Talus slope evolution under the influence of glaciers with the example of slopes near the Hans Glacier, SW Spitsbergen, Norway

Accepted Manuscript Talus slope evolution under the influence of glaciers with the example of slopes near the Hans Glacier, SW Spitsbergen, Norway

Krzysztof Senderak, Marta Kondracka, Bogdan Gądek PII: DOI: Reference:

S0169-555X(16)30938-2 doi: 10.1016/j.geomorph.2017.02.023 GEOMOR 5941

To appear in:

Geomorphology

Received date: Revised date: Accepted date:

28 September 2016 24 February 2017 25 February 2017

Please cite this article as: Krzysztof Senderak, Marta Kondracka, Bogdan Gądek , Talus slope evolution under the influence of glaciers with the example of slopes near the Hans Glacier, SW Spitsbergen, Norway. The address for the corresponding author was captured as affiliation for all authors. Please check if appropriate. Geomor(2017), doi: 10.1016/ j.geomorph.2017.02.023

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ACCEPTED MANUSCRIPT Talus slope evolution under the influence of glaciers with the example of slopes near the Hans Glacier, SW Spitsbergen, Norway Krzysztof Senderak a, Marta Kondracka b, Bogdan Gądek a a

Faculty of Earth Sciences, Department of Geomorphology, University of Silesia, Będzińska str. 60, 41-200 Sosnowiec, Poland

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Faculty of Earth Sciences, Department of Applied Geology, University of Silesia, Będzińska str. 60, 41-200 Sosnowiec, Poland

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Abstract

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On Spitsbergen, which is 60% glaciated, talus slopes have frequently developed in interaction with

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glaciers, which had an influence on the evolution of the internal structure of slopes. This paper presents the results of geophysical surveys (electrical resistivity tomography – ERT and ground-

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penetrating radar – GPR) of the talus slopes near the Hans Glacier (SW Spitsbergen). The aim of investigations was to compare the talus slopes under the influence of glaciers in two different parts

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of the area in order to reveal differences in their internal structure. We assumed that different locations of talus slopes can have an influence on the slope structure, showing different stages of

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evolution of the talus slopes. The maximum thickness of studied slopes ranges from 20 m in a

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marginal zone of the glacier, to up to 35 m without contact with the glacier. Permafrost begins at a depth of 2-3 m and can develop until bedrock is reached. The internal structure of these talus slopes

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contains glacial ice, which is covered by a layer of slope material with a thickness from a few to up to 10 m. The buried glacial ice is slowly melting simultaneously with the deglaciation of the area but can

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remain in the structure of the talus slopes for much longer. Morphogenetic processes, such as avalanches, rockfalls, and debris flows are most visible until the glacial ice is completely melted within the internal structure of the slope. Based on the geophysical and geomorphological data, general models were proposed for the early stages of evolution of talus slopes in valleys under deglaciation.

Keywords: talus slopes; periglacial zone; ERT and GPR surveys; permafrost

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ACCEPTED MANUSCRIPT 1. Introduction Talus slopes occur in all climate zones; however, they do not develop in the same way everywhere (Albjär et al., 1979). These places are sensitive to changes in climatic conditions; therefore information about their evolution is stored on the surface and in the internal structure of the talus slopes (Iturrizaga, 2012). The process of deposition of slope material depends on many

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factors but mainly on the size of the area of sediment supply, the intensity of weathering, the type of material transport, and also the water content in the environment (Rapp, 1960; Jahn, 1967; Selby,

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1993). The greatest dynamics of morphogenetic processes are found in landforms in high mountains

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and glaciated areas (Coque and Raynal, 1970; André, 1997; Kochel and Trop, 2012). The talus slopes

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on Spitsbergen are an inseparable element of the contemporary landscape of the island and occur within almost all-mountain massifs (Åkerman, 1984). These places have been, or still are, moulded by

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the neighbouring glaciers.

The first studies of talus slopes were focused on processes occurring on their surface. Some

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crucial studies concerned the formation of Spitsbergen’s slopes and the dominant processes on them

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(e.g., Jahn, 1960, 1967; Rapp, 1960; Dewolf, 1970; Luckman, 1977; Åkerman, 1984). The estimation of the retreat rockwalls of these forms was extensively discussed (André, 1986, 1988, 1997). Numerous groups of researchers have been interested in the dating of glacial/periglacial landforms

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and processes (André, 1990a, 1990b; Étienne et al., 2008; Moreau et al., 2008; Owczarek et al.,

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2013). Blikra and Nemec (1998) described in detail the sedimentology of talus slopes based on the example of fossil deposits from northern Norway. Contemporary factors causing the deposition of slope material in the periglacial environment are identified by Mercier et al. (2009) and De Haas et al. (2015). With the development of geophysical methods for studying slopes, improved measurement techniques enabled their internal structures to be recognised. Pioneering and well-known research work mainly comes from the area of the Alps (Sass, 2006, 2007). Research on methodologies suggests that electrical resistivity tomography (ERT) and ground-penetrating radar (GPR) methods are the best approach to analysing talus slopes with dominant debris material (Otto and Sass, 2006; 2

ACCEPTED MANUSCRIPT Schrott and Sass, 2008; Van Dam, 2012). Despite the difficulties related to occurring permafrost, talus slopes in periglacial areas are more and more frequently studied using a set of such methods (Haeberli et al., 2010; Haeberli, 2013). In Svalbard, the GPR method was used to identify a pushmoraine (Lønne and Lauritsen, 1996), pingo (Ross et al., 2005), as well as ice and soil wedges (Watanabe et al., 2013). Until recently, the application of the ERT method was limited to studies of

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rock glaciers (Berthling and Juliussen, 2008) and also to estimating the retreat rates of Arctic rockwalls (Siewert et al., 2012), as well as to use in the imaging of areas affected by permafrost

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(Kasprzak, 2015; Kasprzak et al., 2016).

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This article presents the results of geophysical research (ERT and GPR) and geomorphological

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observations made in September 2015 on two talus slopes located near the Hans Glacier on SW Spitsbergen (Fig. 1). The slopes have different exposition, height, size, inclination, and size of

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sediment supply area and lithology (Table 1). In addition, these forms are currently at different distances from the glacier, which implies that they can be at different stages of evolution. Previous

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theories about the early development of the talus slopes were not based on geophysical data (e.g.,

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Ballantyne, 2002). The aim of the study was to apply ERT and GPR methods in places where talus slopes interact with glaciers and was also the creation of a model of the early evolution of talus

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slopes for landforms developing in the glacial valleys on Spitsbergen.

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Fig. 1. Simplified geomorphological map of the study area with localization of geophysical profiles (based on maps by Karczewski et al.

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(1984), extent of ice cliffs based on aerial photos (1936) by Norwegian Polar Institute, and satellite image (2011) by NPI / USGS Landsat 2011). Symbols: 1 – rock slopes, 2 – debris-mantled slopes, 3 – glacier, 4 – lateral moraine, 5 – moraine deposits, 6 – plains of raised marine

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terraces, 7 – see extent of tidewater glacier in the indicated year, 8 – rivers, 9 – geophysical profiles.

Table 1

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A characteristic of talus slopes selected to the geophysical surveys Fugleberget

Fannytoppen

(profiles 1 and 2)

(profile 3)

NE

NW

Inclination

35°

50°

Length of slope

245 m

175 m

Distance to glacier

0m

approx. 500 m

Altitude

190 m asl

175 m asl

Dominant lithology

amphibolites

phyllites, schists

Name of slopes

Exposure

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ACCEPTED MANUSCRIPT 2. Study area 2.1. Regional setting The study area is located around Hornsund Fjord on southwest Spitsbergen. The investigations were made in the northern part of the fjord, on slopes near the Hans Glacier where glaciological observations have been carried out for many years based on the logistic facilities of the

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Stanislaw Siedlecki Polish Polar Research Station. The glacier covers an area of over 50 km2 and flows toward the south. The snout of the glacier is located partially in the water of the fjord, which causes

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intense calving of the glacier (Błaszczyk et al., 2013). The Hans Glacier fills a deep valley oriented N-S,

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and the talus slopes have been developing along the entire glacier within the all-mountain massifs

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(Jania, 1998). The studies were carried out on both sides of the valley and directly under the peaks of Fugleberget 569 m asl (profiles 1 and 2) and Fannytoppen 390 m asl (profile 3) because these places

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have been under the influence of the glacier and, thus, can have different stages of evolution (Fig. 1). The vast majority of the area surrounding Hornsund Fjord consists of crystalline rocks of Proterozoic

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Eon (Czerny et al., 1993). These are mainly metamorphic rocks developed during the last stage of the

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orogeny of western Spitsbergen (Harland, 1997). The basement under the Hans Glacier is characterised by a varied lithology: from amphibolites on the west side to phyllites on the east and thin layers of sedimentary rocks (Birkenmajer, 1990).

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2.2. Outline of geomorphology

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The evolution of the Spitsbergen landscape largely depends on the geology of the area and climatic conditions (Zwoliński et al., 2013). In the vicinity of the Hans Glacier, typical landforms of the periglacial zone are being formed (Fig. 1). The high-mountain morphology clearly dominates in the study area, therefore the talus slopes and rock slopes are the most common and dynamic depositional environment. A marine terrace system including 15 levels from 2 up to 230 m asl occurs on the northern side of the Hornsund Fjord (Karczewski et al., 1984; Lindner et al., 1991). The regression of the Hans Glacier, apart from the retreat of ice cliffs, caused the exposure of moraine deposits and the deposition of wide lateral moraine ridges with buried glacial ice (Jania, 1988). 5

ACCEPTED MANUSCRIPT In the eastern part of the Fugleberget massif, from the highest areas down to a height of about 200 m asl, the rockwalls are cut by a system of deep chutes and corrosive gullies. Transport of the material takes place along these paths and leads to the formation of talus slopes (Nitychoruk and Dzierżek, 1988). When at maximum extent, the glacier rested against the eastern slopes of Fugleberget, leaving lateral moraine deposits and a pronounced trimline. Currently talus slopes

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develop above the trimline, and during extreme events their material shapes the surface of the ridge formed by the lateral moraine of the glacier (Fig. 2).

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Mountain massifs on the eastern side of the glacier are built of less resistant rocks

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(Birkenmajer, 1990). As a result, the slopes of the mountains are less gullied, and spaces are more

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open and subject to the formation of extensive talus slopes from the summit (Jania, 1998). An example of this is the Fannytoppen massif and nearest peaks, which, at least on the surface, are built

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of phyllites and amphibolite shales with an orientation of ~110/50 (dip angle of 50°, dip direction SE). In connection with the dip direction on the opposite side of the massif, the western slopes of the

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Fannytoppen are very steep (to 50°). Such conditions promote intensive slope processes and the

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formation of talus slopes (Chandler, 1973). 2.3. Climate

The Hans Glacier is one of the best monitored glaciers on Spitsbergen as a result of the

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measurement series recorded by observers from the Polish Polar Station in Hornsund, the longest

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series of such records on Spitsbergen. For this area the average multiyear air temperature for 19792009 was -4.3°C: the warmest month is July with a temperature of +4.4°C, while the coldest month is January with a temperature of -10.9°C (Marsz, 2013a). The highest rainfall was recorded during the summer with an annual average of 434.4 mm (Łupikasza, 2013). The humidity value for that period was almost 79% (Marsz, 2013b). The precipitation and humidity are highly dependent on the position of the station by the Hornsund Fjord. The nonglaciated areas in the vicinity of the Hans Glacier are affected by permafrost (Harris et al., 2009; Dolnicki et al., 2013; Kasprzak et al., 2016), which is

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ACCEPTED MANUSCRIPT treated as a state of a ground with temperature below 0°C occurring in a period of 2 years (Dobiński,

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2011).

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Fig. 2. Study sites selected for geophysical investigations (A-D) and the examples of processes in the study area (E-F): (A western part of the Hans Glacier with marked trimline and profiles 1 & 2; (B ERT measurements under Fugleberget; (C eastern part of the study area with

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ACCEPTED MANUSCRIPT marked profile 3 (photos by KS); (D GPR measurements under Fannytoppen (photo by MK); (E outcrop of buried glacial ice in slopes of Fugleberget (photo by A. Piwowarczyk); (F deposition of coarse slope material directly on glacier’s surface (KS); (G internal structures in buried glacial ice under Fugleberget (AP); (H deposition of fresh slope material on lateral moraine deposits under Fannytoppen (MK).

3. Methods 3.1. Morphometric measurements and mapping

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Cartographic materials, such as topographic maps, orthophotomaps, and aerial photographs

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of the area of the Hans Glacier were used in order to select the precise talus slopes to be used for research (e.g., Karczewski et al., 1984; Jania et al., 1992, 2002; Kolondra, 2013). The precise

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morphometric measurements, such as length, width, and relative height, were calculated using

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ArcGIS tools (Table 1).

The interpretation of surface structures and subsurface images was performed based on

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additional photographic and descriptive material, which was collected during the measurements. This was supplemented by field data on the geomorphology and sedimentology of the slopes.

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The locations of the analysed profiles were recorded by a precise GPS receiver. Slope

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inclination was measured using a manual inclinometer, and following this, the measurement was tested in ArcGIS.

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3.2. Electrical resistivity tomography (ERT) The ERT measurements were carried out using the LUND automatic electric imaging system

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with an SAS (Signal Averaging System) 4000 Terrameter produced by the Swedish company ABEM. The system consists of an SAS 4000 Terrameter, electrode selector, 41 electrodes connected with multicore cables, and a battery. The system automatically chooses the appropriate electrodes: the two outer current electrodes and two inner potential electrodes. The device introduces current between the outer electrodes and measures the response voltage at discrete time intervals (ABEM, 2009). A Wenner-Schlumberger type of arrangement was used with 5-m spacing between electrodes, which is moderately sensitive to horizontal and vertical structures (Loke, 2004) and is 9

ACCEPTED MANUSCRIPT dedicated to detecting horizontal structures (Samouëlian et al., 2005). The apparent resistivity () of the Wenner-Schlumberger array was calculated by the Terrameter using the equation:

 =  n (n + 1) a R where R is the measured resistance, a is the spacing between potential electrodes, n is the ratio of the distances between the current and potential electrode and parameter a. The factor n was set at 1

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and increased to 2, 3 … until 8 for all profiles.

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At the Fugleberget massif, two profiles were combined using the roll along method to conduct measurements along the whole slope in order to present the structure of talus on one

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profile (profile 1 – total length 245 m). Profiles 2 and 3 were contacted using 36 electrodes, which

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resulted in a length of 175 m.

The collected data was interpreted in RES2DINV software, which calculates the true

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resistivity and true depth of the ground from the apparent resistivity measured using a Jacobi matrix calculation and forward modelling procedures. Models were provided by the smoothness-

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constrained least-squares method based on a quasi-Newton optimisation technique, which attempts

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to reduce the differences between the calculated and measured values of apparent resistivity (ABEM, 2009). These differences are indicated by the root-mean-squared (RMS) error, which is decreased together with another iteration (repeating a process of inversion). The RMS error for the

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5th iteration was 4.9 (profiles 1 and 2) and 3.3 (profile 3). The resulting inversion model was based on:

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298 points (profile 1), 248 and 192 points (profiles 2 and 3 respectively). In the topographical modelling, the distorted finite-element grid with damped distortion was used (damping factor 0.75). In polar regions geoelectrical surveys are difficult to organize because of low temperatures and logistic constraints related to the transport of the research team and equipment (in this study, a motorboat had to be used). The errors in measurements can be caused by empty space in the slope material with debris and a block fraction (Hilbich et al., 2009); therefore, the position of the electrodes has to be modified and corrected to provide a good electrical contact. 3.3. Ground-penetrating radar (GPR) 10

ACCEPTED MANUSCRIPT In GPR surveys, the transmitter propagates electromagnetic waves with a specific frequency in the medium. The waves are reflected from the boundaries of layers, discontinuities, and internal structures and then are returned to the receiver. The velocity of propagation of the waves depends on the dielectrical properties of the medium. The GPR measurements were conducted using the RAMAC GPR system (Malå GeoScience) with unshielded antennae at a frequency of 30 MHz.

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The profiling GPR was carried out along two selected slopes (profiles 1 and 3). The signal was propagated at an interval of 1 second. Time windows were in the range of 680-760 ns. Typical values

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of dielectric permittivity (ɛ) were used to process GPR data for the materials occurring in the

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periglacial zone (e.g., Neal, 2004). They showed respectively: profile 1 – coarse slope deposits and

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glacial ice ɛ = 9 (v = 0.10 m/ns); profile 3 – fine deposits and glacial ice ɛ = 3 (v = 0.17 m/ns) (Berthling and Melvold, 2008). The values of the parameters were dependent on the localisation of the internal

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structures in the slopes. The processing steps and interpretation of data were conducted in Rad Explorer software produced by DECO Geophysical Ltd wherein a DC-shift correction, a static

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correction of the first onset times, a band pass filter, and a runtime-dependent amplification were

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applied.

4. Results

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4.1. Fugleberget – profile 1 (longitudinal section)

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The longitudinal ERT profile of Fugleberget’s eastern slope indicates that the internal structure is divided into two parts, the upper talus slope (0-100 m) and the lower talus slope below the trimline (100-245 m). Lower values of electrical resistivity in the range of 4-20 kΩm were diagnosed in two zones of the profile (Fig. 3). The first zone occurs along the whole profile to a depth of 2-3 m and only consists of slope material or mixed moraine-slope material. The second zone is located at a depth of 20-25 m in the upper part of the profile and includes strongly weathered bedrock. Higher values of electrical resistivity are also recognised in the upper part of the profile at a depth of 3-20 m and occur in the range of 20-60 kΩm, which corresponds to permafrost. In this part 11

ACCEPTED MANUSCRIPT of the profile, the pore ice is only developing in slope material. In the lower part of the profile, higher electrical resistivity is found at a depth of 3-10 m and performs in the range of 60-750 kΩm, which is a typical value of the electrical resistivity of glacial ice (Kneisel and Hauck, 2008). The GPR profile confirms the bipartite nature of the internal structure (Fig. 4). The zones of the talus slope and the glacial influence are divided by the remnant of the trimline. In the first zone,

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in the subsurface layer of the slope, the internal structures (including slope material) were distinguished up to a depth of 2-5 metres. Within this material, we detected the occurrence of

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boundaries on the whole GPR profile between the active layer and permafrost table at a depth of 2-3

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metres. Material within which the electromagnetic waves are well propagated, was found below the

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trimline, which confirms the presence of glacial ice in the internal structure. The GPR profile allows us to determine that the thickness of this ice is in the range of at least 20-30 m. The ERT measurements

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are compatible with the thickness of the talus slopes being up to 20-25 m.

Fig. 3. Profile 1, ERT method, longitudinal section of Fugleberget’s eastern slopes.

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Fig. 4. Profile 1, GPR method, longitudinal section of Fugleberget’s eastern slopes.

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4.2. Fugleberget – profile 2 (cross section)

The ERT cross section allows us to describe in detail the internal structures in the lower part of Fugleberget’s eastern slopes (Fig. 5). On profile 2 we recognized two layers with different values of

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electrical resistivity. Based on the ERT survey, the first layer was assigned from the surface of slope to a depth of 10 m. Within this material the electrical resistivity increases from 4 kΩm at the surface up to 175 kΩm in the bottom of the layer. The field observations and measured values of electrical

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resistivity indicate moraine and slope material, in which the active layer occurs down to a depth of 3

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m. Below this, permafrost occurs to a depth of 10 m. The second layer in the profile has values of electrical resistivity in the range of 175-750 kΩm, which are typical for glacial ice. Based on the profile, we calculated that the thickness of the ice is at least 20 m. The bottom of the ice was not detected in the ERT measurements. In profile 2 we did not carry out a GPR survey because the ERT measurements of the lateral moraine of Hans Glacier had indicated a simple and predictable model of the internal structure.

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Fig. 5. Profile 2, ERT method, cross section of Fugleberget’s eastern slopes.

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4.3. Fannytoppen – profile 3 (longitudinal section)

Profile 3 is similar to profile 1, which is on the other side of the Hans Glacier. Within

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Fannytoppen’s slope, the evidently bipartite internal structure is visible. The study shows that two

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zones can be distinguished: the zone of talus slopes and glacial influence (Fig. 6). Lower values of electrical resistivity in the range of 2-8 kΩm are recorded in the upper part of the slope. These are located from the surface to a depth of 10 m and also inside the slope below a depth of 35-40 m. A

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zone of higher electrical resistivity in the range of 10-50 kΩm occurs between these locations, which continues along the slope. Glacial ice (which also outcrops at the surface) was only recognised in the lower part of slope with values of electrical resistivity up to 365 kΩm.

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The GPR profile confirms the assumption of variable internal structure in the upper part of the slope (Fig. 7). Based on the GPR survey, we detected several regular horizons, which have an

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inclination similar to the inclination of the slope surface. In the surface layer, the electromagnetic waves are propagated well to a depth of 10-15 m. The subsurface image looks similar to the lower part of the slope in which glacial ice lies at the surface. The thickness of the talus slope is from 8 m in the lower part of the slope up to even 35 m in the upper and central parts.

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Fig. 6. Profile 3, ERT method, longitudinal section of Fannytoppen’s western slopes. Please note the different scale values of electrical

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resistivity than in profiles 1 and 2 (Figs. 3 and 5).

Fig. 7. Profile 3, GPR method, longitudinal section of Fannytoppen’s western slopes.

5. Interpretation and discussion 15

ACCEPTED MANUSCRIPT The talus slopes develop on the all-mountain slopes surrounding the Hans Glacier. In general, those forms occurring near existing glaciers are characterised by high activity of morphogenetic processes (Ballantyne, 1995; Chiarle et al., 2007). These nonglacial processes conditioned by glaciers were defined using the term paraglacial (Church and Ryder, 1972; Slaymaker, 2009); therefore the talus slopes in the deglaciated areas of Spitsbergen develop in a paraglacial sequence of slope

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evolution (Ballantyne, 2002). In marginal zones of the glacier, inside the glacial valley, interaction occurs between the mountain slope and the glacier (Bennett, 1999; Curry, 1999). Any movements of

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the glacier (connected with transgression or regression) directly influence the depositional

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environment of the slopes (Rachlewicz, 2010). A mechanism of interaction also depends on the

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influence of slopes on processes of the glacial system (Jania, 1998). One result of coexistence in such different environments and processes is the diversity of the periglacial landscape (Kochel and Trop,

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2012). Apart from the information about the maximum range and history of glaciations, these places are a valuable source of knowledge about the development of older forms in valleys, which were

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earlier exposed by the glacier (Blikra and Nemec, 1998; Iturrizaga, 2012).

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The studies of the western part of the area illustrate an interaction of environments: talus slopes and the Hans Glacier. The conducted field observations indicate that at the moment of greatest transgression, the glacier reached a point on the slope at a height of 145 m asl (level of

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trimline). The Hans Glacier, in a similar manner to the majority of Spitsbergen’s glaciers, has been in a

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retreat stage since the termination of the Little Ice Age (LIA) at the end of nineteenth century (Hagen et al., 1993; Rachlewicz et al., 2007; Błaszczyk et al., 2009). The change in the position of the glacier snout was ~1800 m in the period 1936-2010 (Kolondra, 2013). At present the glacier is located at a height of 95 m asl near the Fugleberget. The trimline is clearly visible in the field. The internal structures associated with the maximum range of the glacier are visible on profile 1. The ERT and GPR surveys indicate that after glacial recession, in addition to lateral moraine deposits, some glacial ice remained. Ice is the dominant medium in the internal structure of Fugleberget’s slopes. Based on the observations and studies, we believe that the ice has been covered by debris and block material 16

ACCEPTED MANUSCRIPT in various slope processes, such as avalanches, rockfalls, and debris flows. Currently, the ice is located under a cover of slope material with a thickness of 5-10 m. The structure of the ice reaches several metres farther inside the slope than the trimline would suggest. This geometry is related to the shape of the glacier, whose marginal zones are lower than the central part (Grabiec et al., 2012). This causes the surface of the glacier to fall to the west, which has been established in the shape of

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the structure of the covered ice lying within the slope. The part of the slopes of Fugleberget above the trimline are still very active, which is

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confirmed by numerous traces of fresh material supply, an unstable surface of the slopes and an

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irregular trough after debris flows (De Haas et al., 2015). The thickness of the slope material

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determined on the basis of ERT and GPR is up to 25 m (Table 2). The bedrock beneath exhibits the characteristics of a medium with lower values of electrical resistivity than the internal structure of

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the talus slope. The reason for this is provided by several metres of sediment layer affected by permafrost, whose range begins at a depth of 2-3 m. Above this, the material of the active layer does

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not develop pore ice (Leszkiewicz and Caputa, 2004). The ERT measurements of talus cones in

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Longyeardalen, central Spitsbergen, deriving from research conducted by Siewert et al. (2012) concerning Arctic rockwall retreat rates, show similarities to the study site. The internal structure of slopes affected by permafrost confirm the higher resistivity of unconsolidated slope material and

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lower resistivity of bedrock, but measured values of these slopes are significantly lower (in the range

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of 10 kΩm) than in the vicinity of the Hans Glacier. Measurements in the eastern part of the area were carried out in front of the snout of the Hans Glacier. Currently, the glacier has not directly influenced the talus slope at the location of profile 3, as it has in the case of profiles 1 and 2. On the slope surface the trimline is not visible, as it was either destroyed or covered. Glacial ice was found in the lower part of the profile and is exposed on the surface of the layer of fine slope material. In the internal structure, the ice also plunges to the inside slope, referring to the original shape of the surface of the glacier. These observations suggest

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ACCEPTED MANUSCRIPT that it is a later stage in the evolution of the slope where the deep ice gradually melts and disappears in the internal structure. The upper part of profile 3 exhibits the characteristics of a more mature slope. Compared to the other talus slopes in the research area, the slope material of this profile is better sorted. Phyllites building the western slopes of Fannytoppen are subject to much faster weathering. As a result, the

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morphogenetic processes are still very frequent but limited mostly to the gravitational movement of material down the slope during the summer and transported by avalanches in winter and spring

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(Luckman, 1977; André, 1990a, 1990b; Eckerstorfer and Christiansen, 2011). Avalanches dominate

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talus slopes’ processes in the Arctic (De Haas et al., 2015). The result of this is often a smooth slope

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surface, as in the case of profile 3 (Nitychoruk and Dzierżek, 1988).

The thickness of the talus slopes of Fannytoppen reaches 30-35 m. Such a large amount of

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slope material may be explained by the older age of the slopes relative to the study slopes in the western part of the study area. The internal structure is similar to profile 1. The surface layer, up to

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2-3 m, occurs in the absence of pore ice. Below this level the material is covered by permafrost, the

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extent of which continues along the entire structure of the slope. The same fact has been observed on the talus slopes in western Svalbard (Humlum et al., 2003; Åkerman, 2005). This is confirmed by the higher electrical resistivity of the profile between the level of 2-3 m and 30-35 m. The bedrock

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visible in the profiles of ERT and GPR is the background for permafrost. The significantly lower

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electrical resistivity of the bedrock confirms that it is composed of low resistance phyllites. The GPR surveys of profile 3 allowed the dissection of the other internal structures. In the upper part of the slope, at least three horizons are visible with an orientation consistent with the modern surface of the slope. Clear horizons in the internal structure are most commonly associated with the record of historical debris flows, which leave behind gullies with typically regular bottoms (Sass and Krautblatter, 2007). The remains of the gullies can be seen in the GPR image (Sass, 2008). The presence of debris flows in the earlier stages of the development of the talus slope may suggest similar climatic conditions to those currently prevailing, which are associated with an interglacial 18

ACCEPTED MANUSCRIPT period (Gądek et al., 2009, 2016). The debris flow gullies are located in the upper part of the slope, which may indicate that the debris flows only occurred in the higher slope zone, whereas their bottoms were the basis for all of the transported material on the talus slope. Based on the results of the ERT and GPR surveys, as well as observation of geomorphological forms on the surface of the slopes, we can reconstruct the process of initial development of these

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structures on the area covered by the Hans Glacier (Fig. 8). Talus slopes almost certainly were present in the landscape of Spitsbergen before the glaciations that we observe today (André, 1997).

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The original forms were significantly smaller and formed in the direct vicinity of rock outcrops (Fig.

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8A). Similar initial landforms were studied in the Western Grampian Highlands, Scotland and Mynydd

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Du, UK (Brazier et al., 1988; Curry and Morris, 2004).The development of these structures would follow later, depending on the climatic conditions appropriate for warmer or cooler environments

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(Albjär et al., 1979). However, the cooling of the climate and covering of an area of Spitsbergen by glaciers caused the formation of the forms occurring now, which have their own unique and

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distinctive features typical of the talus slopes present in the climatic conditions of the Arctic (De Haas

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et al., 2015).

The transgression of the Hans Glacier destroyed the initial talus slopes (Fig. 8C). The glacier filling the valley eroded the ground directly under the glacier, but also on the surrounding mountain

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ranges (Błaszczyk et al., 2013). Until the maximum range of the glacier, which defines the trimline,

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slope material was continually being deposited on the surface of the ice (Ballantyne and Benn, 1996). Then the deglaciation of the area started and the process of exhumation of the bedrock from under the glacier took place (Fig. 8D). The high rate of regression caused the isolation of glacial ice covered with slope material from the main mass of the glacier (Ballantyne, 2002). The marginal zones stuck in the structure of the talus slopes, whose development has been accelerated by the lifting of the glacial pressure on the bedrock and the start of morphogenetic processes within the sediment supply area (McColl, 2012), was also observed in Chilean Patagonia (Harrison and Winchester, 1997). The slope under Fugleberget (profiles 1 and 2) is currently at this stage. The continuation of deglaciation 19

ACCEPTED MANUSCRIPT and regression causes talus slopes to lie far from the glacier and reduces its impact on the immediate surroundings. Ice trapped beneath the slope material begins to melt (Fig. 8E). The process of ice disappearance causes deformation of the internal structure (Blikra and Nemec, 1998). Over time, changes in the structure are compensated for by a fresh supply of material and accumulation of this material on the surface layer a talus slope, which is typical for paraglacial processes (Fig. 8F). The last

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stage presently specifically concerns the slope of Fannytoppen (profile 3). The regression of the glacier and simultaneous melting of buried glacial ice in talus slopes

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decrease the role of paraglacial processes in slope evolution. Then the factors responsible for typical

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weathering in the periglacial zone, which had had a secondary importance while the glacier was

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active, such as for example exposure and insolation (Åkerman, 1984), will start to dominate in the

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development of mature talus slopes.

Fig. 8. Early stages of evolution of talus slopes near the Hans Glacier on Spitsbergen. Contemporary stage of development of Fugleberget’s slopes is imaged on sketch (D), whereas, Fannytoppen’s slopes are presented in (F).

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Table 2 Comparison of the results of surveys on the talus slopes near the Hans Glacier Fugleberget

Fannytoppen

(profiles 1 and 2)

(profile 3)

Thickness of slope

20-25 m

30-35 m

Kind of material

Slope and moraine

Slope

Buried glacial ice

Yes, below trimline (145 m asl)

Yes, in the lower part of the slope

Active layer

2-3 m

2-3 m

Permafrost

3-20 m

3-30 m

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Name of slopes

6. Conclusions

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The studies of talus slopes near the Hans Glacier confirm that geophysical methods are useful in identifying internal structures. A set of ERT and GPR methods provide great accuracy in

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determining the distribution of structures along the slope.

The early evolution of the talus slopes in glacial valleys depends on interaction with the

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glacier. The initial depositional forms are destroyed by ice and replaced by new slope material, which

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is formed synchronously with the occurrence of the transgression of the glacier. In the marginal zones of the glaciers, material enters the lateral moraine, which is marked by the trimline. The most

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important factors shaping the talus slopes near the Hans Glacier on Spitsbergen, apart from interacting with the glacier, are the size of the sediment supply area and lithology. In the polar

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environment, exposure and insolation are of secondary importance. The activity of paraglacial processes (conditioned by glaciation) is the most intensive by the time glacial ice is completely melted within the internal structure of the slope, which causes stabilisation of the mechanisms responsible for redeposition of slope material and the pursuit of a straight longitudinal profile of the slope in typical climatic conditions of the periglacial zone. The maximum thickness of talus slopes in the vicinity of the Hans Glacier ranges from 20 m, in the marginal zone of the glacier, up to 35 m without contact with the glacier. The internal

21

ACCEPTED MANUSCRIPT structure of these talus slopes contains buried glacial ice underneath the slope material with a thickness from a few to up to 10 m. The position of the glacial ice in the slope depends on the stage of slope evolution and occurs below the trimline or in the lower part of the slope. Permafrost develops at every stage of the evolution of the slopes, regardless of the presence of glacial ice in the internal structure of the slopes; whereas the active layer occurs to a depth of 2-3 m, which is caused

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by thick unconsolidated near-surface sediments with debris and block fractions. The development of talus slopes and the dominant morphogenetic processes in the

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periglacial zone on Spitsbergen are similar to alpine areas where the interaction between the talus

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slope and glacier takes place now or has occurred in the past. Studies are needed of talus slopes

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using noninvasive geophysical methods to be further developed coupled with methods of absolute

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dating, which would reconstruct the timeline of evolution of the slope.

Acknowledgments

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The article was prepared as a part of the doctoral project conducted by KS at the Centre for Polar

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Studies, University of Silesia, Poland. This work was supported by funds from the Leading National Research Centre (KNOW) received by the Centre for Polar Studies for the period 2014-2018. We especially thank the editor Richard A. Marston, Dr. Tjalling de Haas and two anonymous reviewers

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whose suggestions and comments helped to improve the paper. We are also grateful to Marta

the field.

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Bystrowska and Paweł Pilch for help in translation and Arkadiusz Piwowarczyk for his assistance in

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ACCEPTED MANUSCRIPT Highlights

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 

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

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The geophysical surveys (ERT and GPR) of talus slopes at the different stages of early evolution were conducted Internal structure of talus slopes contains the buried glacial ice which is slowly melting simultaneously with the deglaciation of the area Maximum thickness of talus slopes ranges from 20 metres in a marginal zone of the glacier, to up to 35 metres without contact with the glacier Permafrost begins at a depth of 2-3 metres and can develop until bedrock is reached General models were proposed for the early stages of evolution of talus slopes in valleys under deglaciation

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