morphometric, petrographic and vegetation surveys suggest that talus slopes throughout the study area are at
different stages of development, with some ancient and recent
slope deposits. In the field, this results in steep and concave talus
slopes with fresh debris, but also coarse deposits with an openwork texture showing alteration and a more developed vegetation
cover. In addition, the opposite SW and NE sides exhibit significant
differences with respect to most of the parameters surveyed, which
means that their evolution occurred at different time scales.
Following the last deglaciation, paraglacial adjustment could
have enabled slope failures, initiating talus slope formation. Nowadays, evidence of periglacial processes (gelifraction, frost heave)
that have interacted with extensive bedrock jointing has been highlighted. As a result, features such as monoliths that have detached
from the basalt layer and are on the verge of falling, or notches
within the sedimentary rocks, show that rockfalls could occur at
any given time. However, dense vegetation cover, primarily in the
form of dense shrubs, has developed on some of the slopes and at
the base of the talus slopes, meaning that the slope process runout
is limited at the present time.
The present-day slope dynamic has been documented with
the use of automatic cameras over a one-year period during the
2017–2018 winter season. The reworking of snow avalanches on
slope deposits appeared to be a significant factor in the redistribution of debris, especially in the spring when the wet snow avalanche
deposits mostly consisted of dirty snow. Discrete rockfalls have
occurred during the same period, and some of the fallen debris have
been transported down the snow-covered talus slope.
Acknowledgements
Funding for this project was provided by the Natural Sciences
and Engineering Research Council (NSERC), LabEx DRIIHM and
OHMi NUNAVIK-TUKISIG, and IPEV program DeSiGN. The authors
want to thank Félix Faucher, Julien Lebrun and Thorsteinn Saemundsson for their valuable help in the field and the Centre d’études
nordiques (CEN) for its logistical support.
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