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Structure and genesis of the Thabor rock glacier (Northern French Alps) determined from morphological and ground-penetrating radar surveys
Geomorphology 134 (2011) 269–279
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Geomorphology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / g e o m o r p h
Structure and genesis of the Thabor rock glacier (Northern French Alps) determined from morphological and ground-penetrating radar surveys Sébastien Monnier a,⁎, Christian Camerlynck b, Fayçal Rejiba b, Christophe Kinnard a, Thierry Feuillet c, Amine Dhemaied b a b c
Centro de Estudios Avanzados en Zonas Áridas (CEAZA), Campus Andrés Bello, Casilla 599, Raúl Bitrán s/n, La Serena, Chile UMR 7619 Sisyphe, Université Paris 6 Pierre et Marie Curie, 4 place Jussieu, 75252 Paris Cedex 05, France UMR 6554 LETG, Université Paris Sorbonne-Paris IV, 191 rue Saint Jacques, 75005 Paris, France
a r t i c l e
i n f o
Article history: Received 15 December 2010 Received in revised form 29 April 2011 Accepted 6 July 2011 Available online 31 July 2011 Keywords: Rock glacier Ground-penetrating radar Internal structure Glacier–permafrost interactions
a b s t r a c t Landform analysis and ground-penetrating radar (GPR) were used to investigate the Thabor rock glacier, in the Northern French Alps. The surface features of the rock glacier were classified and described, with emphasis on massive ice exposures. The retreat of the former Thabor glacier since the Little Ice Age (LIA) was documented through an analysis of historical sources, and recent movements of the rock glacier were inferred from orthophoto-based measurements. Two-dimensional (2-D) models of the radar wave velocity were derived from the raw GPR data, using the numerous diffraction hyperbolae for local determinations of the velocity and kriging interpolation techniques. Subsequently, the profiles were migrated through a 2-D Kirchhoff migration method using the interpolated velocities. The 2-D velocity models exhibit pronounced spatial variations and, in several locations, high values (N 0.15 m ns−1) potentially corresponding to massive ice. On the other hand, while the migrated profiles show numerous layers, the internal stratigraphy of the rock glacier is dominated by a few prominent internal boundaries. The integration of morphology, radar wave velocity, and internal stratigraphy allowed us to identify the main structural units of the rock glacier as well as to explain its genesis: the rock glacier was formed by the imbrication of a massive ice core, originating from the retreat of the former Thabor glacier since the LIA, into pre-existing glacial deposits. © 2011 Elsevier B.V. All rights reserved.
1. Introduction During the last decade, important insights into the geomorphology of high mountain regions have been gained from the use of groundpenetrating radar (GPR). Previous GPR studies have looked at various morphological features such as talus slopes, glacial deposits and, most commonly, rock glaciers. In particular, the advent of GPR technology in rock glacier research has been decisive. Due to both its complexity and inaccessible nature, the internal structure of rock glaciers is often unknown. Hence GPR is often considered an ideal tool for “taking a look” into the subsurface of rock glaciers. Previous GPR studies of rock glaciers have permitted the examination of the layered structure, the main internal boundaries, and the thickness of the rock glaciers studied (e.g., Isaksen et al., 2000; Hausmann et al., 2007; Monnier et al., 2008; Degenhardt, 2009). However, only rarely GPR data were correlated with direct observations of the internal structure (Fukui et al., 2008). Therefore, previous GPR studies were generally unsuccessful in defining the nature of the inner materials, especially in revealing the presence of ice. In this paper, we present work
⁎ Corresponding author. Tel.: + 56 51 334 869. E-mail address: [email protected] (S. Monnier). 0169-555X/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.geomorph.2011.07.004
conducted on the Thabor rock glacier, Northern French Alps. Taking into account the morphology, the recent glacial history of the site, and the occurrence of several direct observations of the shallow rock glacier structure, we process large and complex common-offset GPR data in order to both estimate the properties of the internal materials and highlight the rock glacier stratigraphy. We then use the combination of these data and observations to understand how the Thabor rock glacier has formed. 2. The Thabor rock glacier The Thabor rock glacier is located in the Mt Thabor Massif, Northern French Alps, south of the Vanoise Massif (Fig. 1). The Thabor Massif is an area with relatively moderate crest elevations (ca. 3000 m), and is known for being a regional dry pole — ~700 mm of annual precipitations in valley stations (Voiron, 1983; Kaiser, 1987). On the basis of meteorological data recorded by Meteo France in the Modane valley-station, and using a lapse rate of −0.6 °C 100 m −1, the 0 °C and −2 °C isotherms may be located at approximately 2460 and 2790 m, respectively. The Thabor rock glacier extends between 2900 and 2650 m, inside a north-facing amphitheater. Surrounding summits (Mt Thabor, 3178 m; Thabor Peak, 3207 m; and Cheval Blanc, 3020 m) are made of Triassic quartzites while the lower slopes
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Fig. 1. Location (a) and geological context (b) of the Thabor rock glacier. The coordinates are expressed in the UTM−WGS84 system (as for all subsequent maps). The geological context is depicted onto a 3-D Google Earth sketch. Q: quartzites. h: carboniferous schists. T.s.: Thabor synclinal. Q-h: geological contact between the quartzites and the carboniferous schists. The Google Earth sketch is 1.4 km wide.
are made of carboniferous schists, as a result of the differential erosion of the local synclinal structure belonging to the Zone houillère Briançonnaise (Gidon, 1977). 2.1. Morphology The Thabor rock glacier is one of the largest rock glaciers of the Northern French Alps, being more than 1 km long and up to 600 m wide, with a surface area close to 60 ha. These large dimensions and the light color arising from the quartzitic composition of the surface compensate for the relatively low height of the front (rarely more than 15 m) and give a striking appearance to the rock glacier, which is clearly visible from aerial view or from distant crests. The rock glacier surface, composed of 20–40 cm long clasts, exhibits several morphological features (Fig. 2). The central part of the rock glacier above 2750 m is depressed and has a poorly differentiated morphology. In the upper part, typical glacial features such as lateral morainic crests and hummocks are present; crevasse-like features reminiscent of true crevasses in glacier ice are also noticeable. Over the whole rock glacier, the two main morphological features encountered are arcuate ridges and long and straight ramps, which extend along both sides of the rock glacier. The arcuate and embedded ridges are 1–3 m high and
10–100 m long, and are mainly present in the lower part and on the western flank of the rock glacier. These often exhibit a dissymmetric cross profile, with a smooth, convex front slope and a steep back slope; the latter precedes a flat and depressed area. The long and straight ramps concentrate on the eastern flank of the rock glacier. Running over several hundred meters, they exhibit a steep, 1–2 m high slope facing the center-line depression and a gentle slope facing the rock glacier margin. The arcuate ridges, according to experiments made on rock glacier rheology (Kääb and Weber, 2004), suggest compression processes, with possible local downwasting on their back side, while the long and straight ramps suggest downwasting processes. Moreover, on its western and especially eastern flanks, the rock glacier connects with other deposits that exhibit morphological signs of permafrost debris deformation (ridges and frontal slopes). 2.2. Diachronic information The Thabor rock glacier is a recent glacial forefield. Indeed, several available documents (Fig. 3) show an important glacial retreat in the amphitheater since the middle of the 19th century. These include the socalled Carte d'État-Major (1843–1867), topographic maps from the beginning of the 20th century (based on late 19th century topographic
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Fig. 2. The Thabor rock glacier morphology and GPR survey. Background: IGN orthophoto of 2007 (as for all subsequent maps). Modified from Monnier, 2007.
surveys), a picture of the massif in a publication of Ferrand (1892), and topographic maps of the late 20th century. Nevertheless, integration of this information requires caution because of potential mapping errors in depicting both the topography (for maps of the 19th century) and the glacier extent (for all dates). It is thus necessary to adjust the information by seeking a coherency between the various documents and by taking into account the geomorphology of the rock glacier. According to the Carte d'Etat-Major (1843–1867), in the middle of the 19th century, which corresponds to the end of the Little Ice Age (Le Roy Ladurie, 2003), the Thabor glacier completely filled the amphitheater and extended beyond the actual limits of the rock glacier. However, according to the topographic data of 1890 and the picture in Ferrand (1892), it had severely reduced in size 40 years later, probably ending halfway up the actual rock glacier tongue. Little Ice Age (LIA hereafter) glacial features are usually well preserved in the area, especially in quartzite accumulations, but there are no glacial features in the amphitheater around the rock glacier terminus, where the slopes are quite gentle. Therefore, the glacial extent interpretation provided by the Carte d'Etat-Major (mid-19th century) is clearly exaggerated and must have resulted from approximations made in locating landforms and from the incorporation of snow patches and/or adjacent small glaciers. The most reasonable interpretation sets the mid-19th century extent of the Thabor glacier near the current rock glacier terminus and all subsequent extents within the current rock glacier margins. Finally, although the 1977 map shows the Thabor glacier as a small residual glacier occupying the upper reach of the rock glacier, aerial photos of the same decade suggest a subtly larger extent. Today, the Thabor glacier has completely vanished.
2.3. Exposures of massive ice Despite the disappearance of the Thabor glacier, buried glacier ice is still present at the site. The upper part of the rock glacier above 2800–2850 m is strongly depressed (~30 m) at its center, and exhibits remnants of a rapidly decaying debris-covered glacier, as demonstrated by the appearance of bedrock outcrops over the last 30 years (Fig. 2). Moreover, while walking down the rock glacier at the end of the warm summer of 2003, we recorded 24 exposures of massive ice in furrows, depressions or spaces between ridges usually occupied by snow (Figs. 2 and 4). In some exposures, the ice was banded. In the upper part (ca. 2800 m), but away from the central furrow, ice was found a few decimeters beneath the surface by means of manual digging. The lowest exposure of massive ice was found within an elliptic depression on the distal part of the rock glacier at 2730 m (pt. 24, Fig. 2). These ice exposures correlate with water temperatures near 0 °C measured in the spring in front of the rock glacier during the same late summer, in complete absence of snow on the upper part. 2.4. Working hypotheses about the Thabor rock glacier formation The evolution of the Thabor rock glacier over the last 150 years and the existence of a massive ice-core are not yet resolved. Monnier (2007) suggested a model of formation in which the Thabor rock glacier represents a transition between a rock glacier and a push moraine complex, following ideas of Haeberli (1979) about similar landforms in the region of the Grubengletscher, Switzerland. Thus, the morphology of the Thabor rock glacier would originate from glacier–
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Fig. 3. Historical documents providing information about the Thabor glacier retreat. a) “Carte d'État Major” (1847–1867). b) Picture of the massif, after a photograph, in Ferrand (1892). c) Topographic map from the beginning of the 20th century (topographic survey: 1890). d) IGN map of the late 20th century (topographic and photogrammetric surveys: 1977).
permafrost interactions: the LIA Thabor glacier advances would have created marginal arcuate ridges by compression/accumulation of frozen sediments; then, the glacier would have progressively decayed both by retreat and burial, leading to differentiated dynamics in the resulting debris-ice tongue — creep in zones with low ice wasting (distal part; amplification of the arcuate ridge morphology) and collapse in zones with pronounced ice wasting (central part; formation of the straight ramps). Such a scenario leaves unknown the origin of the forefield present before the LIA glacier advance.
3. Methods 3.1. Photogrammetric measurements Frontal movements of the Thabor rock glacier were calculated using an aerial photo of 1971 and an aerial orthophoto of 2006. Both photos were provided by the IGN (Institut Géographique National, France). The photo of 1971 was scanned and orthorectified. On the basis of 35 points selected throughout the two orthophotos and corresponding to stable topography (large blocks in flat areas, and outcrops of bedrock), the mean error for the orthorectification of the 1971 photograph is 3.6 m. Both orthophotos have a high spatial
resolution (1 m for 1971, and 0.5 m for 2006). Therefore, the base of the rock glacier front is easily recognizable and was digitally delineated. The displacements were calculated at 106 points sampled along the front.
3.2. GPR (ground-penetrating radar) data acquisition Common-offset GPR measurements were performed on the Thabor rock glacier using a pulseEKKO 100 system (Sensors & Software Inc.) with unshielded 50 MHz frequency antennae. The antennae were separated by 2 m, oriented perpendicular to the survey direction, and traces were recorded every 50 cm. Each recorded trace was the result of 16 stacks in order to increase the signal over noise ratio. GPR data were acquired along four profiles (Fig. 2). The longitudinal profile is 1450 m long and extends between 2675 and 2928 m above sea level (a.s.l.). Three transversal profiles, labeled from north to south #1, #2 and #3, have the following characteristics, respectively: 330 m long between 2735 and 2764 m a.s.l., 390 m long between 2769 and 2799 m a.s.l., and 660 m long between 2821 and 2853 m a.s.l. Elevation data were obtained with a differential GPS system with centimeter resolution. The preliminary processing of the GPR radargrams included zero-time correction, the removal of the DC shift component and that of the very low frequency component using a highpass filter (i.e. dewow filtering).
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Brandt et al., 2007). In the case of a tabular subsurface and for small offsets compared with depth, the stacking velocity is a good approximation of the root mean square (RMS) velocity between the antennae and the hyperbola apex. Whereas ‘true’ interval velocities could be retrieved from stacking velocities (Dix, 1955), in this study the latter were used for migration, and thus the conversion was not performed. Furthermore, referring to the aim of comparing the radar wave velocity with the superficial ice exposures, the shallow RMS velocity is close to the actual interval velocity. That being said, repeated hyperbola fittings throughout a GPR profile may constitute an uneasy and tricky task requiring high attention, since the diffraction hyperbolae may be poorly visible, interfering with one another, or confused with rounded or undulating reflectors. Nevertheless, the Thabor GPR data exhibit a great number of hyperbolae evenly distributed throughout the profiles. The quality of the Thabor GPR data allowed velocity adaptations to be determined for each profile. Shallow areas corresponding to superficial snow patches, easily identified by a strong reflector and by very high velocity adaptations (N0.20 m ns −1), were not taken into account. Two-dimensional (2-D) velocity models were produced by interpolation of the local velocity adaptations to an evenly spaced grid along the GPR profiles, using kriging techniques with a prior analysis and modeling of the semi-variogram (Davis, 2002). When the variogram exhibited a trend, as it was the case for the transversal profiles #1 and #2, the universal kriging technique was chosen, otherwise ordinary kriging was used. The quality of the kriging was assessed by leave-one-out cross validation. Finally, the gridded data were corrected for elevation using the DGPS data and the mean velocity value of the model. 3.4. Reconstruction and imaging of the true subsurface Fig. 4. Massive ice exposures on the Thabor rock glacier. a) Exposure number 18. b) Exposure number 12 (see locations in Fig. 2).
3.3. Hyperbola-based velocity adaptations and interpolation In a GPR survey, the radar wave velocity provides indirect information about the material of the medium investigated. Indeed, the radar wave velocity v is relatedpffiffiffi to ffi the relative dielectric permittivity K by the equation v = c = K , where c is the speed of light in a vacuum (3 × 10 8 m s −1). Relative dielectric permittivity of earth surface materials is highly sensitive to liquid water – and ice – content; K ranges from 1 (air) to 80 (water) with typical values of 3–4 for massive ice and 4–8 for permafrost (Brandt et al., 2007). When direct visual points are available to correlate with GPR data, as is the case with the Thabor rock glacier, radar wave velocity becomes very useful information. Radar wave velocities in a medium are generally acquired using punctual common multi-offset GPR surveys: ‘CMP’ for commonmidpoint or ‘WARR’ for wide angle reflection and refraction (Annan, 2003). However, when the goal is to obtain a two-dimensional velocity distribution, multi-offset acquisitions are required along the whole profile, not at only some selected locations (Bradford et al., 2009; Brown et al., 2009). For such multi coverage GPR surveys, and using standard single-channel apparatus, successive common-offset profiles are performed with varying offsets. Then, proper data concatenation and sorting allow for a complete velocity analysis and multifold stacking. However, in the case of a large and remote survey area such as the Thabor rock glacier, performing such measurements is unrealistic. Fortunately, owing to the heterogeneous media of rock glaciers, numerous and somehow well distributed diffraction hyperbolae are frequently recorded and can be used for local determinations of an appropriate normal move-out (NMO), or stacking, velocity, through a common operation of ‘velocity adaptation’, or ‘hyperbola fitting’ (e.g.,
Migration aims to correctly reposition dipping and arcuate reflectors into their true position, and increases the spatial resolution of the subsurface stratigraphy. Its main visual effect is to focus diffraction hyperbolae. When lateral velocity variations are not significant, a simple Stolt migration using a constant velocity (Stolt, 1978) is satisfactory. In the present study, because of the strong velocity variations in the Thabor rock glacier, we employed a migration on the basis of a 2-D velocity distribution, i.e. a standard 2-D Kirchhoff migration (see, e.g., Yilmaz, 1987). Despite strong global topographic variations, the relevance of a topographic migration (Lehmann and Green, 2000) was considered limited by the generally gentle slope of most significant dipping reflectors. The migrated profiles were filtered using a spatial moving-average filter, and topographically corrected using the field elevation data and the mean velocity value from the 2-D velocity model. At the interpretation stage, the migrated profiles were enhanced using an AGC (Automatic Gain Control) in order to highlight the most continuous and prominent reflectors in the structure. The resulting stratigraphic interpretations (presented below) were finally superimposed onto the velocity models to better understand the rock glacier structure. 4. Results 4.1. Displacement of the terminus part The position of the rock glacier front has significantly changed for the last 40 years (Fig. 2). Along the digitized front lines of the 1971 and 2006 orthophotos, the displacement range is 2.1–17.9 m with a mean value of 8.5 m and a 25–75% interquartile interval of 6.2–9.7 m. About 98% of the frontal displacement values exceed the mean error (3.6 m) for the orthorectification of the 1971 photo. The displacement may be caused by the whole rock glacier activity or, in second place, by a large settling in the frontal part.
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4.2. Velocity distribution models: materials of the Thabor rock glacier The modeled distributions of the stacking velocity (Fig. 5) show pronounced variability, which highlight the heterogeneous character of the rock glacier materials. In the longitudinal profile, for instance, the mean velocity is 0.127 m ns −1, but the velocity values range between 0.089 and 0.171 m ns −1, with a standard deviation of 0.019 m ns −1. For the other profiles, the basic descriptive statistics (Table 1) show analogous results, with a mean velocity between 0.118 and 0.13 m ns −1, a range of 0.07–0.09 m ns −1 and a standard deviation ca. 0.020 m ns −1. Most of the time, the first 5–15 m of the structure exhibit areas of high velocity (0.15–0.18 m ns −1) while in the deepest parts (very) low values (b0.10 m ns −1) may be reached. High velocities near the surface (b1–3 m depth) may arise due to the wave propagating through snow and/or air within the top blocky layer. That being said, lateral variations of the velocity are important, especially throughout the surveyed area from the lower part to the upper part, or in more restricted areas (e.g., between 750 and 1000 m in the longitudinal profile, on the margins of the transversal profiles). In order to compare the radar wave velocity with the Thabor rock glacier structure visible at the surface, the values at 2.50 m depth in the velocity models were plotted along with the locations of the massive ice exposures observed during the warm summer of 2003. The ice exposures correspond to velocities between 0.15 and 0.17 m ns −1 in the longitudinal profile (Fig. 6), and between 0.15 and 0.18 m ns −1 in the transversal profile #3. Such velocity values are consistent with those reported in the literature for buried ice (Brandt et al., 2007; Fukui et al., 2008). Stacking (RMS) velocity never reflects more than the bulk electromagnetic behavior of the material investigated and, as such, it would be premature to state that all velocities N0.15 m ns −1 in the models are evidence for buried massive ice. For example, such velocities would not be incompatible with an ice-quartzite clasts mixture owing to the low relative dielectric permittivity (K = 4.5) of quartzite. However, there is a strong basis for considering them as reflecting a high probability for massive ice occurrence. On this basis, the strongly varying velocity in the rock glacier is thought to reflect extreme contrasts in the constitutive materials. The low velocity values (b0.10 m ns −1) are, indeed, typical of wet deposits or bedrock and may be caused by the presence of water (running or in the form of pockets) in the deep rock glacier or
Table 1 Descriptive statistics from the velocity distribution models. Velocity (m/ns)
LP TP #1 TP #2 TP #3
Mean
Max
Min
Standard deviation
0.127 0.118 0.123 0.130
0.171 0.171 0.169 0.189
0.089 0.089 0.099 0.091
0.019 0.022 0.019 0.019
at/below the rock glacier floor. Such hypothesis is supported by the existence of an extensive central furrow (Fig. 2) along which it is often possible during summer to hear deep running water sounds from the surface. The remaining intermediate range of velocity values (0.10– 0.15 m ns−1) is large; it could theoretically correspond to either unfrozen dry deposits or permafrost (Daniels, 2004; Brandt et al., 2007). 4.3. Migrated and automatic gain control (AGC)-enhanced profiles: reflector patterns The migrated and AGC-enhanced GPR profiles (Figs. 7 and 8) offer a very high quality visualization of the Thabor rock glacier internal stratigraphy. The results of the 2-D Kirchoff migration are very satisfactory, since the hyperbolae initially present in the data are focused and do not perturb the interpretation. The GPR penetration depth is variable: generally 25–30 m, down to 40 m in the distal part of the rock glacier, but only a few meters in some areas of the upper rock glacier. The latter weak penetration is certainly explained by the considerable thinning of the superficial deposits and the shallow bedrock occurrence. There, the straight and 30–45° dipping features are interpreted as stratifications in the quartzite basement. Owing to the AGC enhancement, basal reflectors are well delineated, especially in the lower parts of the rock glacier, and are referred to here as rock glacier-underlying floor interfaces. Thus, the rock glacier thickness ranges from less than 10 m in the central (650–700 m position on the longitudinal profile) and upper parts, to 35–40 m in the distal part. On the whole, the stratigraphy is dense and salient, and the recognition of a few prominent and bounding reflectors highlights the specific stratigraphic organization of the rock glacier. In the longitudinal profile, apart from the attenuation areas, a shallower (15–20 m) part
Fig. 5. Modeled distribution of the stacking velocity obtained from the GPR common-offset data of the Thabor rock glacier. The main illustration is the model of the longitudinal profile. All transversal models (TP #1, TP #2, TP #3) are E (left)–W (right) oriented. See Fig. 2 for the exact locations of the profiles. The full black lines on the models delineate areas with velocity values N 0.15 m ns−1. The dashed black lines delineate areas with velocity values b 0.10 m ns−1. The models derived from the application of kriging techniques to local velocity determinations. The amount of local velocity determinations in the data is the following: 90 in the longitudinal profile (0.062 per horizontal meter), 37 in the transversal profile #1 (0.112 m−1), 50 in the transversal profile #2 (0.128 m−1), and 79 in the transversal profile #3 (0.119 m−1).
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section of the transversal profile #3, which corresponds to an adjacent deposit, shows the same type of stratigraphy, with undulating reflectors at depth and dipping, embedded, and curved reflectors close to the surface (Fig. 8). The significance of the stratigraphic patterns is crucial for our understanding of the rock glacier structure. For example, 1) Straight and parallel reflectors concordant with a basal one express low longitudinal deformation and/or vertical compaction; 2) Chaotic reflectors may reflect local downwasting and perturbation of the shallow structure; 3) Undulating and toplapping reflectors; and 4) upward- or rock glacier center-dipping reflectors rather evocate the occurrence of compressive stresses and thrusting. Indeed, undulating and toplapping reflectors are well known in rock glacier deformation structures (e.g., Degenhardt et al., 2003; Monnier et al., 2008), and the dipping reflectors observed here are analogous to GPR signatures given by thrust planes in glaciotectonized sediments (Overgaard and Jakobsen, 2001; Sadura et al., 2005) or by debris inclusions in glacier ice (Woodward et al., 2003) and debris-covered glacier ice structures in rock glaciers (Fukui et al., 2008). Fig. 6. Plotting of the radar wave velocity values in the longitudinal profile at a depth of 2.50 m with the locations of the 2003 massive ice exposures.
predominantly made of upward-dipping reflectors lies on or imbricates into a deeper part made of straight and parallel reflectors, concordant with the basal one. In the terminus part, undulating reflectors are preferentially encountered. The three transversal profiles have similar organization. Undulating and toplapping reflectors are present in the deeper part. In the central part, where the morphology is poorly differentiated or may consist of straight ramps, the reflectors are mainly chaotic or rock glacier center-dipping. On the lateral margins, where the morphology consists of arcuate ridges or straight ramps, the reflectors are dipping toward the center of the rock glacier. Moreover, the eastern
5. Discussion: structure and genesis of the Thabor rock glacier 5.1. Structural and stratigraphical units of the rock glacier By integrating the main data available – stratigraphy, radar wave velocity, and surface morphology – and according to a good correspondence between the velocity distributions and the main reflectors, four types of structural unit are identified along the entire GPR profiles (Figs. 7 and 8). Units of type 1 correspond to the bedrock floor. The bedrock floor connects with the surface in restricted sections of the upper rock glacier; it is very close to the surface halfway down the longitudinal profile, so that the rock glacier appears almost cut in two, with the upper part being quasi disconnected from the lower part. In the upper parts of the rock glacier the bedrock floor
Fig. 7. Stratigraphic patterns and structural units of the longitudinal profile (migrated, enhanced with AGC). The surface morphology is indicated in the upper part of the illustration. The units 1 to 4 are described as following: 1) Units generally disconnected from the surface and separated from other units by a basal reflector; they have poorly visible patterns of reflectors and generally (but not always) correlates low radar wave velocities (b0.10–0.12 m ns−1). 2) Units of the deep rock glacier, connecting to the surface by the lateral margins (or without connection with the surface), characterized by parallel/concordant, or undulating–toplapping reflectors, and a midrange radar wave velocity (0.10–0.13 m ns−1). 3) Units characterized by arcuate ridges or straight ramps on the surface, a pattern of center−/−up dipping (frequently curved) reflectors, and a quite high radar wave velocity (N 0.13–0.14 m ns−1). 4) Units characterized by a poorly differentiate surface morphology (but with straight ramps and crevasse-like features in some parts) a pattern of center-/updipping reflectors (indeed chaotic, in some cases), and a high radar wave velocity (N 0.14–0.15 m ns−1).
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Fig. 8. Stratigraphic patterns and structural units of the transversal profiles (migrated, enhanced with AGC). The surface morphology is indicated in the upper part of the illustrations. The description of the types of unit is given in the legend of Fig. 7.
is probably frozen (high radar wave velocities) while it is likely unfrozen and even penetrated by melting water in the lower parts (low radar wave velocities). The bedrock floor exhibits few structural features (dipping joints) in the longitudinal GPR profile. Units of type 2 constitute the “basal sole” of the rock glacier. This basal sole connects with the surface at the terminus and at the lateral margins. It is made of debris potentially frozen (typical radar wave velocities between 0.10 and 0.13 m ns −1). Its stratigraphic features (undulating and toplapping) are the expression of compaction and gentle compression. Units of type 3 constitute the “deformed margins” of the rock glacier. At the surface, the morphology is highly differentiated with arcuate ridges and, over the transition with the central part of the rock glacier, straight ramps. The deformed margins are made of
ice-debris mixtures with a large ice content and with possible inclusions of massive ice (radar wave velocity N0.15 m ns −1 in several parts), which correlate with observations at the surface (Fig. 3). The stratigraphic features (strongly curved and dipping reflectors) are the expression of strong compression. Finally, units of type 4 constitute the “shallow debris-covered ice core”. The ice is thought to be massive (radar wave velocity typically between 0.15 and 0.17 m ns −1) but with a high debris content. Indeed, the gently curved and dipping reflectors are best interpreted as debris inclusions along thrust planes in the ice. The architectural model of the rock glacier appears, thus, as one of a debris-covered massive ice core superimposed on, or embedded into, a sediment complex. The debris-covered massive ice core is the remnant of the former Thabor glacier, which stationary
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and re-advancing positions since the end of the LIA can finally been inferred (Fig. 9). 5.2. Origin of the surface features Given our structural interpretation, the surface morphology (Fig. 2) may be interpreted as being mainly the result of glacier– permafrost interactions. As primarily inferred, the hummocks and very upper lateral ridges are of glacial origin. The occurrence of crevasse-like features at the surface corresponds with the inferred extent of debris-covered ice, hence these features probably relate to fractures within the massive ice core. In the lower-central part of the rock glacier, a large convex slope, preceding a crevasse-like feature, evocate an accumulation form and may correspond to a former terminus of the stagnating glacier (Fig. 9). Straight ramps occur both in the deformed margins and in the debris-covered ice core; they are thought to result from downwasting by melting of internal ice. The origin of the arcuate ridges on the Thabor rock glacier was extensively discussed by Monnier (2007). Further understanding of these features can be gained by focusing on the terminus part of the rock glacier in the longitudinal profile (Fig. 10). The location of the ridges correlates exactly with the dipping planes. Some of the ridges, especially number 4 (Fig. 10c,d), exhibit a peculiar morphology with a steep and sharp back slope preceding a flat and depressed area. This morphology was
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first interpreted as the result of differential ablation processes (Monnier, 2007). However, the observed connection between the back slope of the ridge and the dipping reflector leads to a new interpretation: the back slope is a thrusting plane-related scarp, and the entire ridge is the expression of a compressive deformation process. Thus, considering the structure of the rock glacier as a whole and the peripheral, concentric distribution of the ridges, the so-called deformed margins correspond to a belt of frozen/ice-cored sediments (potentially classifiable as moraines) deformed by ice push. The deformations (internal thrusting and surface ridges) may have been subsequently increased by the displacement of the rock glacier tongue in the terminus part. 5.3. History of the Thabor rock glacier At the end, the following genesis model can be established. During the Little Ice Age, as attested by the Carte d'État Major (1847–1864), the Thabor glacier advanced over and deformed preexisting deposits in discontinuous permafrost conditions. Lateral and terminal arcuate ridges were created – or accentuated if they preexisted – by ice pushrelated compressive stresses, perhaps through several cycles of short advances and retreats (Fig. 9). The same conclusion was proposed by Lugon et al. (2004) for rock glaciers in the Posets Massif, Spanish Pyrenees. The preexisting deposits mentioned are likely to be a
Fig. 9. Spatial extent (on 3-D Google Earth sketch) of the debris-covered massive ice in the Thabor rock glacier as determined by matching the structural interpretation of the GPR profiles and the surface morphology. The upper area with thin debris cover over ice or bedrock is also depicted. Consequently, the former terminus positions of the Thabor glacier since the end of the LIA can be inferred using the historical maps and the morphology.
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Fig. 10. Focus on the terminus part of the longitudinal profile showing the relationship between the arcuate ridges and the curved dipping reflectors. a) Numbering and position of the arcuate ridges relative to the longitudinal GPR profile. b) Zoom on the terminus part of the longitudinal profile; the numbers are the same as those in a). c) Detail of an arcuate ridge with a steep and sharp back slope preceding a flat and depressed area. The backpack gives the scale. d) Interpreted version of the box in b).
previous glacial forefield. Indeed, it is well known in the Alps (Patzelt, 1974; Le Roy Ladurie, 2003) and especially in their occidental part (Deline and Orombelli, 2005), that several glacial phases analogous in spatial extent to the LIA occurred during the ten thousands of years preceding the LIA. Moreover, in the valley down the Thabor rock glacier, the only visible glacial deposits are lateral moraines at an elevation of 2450 m that were attributed to the Late-glacial period (Monnier, 2006). The “basal sole” of the rock glacier as well as some parts of the “deformed margins” are, therefore, thought to have been deposited (even shaped?) by one or several glacial advances of the Thabor glacier taking place between the end of the Late-glacial and the beginning of the LIA. Following the end of the LIA and under the progressively warmer climatic conditions, the Thabor glacier began retreating and burying, resulting in the superimposition of the debris-covered ice core onto the preexisting accumulation. The origin of the burial of the ice remains uncertain. It may be related to repeated rock falls in the accumulation area of the former Thabor glacier, which supplied a considerable stock of debris to the glacier–rock glacier system. Indeed, large and clear failure scars are visible on the rock walls of the Cheval Blanc and the Mt. Thabor, and several sudden rock falls were observed during field trips at the site. Additionally, melting out of internal debris in the ice was possible. Considering the occurrence of both origins is appealing: climate warming melts the glacier and the adjacent rock walls and, at the end, produces a rock glacier. Irrespective of the burial processes, the retreat of the Thabor glacier led to the apparition of the central depression in the upper part of the rock glacier. Furthermore, when the burial/retreat of the glacier reached an advanced level (in the 20th century), new dynamics
occurred: the lower part of the rock glacier began to move by permafrost creep, and the upper part began to settle by melting of the buried ice, the volume of which was especially important in the central part (and at the origin of the network of straight ramps). Completely disappeared from the surface at the turn of the 20th and 21st centuries, the Thabor glacier left in the landscape a disconnection between the debris source on Mt Thabor and the rock glacier: the depression at the upper end of the rock glacier reached its maximum dimensions and the bedrock started to appear through thin, discontinuous remnants of the debris-covered glacier. 6. Conclusions We used a complete set of historical, geomorphological, and geophysical data to investigate the structure and genesis of the Thabor rock glacier. The extensive common-offset GPR data collected at the site were used to model the distribution of the stacking radar wave velocity and reveal the stratigraphy of the rock glacier. The GPR results were correlated with the morphological analysis of the surface. Our study strengthens previous morphological analyses (Monnier, 2007) and the following important points deserve to be highlighted. 1) The 2-D models of the radar wave velocity show a very heterogeneous structure, with velocities ranging between 0.08 and 0.18 m ns −1. Following comparison with direct observations of the internal structure, velocities higher than 0.15 m ns −1 can be considered as reflecting a high probability for massive ice occurrence. 2) The stratigraphy of the rock glacier is varied, with most features expressive of compressive stresses and thrusting, and prominent bounding reflectors dividing the structure. The integration of velocity
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modeling, stratigraphic representation and surface morphology information depicts a rock glacier composed of a shallow (b15 m) debris-covered massive ice core overlying a (potentially frozen) sole of sediments, and embedded into deformed permafrost margins. 3) The genesis of the rock glacier has probably involved several temporal phases since the Late-glacial; however, the current rock glacier morphology is mainly the product of glacier–permafrost interactions during the LIA and subsequent deformation processes. During the LIA the Thabor glacier advanced over and deformed preexisting deposits before the resulting system became modified in the last 150 years by ice burial, permafrost creep, and downwasting. At the end, is the Thabor rock glacier still a true one? Haeberli (1979, p. 46) states that “All transitions between push-moraines and rock glaciers may exist in cases where a glacier was in contact with a rock glacier during neo-glaciation or Holocene advances”. This is exactly where the study of the Thabor rock glacier leads us. The actual and impressively esthetic tongue is neither a debris-covered glacier, nor a morainic complex, nor a rock glacier in the sense of an individual creeping permafrost body (Barsch, 1992). The Thabor rock glacier is all of that in one feature, and thus expresses the complexity of the interactions that can occur on the surface as consequences of the environmental changes of the last centuries and millennia. Acknowledgments The authors would like to thank friendly and sincerely Marion Siboni and Guillaume Tanguy who provided valuable and courageous help during the field survey. Moreover, the Thabor rock glacier is part of the Thabor natural site listed by the Ecology and Sustainable Development Minister of France; the temporal implantation of scientific material on this site was permitted by a decree of the Savoy prefecture emitted the 30th of June of 2008. The authors renew their acknowledgements to the Savoy prefecture for this permission. Finally, Shelley MacDonell provided helpful review of the language in the manuscript. References Annan, A.P., 2003. Ground Penetrating Radar principles, Procedures and Applications. Sensors & Software Inc., Mississauga. Barsch, D., 1992. Permafrost creep and rockglaciers. Permafrost and Periglacial Processes 3, 175–188. Bradford, J.H., Nichols, J., Mikesell, T.D., Harper, J.T., 2009. Continuous profiles of electromagnetic wave velocity and water content in glaciers: an example from Bench Glacier, Alaska, USA. Annals of Glaciology 50, 1–9. Brandt, O., Langley, K., Kohler, J., Hamran, S.-E., 2007. Detection of buried ice and sediment layers in permafrost using multi-frequency Ground Penetrating Radar: a case examination on Svalbard. Remote Sensing of Environment 111, 212–227. Brown, J., Nichols, J., Steinbronn, L., Bradford, J., 2009. Improved GPR interpretation through resolution of lateral velocity heterogeneity: example from an archaeological site investigation. Journal of Applied Geophysics 68, 3–8. Ground Penetrating Radar, In: Daniels, D.J. (Ed.), 2nd Edition. The Institution of Electrical Engineers, London. 726 pp.
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Davis, J.C., 2002. Statistics and Data Analysis in Geology, Third edition. John Wiley & Sons, New York. 638 pp. Degenhardt, J.J., 2009. Development of tongue-shaped and multilobate rock glaciers in alpine environments — interpretations from ground penetrating radar surveys. Geomorphology 109, 94–107. Degenhardt, J.J., Giardino, J.R., Junck, M.B., 2003. GPR survey of a lobate rock glacier in Yankee Boy Basin, Colorado, USA. In: Bristow, C.S., Jol, H.M. (Eds.), Ground Penetrating Radar in Sediments: Geological Society of London, Special Publications, 21, pp. 167–179. Deline, P., Orombelli, G., 2005. Glacier fluctuations in the Western Alps during the Neoglacial, as indicated by the Miage morainic amphitheatre (Mont Blanc massif, Italy). Boreas 34, 456–467. Dix, C.H., 1955. Seismic velocities from surface measurements. Geophysics 20, 180–195. Ferrand, H., 1892. Autour de Modane: la chaîne frontalière franco-italienne entre le Mont Thabor et le Mont Cenis. Annuaire du Club Alpin Français, pp. 3–49. Fukui, K., Sone, T., Strelin, J.A., Torielli, C.A., Mori, J., Fujii, Y., 2008. Dynamics and GPR stratigraphy of a polar rock glacier on James Ross Island, Antarctic Peninsula. Journal of Glaciology 54, 445–451. Gidon, M., 1977. Carte géologique simplifiée des Alpes occidentales, du Léman à Dignes (1/2500000). Didier Richard Editions and BRGM, Grenoble and Orléans, 2 maps, instructions 18 p. Haeberli, W., 1979. Holocene push-moraine in alpine permafrost. Geografiska Annaler 61A, 43–48. Hausmann, H., Krainer, K., Brückl, E., Mostler, W., 2007. Internal structure and ice content of Reichenkar rock glacier (Stubai Alps, Austria) assessed by geophysical investigations. Permafrost and Periglacial Processes 18, 351–367. Isaksen, K., Ødegård, S.R., Eiken, T., Sollid, J.L., 2000. Composition, flow and development of two tongue-shaped rock glaciers in the permafrost of Svalbard. Permafrost and Periglacial Processes 11, 241–257. Kääb, A., Weber, M., 2004. Development of transverse ridges on rock glaciers: field measurements and laboratory experiments. Permafrost and Periglacial Processes 15, 379–391. Kaiser, B., 1987. Les versants de Vanoise : enjeux traditionnels et fonctionnement morphoclimatique. Ph.D., University Paris 7, 1088 p. Le Roy Ladurie, E., 2003. Histoire du Climat depuis l'an mil. Flammarion, Paris. 541 pp. Lehmann, F., Green, A., 2000. Topographic migration of georadar data: implications for acquisition and processing. Geophysics 65, 836. Lugon, R., Delaloye, R., Serrano, E., Reynard, E., Lambiel, C., Gonzales-Trueba, J.J., 2004. Permafrost and Little Ice Age glacier relationships, Posets Massif, Central Pyrenees, Spain. Permafrost and Periglacial Processes 15, 207–220. Monnier, S., 2006. Les glaciers rocheux, objets géographiques. Analyse spatiale multiscalaire et investigations environnementales. Application aux Alpes de Vanoise. Ph.D., University Paris 12, 339 p. Monnier, S., 2007. Du glacier au glacier rocheux, depuis la fin du Petit âge glaciaire, au pied du Mont Thabor (Alpes du Nord, France). Quaternaire 18, 283–294. Monnier, S., Camerlynck, C., Rejiba, F., 2008. Ground-penetrating radar survey and stratigraphic interpretation of the Plan du Lac rock glaciers, Northern French Alps. Permafrost and Periglacial Processes 19, 19–30. Overgaard, T., Jakobsen, P.R., 2001. Mapping of glaciotectonic deformation in an ice marginal environment with ground penetrating radar. Journal of Applied Geophysics 47, 191–197. Patzelt, G., 1974. Holocene variations of glaciers in the Alps. Colloques Internationaux du CNRS 219, 51–59. Sadura, S., Martini, I.P., Endres, A.L., Wolf, K., 2005. Morphology and GPR stratigraphy of a frontal part of an end moraine of the Laurentide Ice Sheet: Paris Moraine near Guelph, ON, Canada. Geomorphology 75, 212–225. Stolt, R.H., 1978. Migration by Fourier transform. Geophysics 43, 23–48. Voiron, H., 1983. Les régimes nivométriques de la Vanoise. Travaux scientifiques du Parc national de la Vanoise XIII, pp. 81–89. Woodward, J., Murray, T., Clark, R.A., Stuart, G.W., 2003. Glacier surge mechanisms inferred from ground-penetrating radar: Kongsvegen, Svalbard. Journal of Glaciology 49, 473–480. Yilmaz, O., 1987. Seismic Data Processing (Investigation in Geophysics 10). Society of Exploration Geophysicists, Tulsa. 2027 pp.
Contents lists available at ScienceDirect
Geomorphology j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / g e o m o r p h
Structure and genesis of the Thabor rock glacier (Northern French Alps) determined from morphological and ground-penetrating radar surveys Sébastien Monnier a,⁎, Christian Camerlynck b, Fayçal Rejiba b, Christophe Kinnard a, Thierry Feuillet c, Amine Dhemaied b a b c
Centro de Estudios Avanzados en Zonas Áridas (CEAZA), Campus Andrés Bello, Casilla 599, Raúl Bitrán s/n, La Serena, Chile UMR 7619 Sisyphe, Université Paris 6 Pierre et Marie Curie, 4 place Jussieu, 75252 Paris Cedex 05, France UMR 6554 LETG, Université Paris Sorbonne-Paris IV, 191 rue Saint Jacques, 75005 Paris, France
a r t i c l e
i n f o
Article history: Received 15 December 2010 Received in revised form 29 April 2011 Accepted 6 July 2011 Available online 31 July 2011 Keywords: Rock glacier Ground-penetrating radar Internal structure Glacier–permafrost interactions
a b s t r a c t Landform analysis and ground-penetrating radar (GPR) were used to investigate the Thabor rock glacier, in the Northern French Alps. The surface features of the rock glacier were classified and described, with emphasis on massive ice exposures. The retreat of the former Thabor glacier since the Little Ice Age (LIA) was documented through an analysis of historical sources, and recent movements of the rock glacier were inferred from orthophoto-based measurements. Two-dimensional (2-D) models of the radar wave velocity were derived from the raw GPR data, using the numerous diffraction hyperbolae for local determinations of the velocity and kriging interpolation techniques. Subsequently, the profiles were migrated through a 2-D Kirchhoff migration method using the interpolated velocities. The 2-D velocity models exhibit pronounced spatial variations and, in several locations, high values (N 0.15 m ns−1) potentially corresponding to massive ice. On the other hand, while the migrated profiles show numerous layers, the internal stratigraphy of the rock glacier is dominated by a few prominent internal boundaries. The integration of morphology, radar wave velocity, and internal stratigraphy allowed us to identify the main structural units of the rock glacier as well as to explain its genesis: the rock glacier was formed by the imbrication of a massive ice core, originating from the retreat of the former Thabor glacier since the LIA, into pre-existing glacial deposits. © 2011 Elsevier B.V. All rights reserved.
1. Introduction During the last decade, important insights into the geomorphology of high mountain regions have been gained from the use of groundpenetrating radar (GPR). Previous GPR studies have looked at various morphological features such as talus slopes, glacial deposits and, most commonly, rock glaciers. In particular, the advent of GPR technology in rock glacier research has been decisive. Due to both its complexity and inaccessible nature, the internal structure of rock glaciers is often unknown. Hence GPR is often considered an ideal tool for “taking a look” into the subsurface of rock glaciers. Previous GPR studies of rock glaciers have permitted the examination of the layered structure, the main internal boundaries, and the thickness of the rock glaciers studied (e.g., Isaksen et al., 2000; Hausmann et al., 2007; Monnier et al., 2008; Degenhardt, 2009). However, only rarely GPR data were correlated with direct observations of the internal structure (Fukui et al., 2008). Therefore, previous GPR studies were generally unsuccessful in defining the nature of the inner materials, especially in revealing the presence of ice. In this paper, we present work
⁎ Corresponding author. Tel.: + 56 51 334 869. E-mail address: [email protected] (S. Monnier). 0169-555X/$ – see front matter © 2011 Elsevier B.V. All rights reserved. doi:10.1016/j.geomorph.2011.07.004
conducted on the Thabor rock glacier, Northern French Alps. Taking into account the morphology, the recent glacial history of the site, and the occurrence of several direct observations of the shallow rock glacier structure, we process large and complex common-offset GPR data in order to both estimate the properties of the internal materials and highlight the rock glacier stratigraphy. We then use the combination of these data and observations to understand how the Thabor rock glacier has formed. 2. The Thabor rock glacier The Thabor rock glacier is located in the Mt Thabor Massif, Northern French Alps, south of the Vanoise Massif (Fig. 1). The Thabor Massif is an area with relatively moderate crest elevations (ca. 3000 m), and is known for being a regional dry pole — ~700 mm of annual precipitations in valley stations (Voiron, 1983; Kaiser, 1987). On the basis of meteorological data recorded by Meteo France in the Modane valley-station, and using a lapse rate of −0.6 °C 100 m −1, the 0 °C and −2 °C isotherms may be located at approximately 2460 and 2790 m, respectively. The Thabor rock glacier extends between 2900 and 2650 m, inside a north-facing amphitheater. Surrounding summits (Mt Thabor, 3178 m; Thabor Peak, 3207 m; and Cheval Blanc, 3020 m) are made of Triassic quartzites while the lower slopes
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Fig. 1. Location (a) and geological context (b) of the Thabor rock glacier. The coordinates are expressed in the UTM−WGS84 system (as for all subsequent maps). The geological context is depicted onto a 3-D Google Earth sketch. Q: quartzites. h: carboniferous schists. T.s.: Thabor synclinal. Q-h: geological contact between the quartzites and the carboniferous schists. The Google Earth sketch is 1.4 km wide.
are made of carboniferous schists, as a result of the differential erosion of the local synclinal structure belonging to the Zone houillère Briançonnaise (Gidon, 1977). 2.1. Morphology The Thabor rock glacier is one of the largest rock glaciers of the Northern French Alps, being more than 1 km long and up to 600 m wide, with a surface area close to 60 ha. These large dimensions and the light color arising from the quartzitic composition of the surface compensate for the relatively low height of the front (rarely more than 15 m) and give a striking appearance to the rock glacier, which is clearly visible from aerial view or from distant crests. The rock glacier surface, composed of 20–40 cm long clasts, exhibits several morphological features (Fig. 2). The central part of the rock glacier above 2750 m is depressed and has a poorly differentiated morphology. In the upper part, typical glacial features such as lateral morainic crests and hummocks are present; crevasse-like features reminiscent of true crevasses in glacier ice are also noticeable. Over the whole rock glacier, the two main morphological features encountered are arcuate ridges and long and straight ramps, which extend along both sides of the rock glacier. The arcuate and embedded ridges are 1–3 m high and
10–100 m long, and are mainly present in the lower part and on the western flank of the rock glacier. These often exhibit a dissymmetric cross profile, with a smooth, convex front slope and a steep back slope; the latter precedes a flat and depressed area. The long and straight ramps concentrate on the eastern flank of the rock glacier. Running over several hundred meters, they exhibit a steep, 1–2 m high slope facing the center-line depression and a gentle slope facing the rock glacier margin. The arcuate ridges, according to experiments made on rock glacier rheology (Kääb and Weber, 2004), suggest compression processes, with possible local downwasting on their back side, while the long and straight ramps suggest downwasting processes. Moreover, on its western and especially eastern flanks, the rock glacier connects with other deposits that exhibit morphological signs of permafrost debris deformation (ridges and frontal slopes). 2.2. Diachronic information The Thabor rock glacier is a recent glacial forefield. Indeed, several available documents (Fig. 3) show an important glacial retreat in the amphitheater since the middle of the 19th century. These include the socalled Carte d'État-Major (1843–1867), topographic maps from the beginning of the 20th century (based on late 19th century topographic
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Fig. 2. The Thabor rock glacier morphology and GPR survey. Background: IGN orthophoto of 2007 (as for all subsequent maps). Modified from Monnier, 2007.
surveys), a picture of the massif in a publication of Ferrand (1892), and topographic maps of the late 20th century. Nevertheless, integration of this information requires caution because of potential mapping errors in depicting both the topography (for maps of the 19th century) and the glacier extent (for all dates). It is thus necessary to adjust the information by seeking a coherency between the various documents and by taking into account the geomorphology of the rock glacier. According to the Carte d'Etat-Major (1843–1867), in the middle of the 19th century, which corresponds to the end of the Little Ice Age (Le Roy Ladurie, 2003), the Thabor glacier completely filled the amphitheater and extended beyond the actual limits of the rock glacier. However, according to the topographic data of 1890 and the picture in Ferrand (1892), it had severely reduced in size 40 years later, probably ending halfway up the actual rock glacier tongue. Little Ice Age (LIA hereafter) glacial features are usually well preserved in the area, especially in quartzite accumulations, but there are no glacial features in the amphitheater around the rock glacier terminus, where the slopes are quite gentle. Therefore, the glacial extent interpretation provided by the Carte d'Etat-Major (mid-19th century) is clearly exaggerated and must have resulted from approximations made in locating landforms and from the incorporation of snow patches and/or adjacent small glaciers. The most reasonable interpretation sets the mid-19th century extent of the Thabor glacier near the current rock glacier terminus and all subsequent extents within the current rock glacier margins. Finally, although the 1977 map shows the Thabor glacier as a small residual glacier occupying the upper reach of the rock glacier, aerial photos of the same decade suggest a subtly larger extent. Today, the Thabor glacier has completely vanished.
2.3. Exposures of massive ice Despite the disappearance of the Thabor glacier, buried glacier ice is still present at the site. The upper part of the rock glacier above 2800–2850 m is strongly depressed (~30 m) at its center, and exhibits remnants of a rapidly decaying debris-covered glacier, as demonstrated by the appearance of bedrock outcrops over the last 30 years (Fig. 2). Moreover, while walking down the rock glacier at the end of the warm summer of 2003, we recorded 24 exposures of massive ice in furrows, depressions or spaces between ridges usually occupied by snow (Figs. 2 and 4). In some exposures, the ice was banded. In the upper part (ca. 2800 m), but away from the central furrow, ice was found a few decimeters beneath the surface by means of manual digging. The lowest exposure of massive ice was found within an elliptic depression on the distal part of the rock glacier at 2730 m (pt. 24, Fig. 2). These ice exposures correlate with water temperatures near 0 °C measured in the spring in front of the rock glacier during the same late summer, in complete absence of snow on the upper part. 2.4. Working hypotheses about the Thabor rock glacier formation The evolution of the Thabor rock glacier over the last 150 years and the existence of a massive ice-core are not yet resolved. Monnier (2007) suggested a model of formation in which the Thabor rock glacier represents a transition between a rock glacier and a push moraine complex, following ideas of Haeberli (1979) about similar landforms in the region of the Grubengletscher, Switzerland. Thus, the morphology of the Thabor rock glacier would originate from glacier–
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Fig. 3. Historical documents providing information about the Thabor glacier retreat. a) “Carte d'État Major” (1847–1867). b) Picture of the massif, after a photograph, in Ferrand (1892). c) Topographic map from the beginning of the 20th century (topographic survey: 1890). d) IGN map of the late 20th century (topographic and photogrammetric surveys: 1977).
permafrost interactions: the LIA Thabor glacier advances would have created marginal arcuate ridges by compression/accumulation of frozen sediments; then, the glacier would have progressively decayed both by retreat and burial, leading to differentiated dynamics in the resulting debris-ice tongue — creep in zones with low ice wasting (distal part; amplification of the arcuate ridge morphology) and collapse in zones with pronounced ice wasting (central part; formation of the straight ramps). Such a scenario leaves unknown the origin of the forefield present before the LIA glacier advance.
3. Methods 3.1. Photogrammetric measurements Frontal movements of the Thabor rock glacier were calculated using an aerial photo of 1971 and an aerial orthophoto of 2006. Both photos were provided by the IGN (Institut Géographique National, France). The photo of 1971 was scanned and orthorectified. On the basis of 35 points selected throughout the two orthophotos and corresponding to stable topography (large blocks in flat areas, and outcrops of bedrock), the mean error for the orthorectification of the 1971 photograph is 3.6 m. Both orthophotos have a high spatial
resolution (1 m for 1971, and 0.5 m for 2006). Therefore, the base of the rock glacier front is easily recognizable and was digitally delineated. The displacements were calculated at 106 points sampled along the front.
3.2. GPR (ground-penetrating radar) data acquisition Common-offset GPR measurements were performed on the Thabor rock glacier using a pulseEKKO 100 system (Sensors & Software Inc.) with unshielded 50 MHz frequency antennae. The antennae were separated by 2 m, oriented perpendicular to the survey direction, and traces were recorded every 50 cm. Each recorded trace was the result of 16 stacks in order to increase the signal over noise ratio. GPR data were acquired along four profiles (Fig. 2). The longitudinal profile is 1450 m long and extends between 2675 and 2928 m above sea level (a.s.l.). Three transversal profiles, labeled from north to south #1, #2 and #3, have the following characteristics, respectively: 330 m long between 2735 and 2764 m a.s.l., 390 m long between 2769 and 2799 m a.s.l., and 660 m long between 2821 and 2853 m a.s.l. Elevation data were obtained with a differential GPS system with centimeter resolution. The preliminary processing of the GPR radargrams included zero-time correction, the removal of the DC shift component and that of the very low frequency component using a highpass filter (i.e. dewow filtering).
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Brandt et al., 2007). In the case of a tabular subsurface and for small offsets compared with depth, the stacking velocity is a good approximation of the root mean square (RMS) velocity between the antennae and the hyperbola apex. Whereas ‘true’ interval velocities could be retrieved from stacking velocities (Dix, 1955), in this study the latter were used for migration, and thus the conversion was not performed. Furthermore, referring to the aim of comparing the radar wave velocity with the superficial ice exposures, the shallow RMS velocity is close to the actual interval velocity. That being said, repeated hyperbola fittings throughout a GPR profile may constitute an uneasy and tricky task requiring high attention, since the diffraction hyperbolae may be poorly visible, interfering with one another, or confused with rounded or undulating reflectors. Nevertheless, the Thabor GPR data exhibit a great number of hyperbolae evenly distributed throughout the profiles. The quality of the Thabor GPR data allowed velocity adaptations to be determined for each profile. Shallow areas corresponding to superficial snow patches, easily identified by a strong reflector and by very high velocity adaptations (N0.20 m ns −1), were not taken into account. Two-dimensional (2-D) velocity models were produced by interpolation of the local velocity adaptations to an evenly spaced grid along the GPR profiles, using kriging techniques with a prior analysis and modeling of the semi-variogram (Davis, 2002). When the variogram exhibited a trend, as it was the case for the transversal profiles #1 and #2, the universal kriging technique was chosen, otherwise ordinary kriging was used. The quality of the kriging was assessed by leave-one-out cross validation. Finally, the gridded data were corrected for elevation using the DGPS data and the mean velocity value of the model. 3.4. Reconstruction and imaging of the true subsurface Fig. 4. Massive ice exposures on the Thabor rock glacier. a) Exposure number 18. b) Exposure number 12 (see locations in Fig. 2).
3.3. Hyperbola-based velocity adaptations and interpolation In a GPR survey, the radar wave velocity provides indirect information about the material of the medium investigated. Indeed, the radar wave velocity v is relatedpffiffiffi to ffi the relative dielectric permittivity K by the equation v = c = K , where c is the speed of light in a vacuum (3 × 10 8 m s −1). Relative dielectric permittivity of earth surface materials is highly sensitive to liquid water – and ice – content; K ranges from 1 (air) to 80 (water) with typical values of 3–4 for massive ice and 4–8 for permafrost (Brandt et al., 2007). When direct visual points are available to correlate with GPR data, as is the case with the Thabor rock glacier, radar wave velocity becomes very useful information. Radar wave velocities in a medium are generally acquired using punctual common multi-offset GPR surveys: ‘CMP’ for commonmidpoint or ‘WARR’ for wide angle reflection and refraction (Annan, 2003). However, when the goal is to obtain a two-dimensional velocity distribution, multi-offset acquisitions are required along the whole profile, not at only some selected locations (Bradford et al., 2009; Brown et al., 2009). For such multi coverage GPR surveys, and using standard single-channel apparatus, successive common-offset profiles are performed with varying offsets. Then, proper data concatenation and sorting allow for a complete velocity analysis and multifold stacking. However, in the case of a large and remote survey area such as the Thabor rock glacier, performing such measurements is unrealistic. Fortunately, owing to the heterogeneous media of rock glaciers, numerous and somehow well distributed diffraction hyperbolae are frequently recorded and can be used for local determinations of an appropriate normal move-out (NMO), or stacking, velocity, through a common operation of ‘velocity adaptation’, or ‘hyperbola fitting’ (e.g.,
Migration aims to correctly reposition dipping and arcuate reflectors into their true position, and increases the spatial resolution of the subsurface stratigraphy. Its main visual effect is to focus diffraction hyperbolae. When lateral velocity variations are not significant, a simple Stolt migration using a constant velocity (Stolt, 1978) is satisfactory. In the present study, because of the strong velocity variations in the Thabor rock glacier, we employed a migration on the basis of a 2-D velocity distribution, i.e. a standard 2-D Kirchhoff migration (see, e.g., Yilmaz, 1987). Despite strong global topographic variations, the relevance of a topographic migration (Lehmann and Green, 2000) was considered limited by the generally gentle slope of most significant dipping reflectors. The migrated profiles were filtered using a spatial moving-average filter, and topographically corrected using the field elevation data and the mean velocity value from the 2-D velocity model. At the interpretation stage, the migrated profiles were enhanced using an AGC (Automatic Gain Control) in order to highlight the most continuous and prominent reflectors in the structure. The resulting stratigraphic interpretations (presented below) were finally superimposed onto the velocity models to better understand the rock glacier structure. 4. Results 4.1. Displacement of the terminus part The position of the rock glacier front has significantly changed for the last 40 years (Fig. 2). Along the digitized front lines of the 1971 and 2006 orthophotos, the displacement range is 2.1–17.9 m with a mean value of 8.5 m and a 25–75% interquartile interval of 6.2–9.7 m. About 98% of the frontal displacement values exceed the mean error (3.6 m) for the orthorectification of the 1971 photo. The displacement may be caused by the whole rock glacier activity or, in second place, by a large settling in the frontal part.
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4.2. Velocity distribution models: materials of the Thabor rock glacier The modeled distributions of the stacking velocity (Fig. 5) show pronounced variability, which highlight the heterogeneous character of the rock glacier materials. In the longitudinal profile, for instance, the mean velocity is 0.127 m ns −1, but the velocity values range between 0.089 and 0.171 m ns −1, with a standard deviation of 0.019 m ns −1. For the other profiles, the basic descriptive statistics (Table 1) show analogous results, with a mean velocity between 0.118 and 0.13 m ns −1, a range of 0.07–0.09 m ns −1 and a standard deviation ca. 0.020 m ns −1. Most of the time, the first 5–15 m of the structure exhibit areas of high velocity (0.15–0.18 m ns −1) while in the deepest parts (very) low values (b0.10 m ns −1) may be reached. High velocities near the surface (b1–3 m depth) may arise due to the wave propagating through snow and/or air within the top blocky layer. That being said, lateral variations of the velocity are important, especially throughout the surveyed area from the lower part to the upper part, or in more restricted areas (e.g., between 750 and 1000 m in the longitudinal profile, on the margins of the transversal profiles). In order to compare the radar wave velocity with the Thabor rock glacier structure visible at the surface, the values at 2.50 m depth in the velocity models were plotted along with the locations of the massive ice exposures observed during the warm summer of 2003. The ice exposures correspond to velocities between 0.15 and 0.17 m ns −1 in the longitudinal profile (Fig. 6), and between 0.15 and 0.18 m ns −1 in the transversal profile #3. Such velocity values are consistent with those reported in the literature for buried ice (Brandt et al., 2007; Fukui et al., 2008). Stacking (RMS) velocity never reflects more than the bulk electromagnetic behavior of the material investigated and, as such, it would be premature to state that all velocities N0.15 m ns −1 in the models are evidence for buried massive ice. For example, such velocities would not be incompatible with an ice-quartzite clasts mixture owing to the low relative dielectric permittivity (K = 4.5) of quartzite. However, there is a strong basis for considering them as reflecting a high probability for massive ice occurrence. On this basis, the strongly varying velocity in the rock glacier is thought to reflect extreme contrasts in the constitutive materials. The low velocity values (b0.10 m ns −1) are, indeed, typical of wet deposits or bedrock and may be caused by the presence of water (running or in the form of pockets) in the deep rock glacier or
Table 1 Descriptive statistics from the velocity distribution models. Velocity (m/ns)
LP TP #1 TP #2 TP #3
Mean
Max
Min
Standard deviation
0.127 0.118 0.123 0.130
0.171 0.171 0.169 0.189
0.089 0.089 0.099 0.091
0.019 0.022 0.019 0.019
at/below the rock glacier floor. Such hypothesis is supported by the existence of an extensive central furrow (Fig. 2) along which it is often possible during summer to hear deep running water sounds from the surface. The remaining intermediate range of velocity values (0.10– 0.15 m ns−1) is large; it could theoretically correspond to either unfrozen dry deposits or permafrost (Daniels, 2004; Brandt et al., 2007). 4.3. Migrated and automatic gain control (AGC)-enhanced profiles: reflector patterns The migrated and AGC-enhanced GPR profiles (Figs. 7 and 8) offer a very high quality visualization of the Thabor rock glacier internal stratigraphy. The results of the 2-D Kirchoff migration are very satisfactory, since the hyperbolae initially present in the data are focused and do not perturb the interpretation. The GPR penetration depth is variable: generally 25–30 m, down to 40 m in the distal part of the rock glacier, but only a few meters in some areas of the upper rock glacier. The latter weak penetration is certainly explained by the considerable thinning of the superficial deposits and the shallow bedrock occurrence. There, the straight and 30–45° dipping features are interpreted as stratifications in the quartzite basement. Owing to the AGC enhancement, basal reflectors are well delineated, especially in the lower parts of the rock glacier, and are referred to here as rock glacier-underlying floor interfaces. Thus, the rock glacier thickness ranges from less than 10 m in the central (650–700 m position on the longitudinal profile) and upper parts, to 35–40 m in the distal part. On the whole, the stratigraphy is dense and salient, and the recognition of a few prominent and bounding reflectors highlights the specific stratigraphic organization of the rock glacier. In the longitudinal profile, apart from the attenuation areas, a shallower (15–20 m) part
Fig. 5. Modeled distribution of the stacking velocity obtained from the GPR common-offset data of the Thabor rock glacier. The main illustration is the model of the longitudinal profile. All transversal models (TP #1, TP #2, TP #3) are E (left)–W (right) oriented. See Fig. 2 for the exact locations of the profiles. The full black lines on the models delineate areas with velocity values N 0.15 m ns−1. The dashed black lines delineate areas with velocity values b 0.10 m ns−1. The models derived from the application of kriging techniques to local velocity determinations. The amount of local velocity determinations in the data is the following: 90 in the longitudinal profile (0.062 per horizontal meter), 37 in the transversal profile #1 (0.112 m−1), 50 in the transversal profile #2 (0.128 m−1), and 79 in the transversal profile #3 (0.119 m−1).
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section of the transversal profile #3, which corresponds to an adjacent deposit, shows the same type of stratigraphy, with undulating reflectors at depth and dipping, embedded, and curved reflectors close to the surface (Fig. 8). The significance of the stratigraphic patterns is crucial for our understanding of the rock glacier structure. For example, 1) Straight and parallel reflectors concordant with a basal one express low longitudinal deformation and/or vertical compaction; 2) Chaotic reflectors may reflect local downwasting and perturbation of the shallow structure; 3) Undulating and toplapping reflectors; and 4) upward- or rock glacier center-dipping reflectors rather evocate the occurrence of compressive stresses and thrusting. Indeed, undulating and toplapping reflectors are well known in rock glacier deformation structures (e.g., Degenhardt et al., 2003; Monnier et al., 2008), and the dipping reflectors observed here are analogous to GPR signatures given by thrust planes in glaciotectonized sediments (Overgaard and Jakobsen, 2001; Sadura et al., 2005) or by debris inclusions in glacier ice (Woodward et al., 2003) and debris-covered glacier ice structures in rock glaciers (Fukui et al., 2008). Fig. 6. Plotting of the radar wave velocity values in the longitudinal profile at a depth of 2.50 m with the locations of the 2003 massive ice exposures.
predominantly made of upward-dipping reflectors lies on or imbricates into a deeper part made of straight and parallel reflectors, concordant with the basal one. In the terminus part, undulating reflectors are preferentially encountered. The three transversal profiles have similar organization. Undulating and toplapping reflectors are present in the deeper part. In the central part, where the morphology is poorly differentiated or may consist of straight ramps, the reflectors are mainly chaotic or rock glacier center-dipping. On the lateral margins, where the morphology consists of arcuate ridges or straight ramps, the reflectors are dipping toward the center of the rock glacier. Moreover, the eastern
5. Discussion: structure and genesis of the Thabor rock glacier 5.1. Structural and stratigraphical units of the rock glacier By integrating the main data available – stratigraphy, radar wave velocity, and surface morphology – and according to a good correspondence between the velocity distributions and the main reflectors, four types of structural unit are identified along the entire GPR profiles (Figs. 7 and 8). Units of type 1 correspond to the bedrock floor. The bedrock floor connects with the surface in restricted sections of the upper rock glacier; it is very close to the surface halfway down the longitudinal profile, so that the rock glacier appears almost cut in two, with the upper part being quasi disconnected from the lower part. In the upper parts of the rock glacier the bedrock floor
Fig. 7. Stratigraphic patterns and structural units of the longitudinal profile (migrated, enhanced with AGC). The surface morphology is indicated in the upper part of the illustration. The units 1 to 4 are described as following: 1) Units generally disconnected from the surface and separated from other units by a basal reflector; they have poorly visible patterns of reflectors and generally (but not always) correlates low radar wave velocities (b0.10–0.12 m ns−1). 2) Units of the deep rock glacier, connecting to the surface by the lateral margins (or without connection with the surface), characterized by parallel/concordant, or undulating–toplapping reflectors, and a midrange radar wave velocity (0.10–0.13 m ns−1). 3) Units characterized by arcuate ridges or straight ramps on the surface, a pattern of center−/−up dipping (frequently curved) reflectors, and a quite high radar wave velocity (N 0.13–0.14 m ns−1). 4) Units characterized by a poorly differentiate surface morphology (but with straight ramps and crevasse-like features in some parts) a pattern of center-/updipping reflectors (indeed chaotic, in some cases), and a high radar wave velocity (N 0.14–0.15 m ns−1).
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Fig. 8. Stratigraphic patterns and structural units of the transversal profiles (migrated, enhanced with AGC). The surface morphology is indicated in the upper part of the illustrations. The description of the types of unit is given in the legend of Fig. 7.
is probably frozen (high radar wave velocities) while it is likely unfrozen and even penetrated by melting water in the lower parts (low radar wave velocities). The bedrock floor exhibits few structural features (dipping joints) in the longitudinal GPR profile. Units of type 2 constitute the “basal sole” of the rock glacier. This basal sole connects with the surface at the terminus and at the lateral margins. It is made of debris potentially frozen (typical radar wave velocities between 0.10 and 0.13 m ns −1). Its stratigraphic features (undulating and toplapping) are the expression of compaction and gentle compression. Units of type 3 constitute the “deformed margins” of the rock glacier. At the surface, the morphology is highly differentiated with arcuate ridges and, over the transition with the central part of the rock glacier, straight ramps. The deformed margins are made of
ice-debris mixtures with a large ice content and with possible inclusions of massive ice (radar wave velocity N0.15 m ns −1 in several parts), which correlate with observations at the surface (Fig. 3). The stratigraphic features (strongly curved and dipping reflectors) are the expression of strong compression. Finally, units of type 4 constitute the “shallow debris-covered ice core”. The ice is thought to be massive (radar wave velocity typically between 0.15 and 0.17 m ns −1) but with a high debris content. Indeed, the gently curved and dipping reflectors are best interpreted as debris inclusions along thrust planes in the ice. The architectural model of the rock glacier appears, thus, as one of a debris-covered massive ice core superimposed on, or embedded into, a sediment complex. The debris-covered massive ice core is the remnant of the former Thabor glacier, which stationary
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and re-advancing positions since the end of the LIA can finally been inferred (Fig. 9). 5.2. Origin of the surface features Given our structural interpretation, the surface morphology (Fig. 2) may be interpreted as being mainly the result of glacier– permafrost interactions. As primarily inferred, the hummocks and very upper lateral ridges are of glacial origin. The occurrence of crevasse-like features at the surface corresponds with the inferred extent of debris-covered ice, hence these features probably relate to fractures within the massive ice core. In the lower-central part of the rock glacier, a large convex slope, preceding a crevasse-like feature, evocate an accumulation form and may correspond to a former terminus of the stagnating glacier (Fig. 9). Straight ramps occur both in the deformed margins and in the debris-covered ice core; they are thought to result from downwasting by melting of internal ice. The origin of the arcuate ridges on the Thabor rock glacier was extensively discussed by Monnier (2007). Further understanding of these features can be gained by focusing on the terminus part of the rock glacier in the longitudinal profile (Fig. 10). The location of the ridges correlates exactly with the dipping planes. Some of the ridges, especially number 4 (Fig. 10c,d), exhibit a peculiar morphology with a steep and sharp back slope preceding a flat and depressed area. This morphology was
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first interpreted as the result of differential ablation processes (Monnier, 2007). However, the observed connection between the back slope of the ridge and the dipping reflector leads to a new interpretation: the back slope is a thrusting plane-related scarp, and the entire ridge is the expression of a compressive deformation process. Thus, considering the structure of the rock glacier as a whole and the peripheral, concentric distribution of the ridges, the so-called deformed margins correspond to a belt of frozen/ice-cored sediments (potentially classifiable as moraines) deformed by ice push. The deformations (internal thrusting and surface ridges) may have been subsequently increased by the displacement of the rock glacier tongue in the terminus part. 5.3. History of the Thabor rock glacier At the end, the following genesis model can be established. During the Little Ice Age, as attested by the Carte d'État Major (1847–1864), the Thabor glacier advanced over and deformed preexisting deposits in discontinuous permafrost conditions. Lateral and terminal arcuate ridges were created – or accentuated if they preexisted – by ice pushrelated compressive stresses, perhaps through several cycles of short advances and retreats (Fig. 9). The same conclusion was proposed by Lugon et al. (2004) for rock glaciers in the Posets Massif, Spanish Pyrenees. The preexisting deposits mentioned are likely to be a
Fig. 9. Spatial extent (on 3-D Google Earth sketch) of the debris-covered massive ice in the Thabor rock glacier as determined by matching the structural interpretation of the GPR profiles and the surface morphology. The upper area with thin debris cover over ice or bedrock is also depicted. Consequently, the former terminus positions of the Thabor glacier since the end of the LIA can be inferred using the historical maps and the morphology.
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Fig. 10. Focus on the terminus part of the longitudinal profile showing the relationship between the arcuate ridges and the curved dipping reflectors. a) Numbering and position of the arcuate ridges relative to the longitudinal GPR profile. b) Zoom on the terminus part of the longitudinal profile; the numbers are the same as those in a). c) Detail of an arcuate ridge with a steep and sharp back slope preceding a flat and depressed area. The backpack gives the scale. d) Interpreted version of the box in b).
previous glacial forefield. Indeed, it is well known in the Alps (Patzelt, 1974; Le Roy Ladurie, 2003) and especially in their occidental part (Deline and Orombelli, 2005), that several glacial phases analogous in spatial extent to the LIA occurred during the ten thousands of years preceding the LIA. Moreover, in the valley down the Thabor rock glacier, the only visible glacial deposits are lateral moraines at an elevation of 2450 m that were attributed to the Late-glacial period (Monnier, 2006). The “basal sole” of the rock glacier as well as some parts of the “deformed margins” are, therefore, thought to have been deposited (even shaped?) by one or several glacial advances of the Thabor glacier taking place between the end of the Late-glacial and the beginning of the LIA. Following the end of the LIA and under the progressively warmer climatic conditions, the Thabor glacier began retreating and burying, resulting in the superimposition of the debris-covered ice core onto the preexisting accumulation. The origin of the burial of the ice remains uncertain. It may be related to repeated rock falls in the accumulation area of the former Thabor glacier, which supplied a considerable stock of debris to the glacier–rock glacier system. Indeed, large and clear failure scars are visible on the rock walls of the Cheval Blanc and the Mt. Thabor, and several sudden rock falls were observed during field trips at the site. Additionally, melting out of internal debris in the ice was possible. Considering the occurrence of both origins is appealing: climate warming melts the glacier and the adjacent rock walls and, at the end, produces a rock glacier. Irrespective of the burial processes, the retreat of the Thabor glacier led to the apparition of the central depression in the upper part of the rock glacier. Furthermore, when the burial/retreat of the glacier reached an advanced level (in the 20th century), new dynamics
occurred: the lower part of the rock glacier began to move by permafrost creep, and the upper part began to settle by melting of the buried ice, the volume of which was especially important in the central part (and at the origin of the network of straight ramps). Completely disappeared from the surface at the turn of the 20th and 21st centuries, the Thabor glacier left in the landscape a disconnection between the debris source on Mt Thabor and the rock glacier: the depression at the upper end of the rock glacier reached its maximum dimensions and the bedrock started to appear through thin, discontinuous remnants of the debris-covered glacier. 6. Conclusions We used a complete set of historical, geomorphological, and geophysical data to investigate the structure and genesis of the Thabor rock glacier. The extensive common-offset GPR data collected at the site were used to model the distribution of the stacking radar wave velocity and reveal the stratigraphy of the rock glacier. The GPR results were correlated with the morphological analysis of the surface. Our study strengthens previous morphological analyses (Monnier, 2007) and the following important points deserve to be highlighted. 1) The 2-D models of the radar wave velocity show a very heterogeneous structure, with velocities ranging between 0.08 and 0.18 m ns −1. Following comparison with direct observations of the internal structure, velocities higher than 0.15 m ns −1 can be considered as reflecting a high probability for massive ice occurrence. 2) The stratigraphy of the rock glacier is varied, with most features expressive of compressive stresses and thrusting, and prominent bounding reflectors dividing the structure. The integration of velocity
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modeling, stratigraphic representation and surface morphology information depicts a rock glacier composed of a shallow (b15 m) debris-covered massive ice core overlying a (potentially frozen) sole of sediments, and embedded into deformed permafrost margins. 3) The genesis of the rock glacier has probably involved several temporal phases since the Late-glacial; however, the current rock glacier morphology is mainly the product of glacier–permafrost interactions during the LIA and subsequent deformation processes. During the LIA the Thabor glacier advanced over and deformed preexisting deposits before the resulting system became modified in the last 150 years by ice burial, permafrost creep, and downwasting. At the end, is the Thabor rock glacier still a true one? Haeberli (1979, p. 46) states that “All transitions between push-moraines and rock glaciers may exist in cases where a glacier was in contact with a rock glacier during neo-glaciation or Holocene advances”. This is exactly where the study of the Thabor rock glacier leads us. The actual and impressively esthetic tongue is neither a debris-covered glacier, nor a morainic complex, nor a rock glacier in the sense of an individual creeping permafrost body (Barsch, 1992). The Thabor rock glacier is all of that in one feature, and thus expresses the complexity of the interactions that can occur on the surface as consequences of the environmental changes of the last centuries and millennia. Acknowledgments The authors would like to thank friendly and sincerely Marion Siboni and Guillaume Tanguy who provided valuable and courageous help during the field survey. Moreover, the Thabor rock glacier is part of the Thabor natural site listed by the Ecology and Sustainable Development Minister of France; the temporal implantation of scientific material on this site was permitted by a decree of the Savoy prefecture emitted the 30th of June of 2008. The authors renew their acknowledgements to the Savoy prefecture for this permission. Finally, Shelley MacDonell provided helpful review of the language in the manuscript. References Annan, A.P., 2003. Ground Penetrating Radar principles, Procedures and Applications. Sensors & Software Inc., Mississauga. Barsch, D., 1992. Permafrost creep and rockglaciers. Permafrost and Periglacial Processes 3, 175–188. Bradford, J.H., Nichols, J., Mikesell, T.D., Harper, J.T., 2009. Continuous profiles of electromagnetic wave velocity and water content in glaciers: an example from Bench Glacier, Alaska, USA. Annals of Glaciology 50, 1–9. Brandt, O., Langley, K., Kohler, J., Hamran, S.-E., 2007. Detection of buried ice and sediment layers in permafrost using multi-frequency Ground Penetrating Radar: a case examination on Svalbard. Remote Sensing of Environment 111, 212–227. Brown, J., Nichols, J., Steinbronn, L., Bradford, J., 2009. Improved GPR interpretation through resolution of lateral velocity heterogeneity: example from an archaeological site investigation. Journal of Applied Geophysics 68, 3–8. Ground Penetrating Radar, In: Daniels, D.J. (Ed.), 2nd Edition. The Institution of Electrical Engineers, London. 726 pp.
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