8.24 Evolution of Slopes in a Cold Climate

8.24 Evolution of Slopes in a Cold Climate G Rixhon, University of Cologne, Cologne, Germany A Demoulin, University of Lie`ge, Sart Tilman, Lie`ge, Belgium r 2013 Elsevier Inc. All rights reserved.

8.24.1 8.24.2 8.24.2.1 8.24.2.2 8.24.2.2.1 8.24.2.2.2 8.24.3 8.24.3.1 8.24.3.2 8.24.3.2.1 8.24.3.2.2 8.24.3.3 8.24.3.4 8.24.3.4.1 8.24.3.4.2 8.24.3.4.3 8.24.3.5 8.24.3.5.1 8.24.3.5.2 8.24.4 8.24.4.1 8.24.4.2 8.24.4.3 8.24.4.4 8.24.4.4.1 8.24.4.4.2 8.24.4.4.3 8.24.5 8.24.5.1 8.24.5.2 8.24.5.3 8.24.5.4 8.24.6 References

Introduction Cryoplanation Mechanism and Landforms Cryopediments Cryoplanation Terraces Morphology and controlling factors Processes and (paleo)climatic significance Talus Slopes, Including Stratified Slope Deposits Definition and Geographic Distribution Morphology and Processes Rockfall as primary process and resulting morphology Secondary reworking processes: Debris flows and snow avalanches Surface Processes and Constitutive Materials Stratified Slope Deposits within Talus Landforms (Including the Gre`zes Lite´es) Definition and geographical distribution Bedding and sedimentary structures Forming processes Talus-Slope Evolution Rates of debris supply/rockwall retreat and paraglacial activity Paleoenvironmental implications of talus slopes and stratified slope deposits Blockfields Definition and Geographic Distribution Form Characteristics Block Accumulation Processes Origin, Age, and Paleoenvironmental Significance of Autochthonous Blockfields The Quaternary periglacial model – short to mid-term formation The pre-Pleistocene (Neogene) model – long-term formation A unified scheme of blockfield formation: Neogene inheritance and Quaternary development Block Streams Definition and Geographic Distribution Form Characteristics Processes (Paleo)Climatic Meaning and Dating of Relict Block Streams Research Perspectives

Glossary Marine Isotopic Stage (MIS) (Also named oxygen isotope stages (OIS)) designate alternating cold (even numbers) and warm (odd numbers) periods in the Earth’s paleoclimate, respectively, referring to glacial (cold) and interglacial (temperate) stages. They are inferred from variations in the oxygen isotope ratio (d18O) reflecting temperature curves and measured in deep sea cores.

393 393 394 394 394 395 395 395 397 397 397 397 399 399 399 399 401 401 402 403 403 403 405 405 405 406 407 407 407 408 410 410 411 412

Nivation A concept referring to the weathering and erosion of land surfaces associated with a seasonally late-lying or perennial snowcover. It encompasses the combined action of frost shattering, gelifluction, and slopewash processes (thought to operate in the vicinity of snowbanks). Optically stimulated luminescence (OSL) A dating method determining how long ago certain minerals (quartz

Rixhon, G., Demoulin, A., 2013. Evolution of slopes in a cold climate. In: Shroder, J. (Editor in chief), Giardino, R., Harbor, J. (Eds.), Treatise on Geomorphology, Academic Press, San Diego, CA, vol. 8, Glacial and Periglacial Geomorphology, pp. 392–415.

392

Treatise on Geomorphology, Volume 8

http://dx.doi.org/10.1016/B978-0-12-374739-6.00218-9

Evolution of Slopes in a Cold Climate

and feldspar) were last exposed to daylight. It makes use of electrons, which become trapped in crystal structure defects or ‘holes’ in the mineral crystal lattice. The luminescent signal is a reflection of the number of trapped electrons, which are released from traps by shining a beam of light onto the mineral. Terrestrial or in situ cosmogenic nuclides (TCN) Nuclides (3He, 10Be, 26Al, 36Cl, etc.) produced by the interaction of

393

secondary cosmic rays with exposed target atoms of minerals at the Earth surface. They are used as dating method (surface exposure dating) by measuring their concentration within surficial materials (bedrock, sediment, etc.) as their abundance is proportional to the time that the surface has been exposed to cosmic-ray activity.

Abstract This chapter focuses on the description and the discussion of main periglacial slope forms. They encompass landforms related to an erosional context (cryoplanation) and landforms associated with a ‘depositional’ context (debris accumulation). Emphasis was put on periglacial block accumulations extensively studied in recent times: talus slopes, blockfields, and block streams. For each treated landform, an introduction with the geographical distribution is first presented. Thorough descriptions of landforms and constitutive deposits are then provided, followed by an overview of the processes discussed in the literature. At the last, the authors gave attention to the chronologic evolution of the landforms and their paleoclimatic significance.

8.24.1

Introduction

Periglacial slope processes and the resulting slope forms are generally controlled by the presence of frozen or thawing ground (French, 2007). The processes range from periglacial weathering to mass wasting, either fall or flow processes, delivering, and transporting variable amounts of debris and shaping the slopes in various ways (see Chapter 8.3). Frost wedging and frost shattering (also called macrogelivation) represent the main mechanical weathering agents under periglacial conditions (see Chapter 4.15). Mostly resulting from seasonal and daily freeze/thaw cycles, they are thought to produce large quantities of angular block debris (occurring, e.g., in blockfields and block streams). Mechanical weathering on exposed rockwalls also induces gravitational mass movements, mainly rockfalls (giving birth to talus slopes), whereas snow avalanches are believed to play a subsidiary role. Flowing masswasting processes are diverse, depending especially on the thickness and the moisture degree of the thawing active layer. They include solifluction sensu lato, contributing to the formation of stratified slope deposits and block streams through solifluction, (con)gelifluction, and frost creep (see Chapter 7.13), debris flow involved in the formation of talus slopes, and slush flow or slopewash in case of an impermeable permafrost table overlain by a shallow active layer. Although a gradual and sequential reduction of the relief is commonly thought to represent an important evolutive trend of periglacial hillslopes, no slope form or sequence should be, however, considered as uniquely and exclusively periglacial in nature (French, 2007). For instance, the morphological similarity between cryopediments and pediments developed in warm semiarid environments suggests that the former might be partially inherited features. More importantly, many periglacial landforms are inactive or weakly active today. They may be regarded as relict features indicating that the periglacial landscape is generally in disequilibrium with its presentday cold-climate conditions (e.g., Andre´, 2003). It is now recognized that periglacial slope processes may operate on preexisting landforms (Harris, 2007), inherited from either

former glacial action (paraglacial adjustment, see Ballantyne, 2002) or pre-Quaternary chemical weathering processes (see Chapter 7.29). Given the diversity of periglacial slope processes and landforms, this chapter cannot pretend to be exhaustive. It rather focuses on the description and the discussion of major slope landforms and deposits that were sometimes hotly debated in the literature (rock glaciers being separately treated in Chapter 8.16). They can be roughly divided in two groups: erosional landforms associated with cryoplanation and landforms resulting from debris accumulation. We put special emphasis on the recently much studied block accumulations, which include talus slopes, stratified slope deposits, blockfields and block streams. Block streams differ from blockfields in that they imply a longitudinal, downslope transport of the rock debris. Talus slopes form at the foot of steep rockwalls, primarily by rockfal. In some cases, the debris accumulation displays a typical layering, forming stratified slope deposits. For each treated landform, geographical data are first provided, then a detailed description of forms and deposits and a brief overview of the processes discussed in the literature. We also give a briefing about the temporal evolution of the landforms, including the few dates available, and their paleoenvironmental meaning (see Chapter 8.26).

8.24.2

Cryoplanation Mechanism and Landforms

Cryoplanation is a traditional concept recurrently evoked in the periglacial literature about landscape and slope evolution under nonglacial conditions, and implying a progressive and sequential reduction of relief with time (see French, 2007). This model refers to a general mechanism creating low-angled and level bedrock surfaces (e.g., Dylik, 1957; Washburn, 1979; Grosso and Corte, 1991). Two main kinds of landforms are commonly believed to derive from this global mechanism: cryopediments and cryoplanation terraces. Both of them are presented and discussed in this section.

394

Evolution of Slopes in a Cold Climate

However, although the widespread existence of such landforms is no more questioned, doubt has recently been cast on the meaning of cryoplanation as a distinct morphogenetic process (Hall, 1998; Thorn and Hall, 2002). Indeed, studies about cryoplanation abundantly refer to nivation processes as the driving force, notably for backward development of landforms. Both concepts seem unrelentingly interconnected and it has been proposed that cryoplanation and nivation represent two end-members of a single processlandform continuum (Hall, 1998; Hall and Andre´, 2010). Cryoplanation would correspond to the dry, permafrost-derived end member and nivation to the wetter, snow-derived one, the moisture degree being thus the differential factor (Hall, 1998). This seems to require at least an updating of the cryoplanation concept applied to periglacial landscape evolution, which needs new data from presently active cryoplanation surfaces (Thorn and Hall, 2002).

8.24.2.1

Cryopediments

Mostly occurring in unglaciated arid periglacial environments (e.g., central Siberia, northern Yukon Territory, see Czudek and Demek, 1973; Priesnitz, 1981; French and Harry, 1992), cryopediments are extensive (up to tens of km in length), lowangled (o101) erosion surfaces (Figure 1), generally planar or concave-up in profile (Priesnitz, 1981). They characteristically truncate the underlying geological structures below a thin (o1m) veneer of locally derived colluvial material (Figure 1, French and Harry, 1992; Czudek, 1993; Vandenberghe and Czudek, 2008) that commonly contributed to making them to be considered as transportation surfaces (French and Harry, 1992). Sediment transport and removal from the cryopediment surfaces have been notably attributed to surficial runoff, acting either as rill or sheetwash (Vandenberghe and Czudek, 2008). Although Dylik (1957) thought that such pediments were similar to those occurring in warm semiarid environments, many authors believe that they result from a specific cryogenic activity (Czudek and Demek, 1973; Priesnitz, 1981). A mechanism of parallel scarp retreat caused by frost action and extending the pediment upslope was proposed, occasionally complemented by downwearing in nonresistant substratum (Czudek, 1993). The role devoted to frost action (frost wedging and nivation) should nonetheless be downplayed in arid continental environments (French and Harry, 1992; French, 2007). In central Europe, the cryopediments are assumed to have developed in relation with the presence of a deep permafrost (see Chapter 8.14), mainly during the Late Pleistocene (Czudek, 1993). French and Harry (1992) also claimed that cryopediments in northern Canada were mostly inactive under the current climatic conditions, adding that their episodic activity took place in every Quaternary period when high moisture levels in the substratum favored severe frost shattering of upper slopes and enhanced sediment transport by solifluction. Similar conclusions were reached for the cryopediments of central and northern Europe, where the periodic return of periglacial conditions combined with a gradual baselevel lowering induced the development of polycyclic features

Figure 1 Dissected cryopediment in the Barn mountains (Northern Yukon Territory, Canada). The truncation of the underlying geological structure and the thin veneer of colluvial material at the surface should be noticed. Reproduced from French, H.M., Harry, D.G., 1992. Pediments and cold-climate conditions, Barn Mountains, unglaciated northern Yukon, Canada. Geografiska Annaler A 74, 145–157.

(Vandenberghe and Czudek, 2008). An alternative interpretation of the cryopediments, much in the sense of Dylik (1957), has however been proposed in the Northern Yukon Territory, making them derive from Tertiary pediments remodeled under periglacial conditions (French and Harry, 1992; French, 2007).

8.24.2.2 8.24.2.2.1

Cryoplanation Terraces Morphology and controlling factors

Likewise occurring within unglaciated periglacial environments (e.g., northern Yukon Territory, see Lauriol and Godbout, 1988; Lauriol et al., 1997), cryoplanation terraces, also called altiplanation or nivation terraces, are nearly horizontal or smoothly inclined bedrock-cut benches (Figure 2; French, 2007). They are intimately related to cryopediments but generally occur at higher relative elevations (Lauriol and Godbout, 1988), in some places even on ridge tops (‘cryoplanation summit flat’, see Reger and Pewe, 1976; Czudek, 1995). A typology of these features was consequently proposed according to their respective topographic situation (Nelson, 1998). Although cryoplanation terraces display slope angles comparable to those of cryopediments (from 11 to 121, French, 2007), they are generally of smaller size, not exceeding a few hundred meters in length (Figure 2). They display various planform shapes (elongated, sicklelike, etc.) but a rather constant profile, with a gently sloping surface backed by a steep riser or a scarp (from 1–2 to 20 m, even up to 75 m in height) that marks the upper end of the terrace (Figure 2, Reger and Pewe, 1976; Nelson, 1989). Stepped cryoplanation terraces (up to 25 levels, see Czudek, 1995) mimicking fluvial terrace sequences may also occur (Figure 2; Hall and Andre´, 2010). Cryoplanation terraces may be carved into every rock type (Grosso and Corte, 1991). They are however best developed on fine-grained and closely jointed rocks susceptible to blocklike disintegration, such as basalt, diorite, gabbro, granite, and quartzite (Reger and Pewe, 1976; Czudek, 1995). In

Evolution of Slopes in a Cold Climate

395

T9 T8 T7

T6

Syrtis hill T9 37 m

T8 3 m T7 2m

T6 63 m

T5 46 m

T4 T3 26 m

T2 T1

21 m 65 m

32 m Not to scale Figure 2 Cross-section displaying stepped cryoplanation terraces (nine levels) carved in the northwest facing slope of Syrtis Hill (Alexander Island, Antarctica), with a closer view on the upper sequence of terraces. Note the small dimensions of the terraces (not exceeding 70 m in length), the gently sloping surface, terrace scarps and the general stepped profile mimicking a fluvial terrace sequence. Reproduced from Hall, K.J., Andre´, M.-F., 2010. Some further observations regarding ‘‘cryoplanation terraces’’ on Alexander Island. Antarctic Science 22(2), 175–183, with permission from Cambridge University Press.

Antarctica, field observations indicate that terraces grew from lithological junctions associated with local facies variations in sandstones (Hall and Andre´, 2010). The first-order (but nonexclusive) controls on the formation of cryoplanation terraces seem indeed to be the underlying bedrock structure and the slope aspect (Figure 2; Lauriol and Godbout, 1989; Czudek, 1993, 1995; Hall and Andre´, 2010). In Alaska for instance, north-facing slopes are best oriented for terrace development, probably in relation to the wind patterns dominating during cold periods (Nelson, 1989, 1998). In the northern Yukon Territory, the predominance of cryoplanation terraces on NWand SE-facing slopes is not only related to the prevailing wind directions but also reflective of the orientation of major structural axes in this area (Lauriol and Godbout, 1988). As for the probable link between cryoplanation terraces and the presence of a permafrost, it is still a debated issue, with pros (e.g., Reger and Pewe, 1976) and cons (e.g., Demek, 1969).

8.24.2.2.2

Processes and (paleo)climatic significance

Like cryopediments, cryoplanation terraces are thought to evolve mostly through backwearing (Czudek, 1995). Various processes acting in combination, such as solifluction and gelifluction, frost weathering and sorting, slope- and sheetwash or even piping, have been cited in the literature (see Lauriol and Godbout, 1988; Czudek, 1993). Their relative efficiency depends on the magnitude and frequency of freeze/ thaw cycles but the presence of snow also appears as definitely crucial (Czudek, 1995). As stated above, nivation has been recognized as the main process involved in scarp retreat and, consequently, in cryoplanation terrace development (Reger and Pewe, 1976; Nelson, 1989, 1998; Czudek, 1993). However, initial nivation hollows may evolve in cryoplanation terraces even on surfaces without any break of slope, the snow supplying the large amounts of water needed to further develop the landforms (Czudek, 1995). Isolated snow patches subsisting late into the summer are believed to control the

effectiveness of nivation (Nelson, 1998). However, the weathering component of the process remains contentious and requires further clarification (see Thorn and Hall, 2002). In Alaska and the northern Yukon Territory where cryoplanation terraces were studied in detail (Reger and Pewe, 1976; Lauriol and Godbout, 1988; Nelson, 1989, 1998), it has been shown that, except for a few places where they might be currently active (see Lauriol and Godbout, 1988), all cryoplanation terraces are relict landforms, as indicated by the abundant vegetation and the lichens covering their surface (Nelson, 1998). This conclusion can be extended to all features in the temperate latitudes of Europe and North America (Reger and Pewe, 1976). By contrast, cryoplanation terraces located in the current periglacial areas of eastern Siberia seem to be active nowadays (Czudek, 1995). Cosmogenic nuclide exposure dating (36Cl) performed on two cryoplanation summit flats in Missouri Mountains (Cremeens et al., 2005) yielded Middle Pleistocene ages (marine isotopic stages (MISs) 9 and 14), showing that they had experienced periglacial conditions during at least two cold periods. However, such dates are still the exception, so that the age of cryoplanation features remains largely unknown (Lauriol and Godbout, 1988; French, 2007).

8.24.3 8.24.3.1

Talus Slopes, Including Stratified Slope Deposits Definition and Geographic Distribution

Scree (English terminology) or talus (American terminology) slopes are landforms induced by mass-wasting processes. They are formed by accumulations of coarse, loose, usually angular rock debris at the foot of steep, bare, rock slopes (Figures 3(a) and 3(b), Luckman, 2007). Although other names were also given to these landforms (e.g., debris slopes (Gardner, 1980), colluvial fans (Blikra and Nemec, 1998)), the two

396

Evolution of Slopes in a Cold Climate

synonymous terms ‘scree’ and ‘talus’ equally refer to the landform and to its constitutive materials. Talus slopes and the correlated processes have been studied extensively in many and diverse periglacial and alpine settings (see Chapter 7.31). These regions encompass arctic and mountainous areas, such as the Alps (Francou and Mante´, 1990; Jomelli and Francou, 2000; Otto and Sass, 2006; Sass, 2006, 2007; Sass and Krautblatter, 2007; Pieracci et al., 2008), the Andes (Perez, 1993), Baffin Island (Church et al., 1979), Canadian Rockies (Luckman, 1978, 1988; Gardner, 1980; Luckman and Fiske, 1997), New Zealand (Hales and Roering, 2005), Que´bec (He´tu and Gray, 2000), Norway (Blikra and Nemec, 1998; McCarroll et al., 1998), Scotland (Luckman, 1992; Hinchliffe et al., 1998; Hinchliffe and Ballantyne, 2009), Svalbard (Rapp, 1958; Andre´, 1997), Sweden (Rapp, 1960) and Wales (Curry and Morris, 2004). Talus slopes are closely related to environments where physical weathering

and gravitational processes prevail (see Chapter 7.4) and where the balance between production and removal of debris is highly positive, so that thick accumulations evolve in a distinct morphology, whatever be the slope and the underlying bedrock. Furthermore, given the coarseness of most talus deposits, this morphology is resistant to posterior erosion processes and talus slopes are generally considered as stable and long-lasting landscape elements, many of them being probably preserved as inactive or relict forms (Figure 3(c), e.g., Hales and Roering, 2005). Owing to the highly variable morphology, composition, and geomorphologic settings of talus slopes, their typology relies primarily on the morphology and the composition of the original cliff, the geomorphic processes involved in their development (rockfall, debris flow, etc.), and the topography/morphology of the accumulation zone at the foot of the rock slope (Figure 3).

Figure 3 Various forms of talus-slope induced by the morphology of the original cliff and the processes involved: (a) Typical sheet and regular talus slope developed at the foot of subvertical cliff of Mount John Laurie (Alberta, Canada). Photo taken by Keven Lenz, 2006 (Creative Commons license). (b) Sheet talus slope primarily formed by rockfalls at Small River (British Columbia, Canada). The straight upper segments contrasts with the slight concavity at the base. The arrow indicates a person for scale. (c) Partially relict (grass covered) talus slope at Lairig Ghru (Cairngorn Mountains, Scotland), reshaped by snow avalanche activity, which produces avalanche boulder tongue visible on the centre and the right and finally reworked by hillside debris flows. (d) Succession of coalescing talus cones resting on raised beaches (Templefjorden, Svalbard). The cone on the left displaying lower slope angle and traces of debris flow reworking strongly point to a multiprocessed morphology. (e) Strong reworking activity of debris flow on coalescing talus cones (Tanquary Fjord, Canada). Gullies dissecting the upper part of cones heading at the cliff base, debris flow-generated levees, terminal lobes and the lower mean slope angle of the sharply reworked ‘alluvial talus’ in the centre should be noticed. (b)–(e) Reproduced from Luckman, B.H., 2007. Talus Slopes. In: Elias, S.A. (Ed.), Encyclopedia of Quaternary Science/Periglacial Landforms, Rock Forms. Elsevier, Amsterdam, pp. 2242–2249.

Evolution of Slopes in a Cold Climate 8.24.3.2 8.24.3.2.1

Morphology and Processes Rockfall as primary process and resulting morphology

Generally defined as the free or bouncing fall of rocky materials down a steep slope (Luckman, 2007), rockfall is believed to be the primary process producing talus slopes (Figure 4). It covers diverse kinds of processes, ranging from massive and catastrophic failures (also named rock valanches) to the fall of individual clasts (see Chapter 7.17). According to Rapp (1960), relatively small rockfalls (o10 1–102 m3) would be mainly responsible for the formation of talus slope in two different ways. Primary rockfalls mainly occur on uniform, massive, and (sub)vertical rock outcrops, or cliffs (Figures 3(a) and 3(b)). Triggered by freeze–thaw alternations, pressure release, or weathering of the outcrop face, they are characterized by a direct downslope transfer of newly detached rock debris. Secondary rockfalls take place on more complex rock slopes where debris may accumulate on rock face irregularities and be dislodged by subsequent rockfalls, surface water flow, snow avalanches, etc. When produced by rockfalls, a rockfall talus slope is the result of repetitive events that caused an accumulation over a long period of time. The plan view morphology of rockfall talus slopes depends not only on the planform of the cliff face delivering the rock debris, but also on the morphology of the footslope surface where the debris are accumulating (Luckman, 2007). Commonly lacking significant lateral variation of facies, sheet, and regular talus slopes (Figures 3(a) and 3(b)) essentially develop at the foot of uniform cliffs cut in a single rock type. By contrast, where the cliff face displays a more complex or dissected form, the rockfalls tend to be channeled into accumulation corridors that open downslope into talus cones (Figures 3(d) and 3(e)). In this case, other transport processes (e.g., stream flow, etc.) are also channeled into these corridors, leading to a potentially significant reworking of the talus cone (Figure 3(e), see below). Therefore, talus cones formed at the foot of dissected cliffs are generally a multiprocessed morphology (Figures 3(d) and 3(e)). In profile, talus slopes are generally characterized by long, straight segments at an angle of 33–351, commonly terminated by a marked basal concavity (Figure 4). Including this concavity, mean talus-slope gradients are generally comprised of slopes between 251 and 301. However, some authors construe the mean angle variability to be a consequence of diverse process combinations (Church et al., 1979; Sauchyn, 1986). It has also been shown that the steepest segments occur on fresh rockfall talus whereas talus reworked by processes such as snow avalanches (see below) display lower gradients (Jomelli and Francou, 2000). Detailed measurements of talus-slope angles have reported values as high as 401 for the straight upper part of the talus profile and 32–371 for the mean gradient (Sauchyn, 1986; Francou and Mante´, 1990). According to Francou and Mante´ (1990), talus slopes display a split-up profile, with a straight upper part mainly devoted to transport and, below a threshold angle of 33–341, a lower concave section resulting from both transport and depositional processes (Figure 4). It must also be noted that length and gradient of the lower concavity depend on various factors, such as the maturity of the talus evolution

397

(Figure 4), the site characteristics and the depositional processes involved.

8.24.3.2.2

Secondary reworking processes: Debris flows and snow avalanches

Small gullies dissecting an existing talus slope and heading at the base of the bare rock slope may be used as transport pathways by debris flows generated upslope by heavy rainstorms (Figure 3(e), see below). The reworking of significant amounts of debris ultimately produces an irregular pattern of debris flow deposits extending over the lower part of the talus slope and substantially covering the earlier deposited rockfall materials (Figure 4). Specific features characterizing this morphology are, e.g., levees margins and terminal lobes (Figures 3(c) and 3(e)). However, multiple reworking of these debris flow sediments by later debris flows and snow avalanches may also eventually lead the talus slope to evolve toward an ‘alluvial talus’ (Figure 3(e)), susceptible to turn into alluvial cones in the long-term. According to Rapp (1960), snow avalanches in alpine or arctic settings may remove significant amounts of materials from the cliff face and deposit them on the talus slope below when they slide down. In case of repeated heavy avalanches, the debris may be carried far beyond the rockfall talus boundaries, resulting in the formation of ‘avalanche boulder tongues’ that exhibit an asymmetric cross-section, a lobate front and steep side slopes (Figure 3(c), Rapp, 1958; Luckman, 1978; Jomelli and Francou, 2000). Therefore, avalanchereworked talus slopes are characterized by a marked concavity at the base of the slope (Figure 3(c)) and commonly by a considerable extension beyond the toe of the normal rockfall talus.

8.24.3.3

Surface Processes and Constitutive Materials

Typical talus slopes are characterized by a veneer of coarse debris and large boulders spread over their surface. Because of the sieving effect of such an openwork surface, fine elements, and smaller rock clasts arriving onto the talus surface are washed or trapped into large interstitial voids where they accumulate and progressively fill the interstitial space at depth. In this respect, it has been shown that rockfall or niveoaeolian processes generate considerable quantities of fine and very fine sediments that are deposited on talus slopes (He´tu and Gray, 2000). Creeplike displacements of individual elements resting on the talus surface are the result of a combination of various processes, such as rockfall impact, frost action, needle ice (pipkrakes), running water, snow avalanche, etc. Long-term studies have used painted markers to monitor the creeping displacements of individual clasts on upper talus slopes (e.g., Gardner, 1980; Perez, 1993). They reported mean motion rates of 1–10 cm year 1 for a mixed cover of small clasts within a fine matrix, noting that the higher values occur where more fine sediments are present. Although rockfall or avalanche impacts may cause very long displacements of individual clasts, transport distances of the largest elements on the lower slopes are much smaller because they depend on repeated dislodgment by direct impact. However,

398

Evolution of Slopes in a Cold Climate

Downlap t0

k

Slope knickline Grain flows, cohesive debris flows, creep lobes Deposits of cohesive debris flows and torrential floods Rockfalls, alluvial fan deposits

t1

t0 (a)

Increments of cliff retreat t3

t2

t1

t0

Proximal slope segment

t4

35

-30

t3

°

Distal slope segment(s)

t2 k

20-

t1

10° k

t0 (b)

Figure 4 Idealized sketch of talus slope morphology, deposits, and development. In an early stage (a), debris supply is dominated by rockfall. In a second stage (b), other processes (in particular debris flows) are then involved in talus development, also characterized by rockwall retreat. Under continued debris input, talus slope progrades and shows a split-up profile, with a straight upper part (proximal slope segment) essentially devoted to transport and a lower concave section (distal slope segments) resulting from both transport and depositional processes. Reproduced from Sanders, D., 2010. Sedimentary facies and progradational style of a Pleistocene talus-slope succession, Northern Calcareous Alps, Austria. Sedimentary Geology 228, 271–283.

one commonly finds a run-out area or basal fringe beyond the talus foot, where the largest outrunning boulders escaped the talus limits before coming to rest (Blikra and Nemec, 1998). Talus-slope deposits mostly consist of irregular and angular rock debris characterized by a wide range of sizes. Size and shape are primarily determined by the bedding and joint properties of the exposed rock. Most talus slopes display a

specific ‘fall sorting’, i.e., an increase of the surficial clast size from the upper part to the lower section of the slope. This is particularly well expressed in the basal part of the talus, where clast size usually depends logarithmically on distance (e.g., Statham, 1976; Church et al., 1979). Although the sorting degree may vary as a function of the cliff height, the dominant clast size and shape, and the length of the talus-slope, fall sorting mostly results from two main mechanisms. Firstly,

Evolution of Slopes in a Cold Climate

because of their higher momentum, larger elements (boulders, entire rock panels) tumble down greater distances by rolling, sliding or bouncing over the slope. Secondly, the roughness of the surface that the boulders travel along also plays an important role (Luckman, 2007), as the frictional resistance to displacement depends on the ratio between the dimension of the moving boulder and that of the slope surface irregularities caused by the underlying boulders and interstitial voids. As a consequence, big boulders or panels are solely stopped when they fit in gaps large enough to trap them or when they hit other main obstacles (e.g., other boulders) that make them to dissipate momentum. Hence, similar-sized boulders tend to come to rest and gather in the same zones of the talus-slope.

8.24.3.4 8.24.3.4.1

Stratified Slope Deposits within Talus Landforms (Including the Gre`zes Lite´es) Definition and geographical distribution

Stratified talus-slope deposits (stratified screes) are made up of sediment layers of varying grain size sorting (De Wolf, 1988) and, as such, they have received specific interest. They are characterized by a wide range of sedimentary facies, including all forms of stratified sediments from hardly apparent stratification (e.g., Hanvey and Lewis, 1991) to regular and wellbedded layers, first described in French lowlands as gre`zes lite´es (Figure 5(a); Guillien, 1951, 1964; Journaux, 1976). Stratified slope deposits are representative of former or current cold, mostly unglaciated environments and involve common periglacial slope processes. Stratification in talus sediments has been extensively reported during the last three decades from various mountain environments of mid- and low latitudes (Figure 5(b), Alps: Van Steijn et al., 1988; Bertran et al., 1993; Van Steijn and He´tu, 1997; Pappalardo, 1999; Andes: Francou, 1988, 1989; Van Steijn et al., 1995; Apennines: Coltorti et al., 1983; Coltorti and Dramis, 1995; Appalachian Mtns.: Gardner et al., 1991; Eaton et al., 2003; Drakensberg: Hanvey and Lewis, 1991; Himalaya – Tibet: Wasson, 1979; Gengnian et al., 1999; Pyrenees: Garcia-Ruiz et al., 2001; St. Elias Range: Harris and Prick, 2000; Tasmanian Range: Colhoun, 2002; Taurus: Nemec and Kazanci, 1999). Besides stratified slope deposits formed in close association with a talus morphology, similar forms (including the gre`zes lite´es) were also reported on gentler slopes in lowlands of Central and Western Europe (Figures 5(a) and 5(c), Watson, 1965; Boardman, 1978; Van Steijn et al., 1984; Pissart, 1995) and Canada (He´tu, 1991; He´tu et al., 1994), near the margins of former ice caps. However, stratified slope deposits seem less common in current high latitude periglacial environments (Francou, 1990), although such forms have been reported from the maritime Sub-Antarctic (Boelhouwers et al., 2003). Rock type appears to be a first-order control on the development of stratified slope deposits. Weathered and cracked rocks very sensitive to frost wedging and shattering, such as limestone (Figure 5(b)) or schist (Figure 5(c)), potentially produce large amounts of fine-grained material and appear to be most prone to favor such features (De Wolf, 1988).

8.24.3.4.2

399

Bedding and sedimentary structures

Layers within stratified slope deposits can be distinguished based on their fine fraction (diametero2 mm). End-members are represented on one hand by openwork, clast-supported beds (nearly) devoid of fine-grained particles, however by matrix-supported beds (Francou, 1990). Individual layers generally preserve their grain size properties greater than several meters so that bedding planes are mostly well-defined, especially at the top of some matrix-supported beds, which may form a true discontinuity surface (Figure 5(d), Francou, 1989; Bertran et al., 1992). This results both from matrix compaction by cryoaggregation and deposition of eluviated silt and clay forming a cohesive coating after drying. Therefore, the contact surface between the top of a matrix-rich bed and the base of the overlying openwork layer tends thus to be straight and tabular whereas the limit between the top of an openwork layer and the base of an overlying matrix-rich bed is more progressive and undulating (Figure 5(e); Francou, 1989; Bertran et al., 1992; Pappalardo, 1999). Detailed sedimentological studies have reported intralayer sorting processes that may be good indicators of the mechanisms involved (Figure 5(b), Francou, 1989; Gardner et al., 1991). Although the coarse fraction of matrix-supported beds appears to be smaller than that in openwork layers, an enrichment in fines commonly occurs in the upper part of matrix-rich beds (Francou, 1989; Bertran et al., 1992). As for clast-supported beds, they commonly display either normal or reverse grading (Figure 5(b), Francou, 1989; Bertran et al., 1992) or, in rarer cases, alternate grading (Francou, 1989). Moreover, one may encounter lateral variations or even inversion of grading within one individual bed (Bertran et al., 1992). Finally, successive layers may also be differentiated by clast orientation measurements (Francou, 1988). Though a trend toward isotropy occurs in some openwork layers, some upraised and imbricated clasts may produce a tiling effect (Bertran et al., 1992). In matrix-supported beds, the clasts’ A-axis generally parallels the slope angle, which is interpreted as a signature of frost creep (Francou, 1989).

8.24.3.4.3

Forming processes

Multiple explanations of the stratification of slope deposits have been proposed. Until the 1980s, they could be grouped in two classes. Firstly, some authors considered that an initially heterogeneous mass of debris might acquire a stratification either during its downslope movement or after deposition. Guillien (1951, 1964) proposed a cryonival hypothesis where layering occurred during transport by differential slopewash linked to the melting of semipermanent snow patches. By contrast, Journaux (1976) attributed the formation of openwork layers on top of soliflucted sheets of heterogeneous material to superficial slopewash that removed the fine fraction. In this case, water would come primarily from the seasonal thaw of the active layer and subsidiarily from snow melting and summer rainfall. However, both hypotheses failed to explain observed structures such as downslope and vertical graded bedding or cross bedding (Bertran et al., 1992). A second group of explanations assumes that clast-supported and matrix-supported beds are formed by distinct processes alternating through time. The supply of debris by rockfall (Tricart and Cailleux, 1967; Coltorti et al., 1983) or dry grain flow (Wasson, 1979;

400

Evolution of Slopes in a Cold Climate

Figure 5 Stratified slope deposits: general morphology and detailed sedimentary structures. (a) Typical gre`zes lite´es deposits with layer downbending at the top (Chaˆtillonais, France). (b) Late Glacial stratified slope deposits developed in limestone showing cementation of the matrixsupported layer and normal sorting (small clasts overlie larger ones) of the openwork layer (Southern Alps, France). (c) Closer view on stratified slope deposits developed in schist displaying bed alternance similar to that of gre`zes lite´es deposits (Famenne, Belgium). (d) Section in the frontal part of an active stone-banked solifluction sheet showing a discontinuity plan (sub-rectilinear ‘table’) at the top of matrix-rich layer (Eastern Andes, Bolivia). (e) Close section of gre`zes lite´es deposits displaying a straight and almost tabular boundary at the top of the matrix-supported layer (foreground in relief) and a strongly undulated contact at the top of the clast-supported layer (Charentes, France). (f) Succession of stone-banked solifluction lobes moving downslope on steep mountain side. Subsequent deposit stratification has been interpreted as the result of their progressive stacking up (Central Andes, Peru). (a) Reproduced from Journaux, A., 1976. Les gre`zes lite´es du Chaˆtillonais. Bulletin de l’Association franc- aise pour l’e´tude du Quaternaire 3–4, 123–138; (b), (d), and (f) Reproduced from Francou, B., 1989. La stratoge´ne`se dans les formations de pente soumises a` l’action du gel. Une nouvelle conception du proble`me. Bulletin de l’Association franc- aise pour l’e´tude du Quaternaire 26(4), 185–199; (c) Reproduced from Pissart, A., 1995. L’Ardenne sous le joug du froid. Le modele´ pe´riglaciaire du massif ardennais. In: Demoulin, A. (Ed.), L’Ardenne, Essai de Ge´ographie Physique. Universite´ de Lie`ge, Lie`ge, pp. 136–154, and (e) Reproduced from Bertran, P., Coutard J.-P., Francou, B., Ozouf, J.C., Texier, J.-P., 1992. Donne´es nouvelles sur l’origine du litage des gre`zes: implications pale´oclimatiques. Ge´ographie Physique et Quaternaire 46(1), 97–112.

Van Steijn et al., 1984) would form openwork layers, whereas solifluction lato sensu (Tricart and Cailleux, 1967) or rapid masswasting processes such as debris flows (Van Steijn et al., 1988) would be responsible for the presence of matrix-rich beds.

The main role of freeze/thaw cycles in the development of stratified slope deposits has thus long been recognized. Based on the analysis of active processes in Andean periglacial areas where freeze/thaw cycles are particularly effective (i.e., in

Evolution of Slopes in a Cold Climate

cryonival environments), Francou (1988, 1989, 1990) elaborated a more refined formation scheme for the stratified slope deposits. According to him, stratification primarily results from the stacking up of solifluction sheets (i.e., stonebanked solifluction lobes) moving downslope at rates of cm year 1 (Figure 5(f)). In his mind, the term solifluction includes not only gelifluction and frost creep but also pipkrakes (needle ice). Sorting occurs firstly during transport by frost heaving of the coarse elements combined with washing of the fines. The coarse clasts tend thus to be accumulated at the front of the flowing sheet, where a localized washing of the matrix takes place. The resulting clast-supported lower layer is later buried by the downslope movement of the fine-grained layer. This formation model has now been extended to explain the genesis of stratified slope deposits in other settings, such as the Alps (Bertran et al., 1993) or the relict gre`zes lite´es of Western Europe (Bertran, 1992). More recently, based on studies in mountain areas, variants on this mechanism have been proposed. Successive occurrences of dominant rockfalls, cohesive debris flow, and sheetwash processes, accompanied by minor grain flow and slush flow (wet snow flow), have been highlighted in the Taurus Mountains (Nemec and Kazanci, 1999). In the St. Elias Range of the Yukon Territory, an annual cycle based on snowfree slopes during warm and dry springs has been proposed (Harris and Prick, 2000). In this context, dry grain flow is particularly active and involves a preferential downslope transport of the coarsest particles toward the footslope. During the following summer, the matrix-rich deposits remain upslope and in turn are mobilized through rainfall-induced debris flows and are deposited above the coarser layer, which is then stabilized when the overlying matrix-rich layer dries and hardens. This model differs from the stone-banked lobe model of Francou (1990) in that the successive beds are alternatively formed by dry grain flow and thin debris flow.

8.24.3.5 8.24.3.5.1

Talus-Slope Evolution Rates of debris supply/rockwall retreat and paraglacial activity

Using either marked surfaces (e.g., Rapp, 1960 ; Luckman, 1988) or lichenometry (e.g., Luckman and Fiske, 1997; McCarroll et al., 1998), short-term rate estimates of debris accumulation on talus slopes were produced in an attempt to determine rates of rockwall retreat, and therefore rates of landscape change. Depending mainly on the frequency of snow avalanches, mean accumulation rates of up to B5m ka 1 have been inferred from a 13-year survey in Canada (Luckman, 1988). Based on lichen size measurements (Rhizocarpon) covering the last 400 years, the estimated rate of late Holocene rockfalls in Norway would result in covering B4% of the talus surface each 25 years by new boulders (McCarroll et al., 1998). However, the scaling up of results obtained from measurements made at very small spatial and temporal scales is problematic. Indeed, only a small portion of a talus is generally sampled, and the time range of lichenometry does not exceed the last few centuries. During the last decade, several authors used various geophysical techniques, such as 2D-electric resistivity, seismic

401

refraction or ground penetrating radar (GPR), in alpine environments, in order to investigate talus internal structure and especially to infer sediment thicknesses (Sass and Wollny, 2001; Hoffmann and Schrott, 2002; Sass, 2006, 2007; Otto and Sass, 2006; Sass and Krautblatter, 2007; Pieracci et al., 2008). Such studies, particularly where seismic surveys and GPR were used in combination (Otto and Sass, 2006; Sass, 2007), were able to provide reliable estimates of the sediment thickness on talus slopes, from which long-term rates of rockwall retreat have been inferred. In the Bavarian Alps, backwearing rates were comprised between 0.15 and 0.73 m ka 1 during the last glacial period (Sass and Wollny, 2001) and Holocene rate estimates vary from 0.1–0.3 m ka 1 (Sass and Wollny, 2001) to an average value of 0.5 m ka 1 (Hoffmann and Schrott, 2002). Similar rates of B0.7 and 0.5–0.8 m ka 1 also occur, respectively, in the Swiss and Austrian Alps (Otto and Sass, 2006; Sass, 2006). These estimations are consistent with the rates inferred for the last two millennia in the Arctic region of Svalbard (Andre´, 1997) where they vary from 0.1 m ka 1 (under frost shattering) up to 1 m ka 1 (where postglacial stress relaxation dominates). Obanawa and Matsukura (2006) recently developed a model of rockwall retreat (including both declining and parallel retreat) calibrated by field experiments of talus development. In summary, whatever the setting, the estimated rates of rockwall retreat and debris production are invariably rather low. Furthermore, in formerly glaciated areas, many talus slopes are at present inactive relict features (e.g., vegetated talus). Admitting a limited efficiency of the periglacial processes to deliver high amounts of sediments to talus slopes (e.g., through intense freeze/thaw cycles in cold postglacial climates), this has led some authors to consider a paraglacial origin of the talus slopes, with brief episodes of high rates of debris production (see Luckman and Fiske, 1997; Ballantyne, 2002). In their opinion, a rapid and massive, but brief accumulation took place immediately after the glacial recession, when stress release and re-exposure of the rockwalls to weathering combined to produce conditions highly favorable to cliff failure and mass wasting. This paraglacial activity has been recently incorporated in a four-step evolution model for relict talus slopes in upland Britain (Figure 6, Hinchliffe and Ballantyne, 2009). During initial paraglacial stress release, a rapid and intense accumulation episode of talus debris by rockfall and granular disintegration occurs, also facilitated by enhanced freeze–thaw activity at that time (‘paraglacial’ talus, Figure 6(a)). During the second stage, delivery of debris through rockfall almost ceases as the rockwall is now in equilibrium with postglacial conditions (mature and ‘relict’ talus, Figure 6(b)). During the third stage, gully initiation and localized failures take place in the uppermost part of the talus accumulation. Channeled in these incipient pathways, consequent debris flow and slopewash processes rework upslope materials (Figure 6(c)). The final stage is associated with gully development and formation of downslope debris cones as reworked sediments are redeposited and accumulated at the slope-foot through debris flows, sediment failures and runoff occurring within gullies (Figure 6(d)). Provided that talus slopes are located within formerly glaciated areas, this four-step evolution scheme may presumably be extended to other regions.

402

Evolution of Slopes in a Cold Climate

Rockwall

Rockwall

Retreating glacier

Mature talus

Paraglacial talus

(a)

(b)

Rockwall

Rockwall

Debris-flow and slopewash deposits

Gullies Debris cones

(c)

(d)

Figure 6 Four-step evolution scheme of relict talus slopes in upland Britain: (a) Initial paraglacial stress release inducing enhanced rockfall activity results in rapid talus formation (‘paraglacial’ talus), granular weathering and increased freeze–thaw activity playing a secondary role. (b) Near-cessation of rockfall supply as a consequence of cliff equilibration to postglacial conditions, resulting in ‘relict’, soil-covered (vegetation) talus slope. (c) Debris-flow and slopewash processes are channeled in gullies and localized failures occurring in the uppermost part of the talus accumulation, inducing a reworking of upslope materials. (d) Gullies development through debris flows, sediment failures, and runoff occurring within gullies is associated to the formation of downslope debris cone as a result of slope-foot accumulation of reworked materials. Reproduced from Hinchcliffe S., Ballantyne, C.K., 2009. Talus structure and evolution on sandstone mountains in NW Scotland, The Holocene 19(3), 477–486, with permission from SAGE.

8.24.3.5.2

Paleoenvironmental implications of talus slopes and stratified slope deposits

The paleoclimatic significance of talus slopes, especially of inactive forms, has raised much interest in recent years (e.g., Hinchliffe, 1999; Curry and Black, 2003). Indeed, the role of climatic variations in the talus development has been recently underlined under various settings. Periods of alternating rockfall activity and cliff stability have been inferred from layers of openwork coarse debris alternating with paleosoils, peat or finer grade materials (see below). Rockfall events were correlated to cool/cold climate, and the stabilization of the talus was supposed to occur during wetter and warmer conditions. For instance, paleosoils interlayered within colluvial sequences of northern Quebec have been attributed to alternating process regimes linked to the Little Ice Age (He´tu and Gray, 2000). A similar conclusion has been reached in the eastern European Alps (Sass and Krautblatter, 2007). During episodes of cold conditions and enhanced freeze/thaw activity such as the Little Ice Age, the production of coarse debris by rockfall would prevail, whereas finer debris (o2 cm) taken

from stocks of fine-grained sediments in intermediate places would be delivered to the talus by the more frequent rainstorms under warmer climates. In this respect, lichenometryderived rates of debris supply by rockfall during the eighteenth century, the coldest phase of the Little Ice Age, are estimated up to five times higher than the mean late Holocene rate (McCarroll et al., 1998). These results are consistent with talus formation under periglacial conditions during the Holocene. As for the stratified slope deposits, it is observed that the stratification is better developed when the annual rate of daily freeze/thaw cycles is high (4200 days year 1). This emerges from experimental studies (Guillien and Lautridou, 1970) and from a comparison between Andean and Alpine settings, where daily freeze/thaw occurs, respectively, all the year round and only during spring and autumn (Francou, 1989, 1990; Bertran et al., 1993). The layering is also more apparent for freezing depths not exceeding a few decimeters, i.e., in shallow seasonally frozen ground. This has led Bertran et al. (1992) to assign the formation of the French gre`zes lite´es to periods of moderate cooling within the glacial cycles. According to them, the glacial

Evolution of Slopes in a Cold Climate

maximum would be recorded in the thick uppermost cryoturbated layer whose formation probably required a deep seasonally frozen ground or even a perennially permafrost. Palynological analyses also show that the gre`zes lite´es formed under a low-growing steppe with scarce pine trees (Bastin and Guillien, 1971). Similarly, the Alpine stratified slope deposits are related to midlatitude periglacial environments with a subMediterranean or a suboceanic flavor (Bertran et al., 1993). By contrast, 14C dating, U/Th dating of crypto-speleothems and pollen spectra indicate that, in the Pyrenees, stratified slope deposits developed mainly during the last glacial maximum (LGM) and the cold episodes of the Late Glacial, in particular the Older Dryas (Garcia-Ruiz et al., 2001). However, there are restrictions to the validity of such paleoclimatic inferences for talus development (Luckman, 2007):

• • •

Age constraints on talus sediments are characteristically limited, relying essentially on 14C dating. Deposition events on talus slopes are episodic and discontinuous in time, implying that individual stratigraphic units can be very thin. A thicker sediment accumulation may also result from one single and nearly instantaneous process (e.g., rockfall, debris flow, snow avalanche) rather than from a long period of time including climate variations.

As it is very difficult to ascertain whether individual deposits correspond to a single or several events, and as the magnitude/ frequency spectra of the sediment-delivering processes are rather badly constrained, the paleoenvironmental interpretation of talus slopes and their deposits remains thus a highly delicate task needing much further careful observation and analysis.

8.24.4 8.24.4.1

Blockfields Definition and Geographic Distribution

Generally defined as a veneer of coarse regolith (Goodfellow, 2007), blockfields correspond to large, sheetlike expanses of weathered blocks and debris covering the bedrock in lowgradient areas (Figures 7(a) and 7(b)). Several synonyms of the term ’blockfield’ are found in the literature: felsenmeer (e.g., Sudgen and Watts, 1977; Marquette et al., 2004), mountain-top detritus (Ives, 1958; Ballantyne, 1998), boulder field (Kleman and Borgstro¨m, 1990) or even blockmeer and stone field. This morphology occurs predominantly in (sub)arctic regions such as Canada, Finland, Norway (including Svalbard) and Sweden (e.g., Dahl, 1966; Sudgen and Watts, 1977; Nesje, 1989; Kleman and Borgstro¨m, 1990; Dyke, 1993; Rea et al., 1996a, b; Dredge, 2000; Fjellanger et al., 2006) or in midlatitude regions such as Scotland (e.g., Mellor and Wilson, 1989; Ballantyne, 1984, 1998). However, some examples have also been reported from low latitude areas (e.g., Lesotho, Namibia, and Tasmania, see Caine, 1983; Boelhouwers, 2004). Topographically, blockfields are preferentially located on upland plateaus or mountain summits (Figure 7(a); Rea et al., 1996a; Whalley et al., 2004), hence the name mountain-top detritus (see Ballantyne, 1998), but some also occur at lower elevations, particularly in Scandinavia (Piirola, 1969; Ha¨ttestrand, 1994; Ha¨ttestrand and Stroeven, 2002).

403

Although they are always composed of rock types resistant to weathering, a practical distinction can be operated between autochthonous and allochthonous blockfields (e.g., Boelhouwers, 2004; Goodfellow, 2007). An autochthonous blockfield is formed as an in-situ residue of the underlying weathered bedrock (Figures 7(c) and 7(d), see Ballantyne, 1998) and is mostly restricted to tectonically stable environments (see Goddard, 1989). The allochthonous type corresponds to blockfields formed from either the downslope motion of weathered debris, or frost heaving of till clasts (e.g., Potter and Moss, 1968; Van Steijn et al., 2002). There are still many controversial issues regarding origin and age of the blockfields, especially in high to midlatitudes, where both parameters are indeed interrelated so that an age determination may help unraveling the origin, and vice versa (Roaldset et al., 1982; Goddard, 1989; Rea et al., 1996a, b; Ballantyne, 1998, 2010).

8.24.4.2

Form Characteristics

Geomorphologically, a blockfield corresponds to a surface covered with angular blocks that either rest directly on the bedrock, or overlie (and are more or less embedded in) a layer of finer material, ranging from smaller clasts to sand and clay (Rea, 2007). As for the block cover, it can be made of either similar-sized (well sorted) boulders embedded in a fine matrix, or an openwork of debris of variable size, from cobble to boulder (Figure 7(a)). In this respect, the presence of largesized boulders tends to favor the formation of the second type of blockfield. One will also note that some blockfields occur as extensive blockspreads without any conspicuous source area (Figure 7(a)) whereas others are clearly associated with tors (Figure 7(b), see below, e.g., Sudgen and Watts, 1977; Kleman and Borgstro¨m, 1990; Dyke, 1993). Slope angle obviously appears as a first-order control on the formation and the preservation of blockfields (Rea, 2007). Indeed, too high an angle inhibits the accumulation of blocks because it increases the shear stress and, consequently, the rate of downslope transport of the blocks. An upper threshold angle of 251 has been reported for the formation of blockfields, beyond which talus slopes are created (Dahl, 1966). Blockfields preferentially develop on surfaces sloping less than 101, in particular on plateau areas (Rea et al., 1996b). It has also been shown that openwork blockfields tend to predominate at higher slope angles, as a consequence of the more active removal of fines on steeper slopes (Rea et al., 1996a, b). The average thickness of blockfields is B1m, but thicknesses greater than 1.5 m have also been reported (Dahl, 1966; Ballantyne, 1998). The blockfield thickness depends on many parameters, notably the slope angle, the rock type, the age, and the erosional context (Rea, 2007). Initially based on forms studied in Scotland, a tripartite classification of blockfield profiles proposed by Ballantyne (1984, 1998) can actually be applied in other settings:

• •

Openwork cover of boulders devoid of fine-grained sediments at the surface (‘clast-supported’, Figure 7(e)). Rock debris of varying size embedded within a cohesionless matrix of coarse sand, which continues downward as a sand layer (‘matrix-supported’, Figure 7(f)).

404

Evolution of Slopes in a Cold Climate

Figure 7 Autochthonous blockfields: general morphology, evidences for their in-situ origin and cross-sections: (a) Typical plateau-top blockfield (Øksfjordjøkelen, Norway). The arrow indicates a person for scale. (b) Blockfield-mantled slope running down a tor (Norway). The arrow indicates a person for scale. (c) Line-disposed pegmatite blocks at the surface within a gneissic blockfield, underlining the presence of an underlying pegmatite vein (Jotunheimen, Norway). (d) Sharp modification in blockfield lithology at a geological contact (Torlesse Range, New Zealand). This picture and the previous one undoubtedly attest of an in-situ weathering of the underlying bedrock and an in-situ origin for these blockfields. (e) Pit dug through an openwork, clast-supported quartzitic blockfield (An Teallach, Scotland). A slight rounding of the block edges exposed at the surface should be noticed. (f) Pit dug through a matrix-supported blockfield formed on sandstone bedrock (An Teallach, Scotland). The contrast between rounded edges of surficial blocks and dominant angular clasts extracted from the pit should be noticed. (a) Reproduced from Rea, B.R., Whalley, W.B., Rainey, M.M., Gordon, J.E., 1996a. Blockfields, old or new? Evidence and implications from some plateaus in northern Norway. Geomorphology 15, 109–121; (b) Reproduced from Rea, B.R., Whalley, W.B., Porter, E.M., 1996b. Rock weathering and the formation of summit blockfields slopes in Norway: examples and implications. In: Anderson, E., Brooks (Eds.), Advances in Hillslope Processes. John Wiley and Sons Ltd., Chichester, pp. 1257–1275; and (c)–(f): Ballantyne, C.K., 2010. A general model of autochthonous block field evolution. Permafrost and Periglacial Processes 21, 289–300.

•

Blocks embedded within a frost-sensitive silt/clay matrix that commonly displays evidence of vertical frost sorting, the coarser debris being accumulated nearby the surface (‘matrix-supported’).

Although these profile types may be helpful to characterize the internal structure of blockfields, they are of course only end-members of a continuum, and every intermediate type may actually occur.

Cold-climate mechanical weathering is commonly thought to produce angular clasts (see Chapter 4.15). However, some authors are not so firm on this point (see Hall et al., 2002; Boelhouwers, 2004). In this respect, rounded elements are likely to be shaped by granular disintegration in cold climates (Figures 7(e) and 7(f), Ballantyne, 1998). At the grain scale, one associates also general angularity with mechanical weathering and roundness with chemical etching (Rea, 2007). In brief, element angularity reliably betrays a mechanical

Evolution of Slopes in a Cold Climate

production but cannot be unequivocally linked to specific (warm or cold) climatic conditions. Finally, a difference in roundness between exposed and buried clasts may be a very relevant indicator of blockfield inactivity (Figure 7(f)), notably when the latter is produced by frost heave processes. Indeed, the contrast between well-rounded exposed clasts resulting from prolonged granular disintegration at the surface, and more angular subsurface debris points to the profile stability of a relict blockfield (Ballantyne, 1998).

8.24.4.3

Block Accumulation Processes

Frost shattering of the bedrock under severe periglacial conditions has been traditionally considered as the main block production process for autochthonous blockfields (e.g., Ives, 1958; Dahl, 1966; Kleman and Borgstro¨m, 1990; French, 2000). Frost shattering of intact porous sedimentary rocks produces brecciated material whereas rock detachment by frost wedging operating in preexisting joints and cracks delivers surface clasts (Walder and Hallet, 1985, 1986; Ballantyne, 1998; Matsuoka, 2001, 2008; Boelhouwers, 2004; Murton et al., 2006). This is especially efficient for blockfields located above permafrost, provided that large amounts of water are available at the weathering front during winter freezeback. In this case, the blockfield base is saturated at the start of freezeback as a consequence of hampered drainage (Dahl, 1966), thus creating optimal conditions for frost crack generation or enlargement and, consequently, rock detachment (Goodfellow et al., 2008). Frost wedging activity near the weathering front is also enhanced by the abrupt temperature drop typical of the thermal regime of openworkdebris blockfields (see Harris and Pedersen, 1998; Gorbunov et al., 2004; Juliussen and Humlum, 2008) and by upwards freezing from the permafrost table in cold permafrost regions (Ballantyne, 2010). Yet, the need for a rigorous periglacial climate as a condition for effective freeze–thaw weathering and significant block production is a matter of debate. On one hand, some authors suggested that the number of cycles around 0 1C appears to be more important than the duration and/or the intensity of freezing (Wiman, 1963; Potts, 1970; Whalley and McGreevy, 1987). On the other hand, blockfield formation on resistant and well-jointed rock types (e.g., microgranite, quartzite) has been reported to have occurred under severe periglacial conditions, for instance directly following deglaciation in northwest Scotland (Ballantyne, 1998; Ballantyne et al., 1998). Hall et al. (2002) have also challenged the commonly accepted idea that blockfield production predominantly results from freeze–thaw weathering. In this respect, the existence of blockfields in nonperiglacial arid settings, like the desert of Namibia (see Van Steijn et al., 2002; Boelhouwers, 2004), undoubtedly attests that other processes may also produce autochthonous features, which are not therefore necessarily indicative of periglacial conditions (French, 2000). Moreover, physical evidence of the efficiency of frost shattering of highstrength, low-porosity crystalline rocks such as quartzite, granite, granulite, is still lacking (Lautridou and Seppa¨la¨, 1986; Matsuoka, 2001; Andre´, 2004). Alternative mechanisms of block production might be thermal shock and thermal

405

stress (Hall, 1999; Hall and Andre´, 2003; Sumner et al., 2004; Boelhouwers, 2004). Induced by the large range of diurnal and seasonal temperature variations at the rock surface occurring in various settings, from (sub)polar and alpine to desert environments, these mechanisms are however no more definitely convincing as the capability of thermal stress to overcome the rock tensile strength is still questioned, and possible effects of thermal shocks are likely to be superficial (Boelhouwers, 2004).

8.24.4.4

Origin, Age, and Paleoenvironmental Significance of Autochthonous Blockfields

The origin and evolution of autochthonous blockfields, and the related question of their age, have always been, and still are a contentious issue. Two main theoretical models and chronological frameworks have been proposed (see Ballantyne, 2010):

•

•

A Quaternary periglacial model that attributes the formation of autochthonous blockfields to frost-induced weathering and sorting processes (see above) during cold episodes of the Quaternary (e.g., Kleman and Borgstro¨m, 1990; Dredge, 1992; Ballantyne and Harris, 1994; Ballantyne et al., 1998; Goodfellow et al., 2008). A pre-Pleistocene (Neogene) model that assigns the production of rock debris to pre-Quaternary chemical weathering processes (‘inherited’ materials, see Chapter 7.29) and restricts the frost action to subsequent minor reworking and sorting of the preexisting regolith (e.g., Dahl, 1987; Nesje, 1989; Andre´, 2003, 2004; Marquette et al., 2004; Paasche et al., 2006).

8.24.4.4.1

The Quaternary periglacial model – short to mid-term formation

Adherent to the idea of block production by frost-induced weathering, most earlier works (e.g., Lundqvist, 1944; Rapp and Rudberg, 1960; Dahl, 1966) also assumed a strong erosive action of ice sheets of the LGM that compelled them to place the formation of autochthonous blockfields during the Late Glacial and Postglacial (Holocene) periods. More recently, other researchers (e.g., Goodfellow, 2007) suggested however that blockfields might also be Pleistocene features having survived one or more glaciations (Figure 8). One has now recognized that the erosive power of coldbased glaciers is very weak, so that they can indeed rather well preserve preexisting landforms (Figure 8). However, field studies obviously point to rapid and recent blockfield developments, e.g., in northwest Scotland where they formed under severe periglacial conditions in the time between the LGM and the Younger Dryas (see above, Ballantyne, 1998; Ballantyne et al., 1998). Another evidence of recent blockfield activity is provided by autochthonous forms currently developing in the eastern Canadian Arctic (Dredge, 1992, 2000). Under permafrost conditions, they appear to result from either mechanical weathering on frost-susceptible carbonates or chemical weathering on frost-resistant granites and gneisses (see Chapter 4.14), but they are probably more an exception than representing a general trend (Ballantyne, 2010). Indeed, current activity of autochthonous blockfields and

406

Evolution of Slopes in a Cold Climate

Meiklejohn, 2004; Fjellanger et al., 2006; Paasche et al., 2006). Such an antiquity of autochthonous blockfield elements is supported by the following observations (Ballantyne, 2010):

Figure 8 Glacially transported block (erratic) overlying a plateau blockfield on Grytøya (NW Norway). This situation strongly points to the former presence of a cold-based glacier, whose erosive action was very limited, and therefore preserved the preexisting landform. This suggests that the underlying blockfield has survived one or more glaciations. Reproduced from Ballantyne, C.K., 2010. A general model of autochthonous blockfield evolution. Permafrost and Periglacial Processes 21, 289–300.

development of new forms are not known elsewhere (Boelhouwers, 2004), even in a priori favorable settings, such as the sub-Antarctic islands, where frost-induced weathering or thermal effects should be able to produce rock debris (Sumner and Meiklejohn, 2004). It has thus been proposed that, in most cases, the blockfields formed during the Pleistocene through frost-induced weathering of the bedrock, and outlasted at least one glacial/ interglacial cycle (e.g., Lautridou and Seppa¨la¨, 1986; Rea et al., 1996b; Sumner and Meiklejohn, 2004; Whalley et al., 2004). Some of these authors suggested that the blockfields developed even over several glacial cycles, being preserved beneath cold-based glaciers during the cold periods (Figure 8, e.g., Nesje et al., 1988; Kleman and Borgstro¨m, 1990; Rea et al., 1996a; Ha¨ttestrand and Stroeven, 2002). This would then lend some paleoenvironmental meaning to the blockfields situated in ice-covered areas, which might be indicative of the ice-sheet geometry and the ice dynamics (Rea, 2007). Other studies pointed out the role of nunataks, notably showing that frost shattering of nunatak rocks during cold episodes may have resulted in blockfield formation (e.g., Ballantyne and Harris, 1994). Furthermore, intact ‘nunatak blockfields’ may help determining the trimline of a subsequent glaciation (see Dahl, 1987), an important paleoenvironmental information that allows an accurate estimate of the ice thickness (Nesje et al., 1988; Nesje, 1989; McCarroll et al. (1995); Ballantyne, 1998).

8.24.4.4.2

The pre-Pleistocene (Neogene) model – long-term formation

Many authors do not believe that the blocks of blockfields were produced by frost shattering, and suggest instead that they represent remnants of pre-Pleistocene regoliths subsequently reworked by Quaternary frost action (e.g., Whalley et al., 1997, 2004; Andre´, 2003; Marquette et al., 2004; Sumner and

1. Frost wedging is hardly present on glacially smoothed rock outcrops in some cold settings (Andre´, 2002, 2003), suggesting that rock debris are unlikely to result from postglacial frost-induced weathering and that blockfields are also unlikely to have developed during former interglacials. 2. Many blockfields are located on elevated plateaus commonly interpreted as uplifted Tertiary planation surfaces, especially in Norway (Nesje, 1989; Whalley et al., 2004; Paasche et al., 2006), and are spatially related to tors and inselbergs of presumed pre-Quaternary age (e.g., Andre´, 2003). 3. Although high silt and clay proportions (from 15% to 80%) characterize the matrix of blockfields (e.g., Paasche et al., 2006), some laboratory experiments have shown that frost wedging is not able to deliver large amounts of fine material (Rea et al., 1996b; Whalley et al., 2004). In this respect, pockets of clay-rich saprolite stemming from in situ chemical weathering of the bedrock and locally associated with blockfields may be considered as a robust marker of preglacial regolith, thus indicating a Neogene, or even older origin of the blocks (e.g., Hall et al., 1989; Paasche et al., 2006). 4. The presence of kaolinite and gibbsite within blockfield soils has also been recurrently evoked to confirm a Tertiary inheritance. Indeed, these clay minerals most probably result from weathering under conditions wetter and warmer than those that prevailed during the Quaternary in the regions where the blockfields are present, and they support the claim of a block production taking place under warmer preglacial climates (Roaldset et al., 1982; Whalley et al., 2004; Marquette et al., 2004; Paasche et al., 2006). Though the above arguments collectively tend to prove the Neogene origin of blocks through intense chemical weathering, they nevertheless need to be discussed (Ballantyne, 2010). For instance, some authors argue that the preservation of preglacial upland plateaus is contradicted by the lowering of surrounding areas notably suggested by the emergence of tors (Ballantyne, 2010). The issues of fine production by frost weathering and efficiency of chemical weathering during Quaternary periglacial conditions are also still debated. Indeed, other laboratory experiments reached the conclusion that significant amounts of silt may be produced by frost-induced weathering (Lautridou and Seppa¨la¨, 1986; Smith et al., 2002), and some studies showed that chemical weathering under periglacial conditions in arctic or alpine settings can potentially produce non-negligible amounts of silt and clay (e.g., Smith et al., 2002; Dixon and Thorn, 2005). Finally, although high concentrations of kaolinite and gibbsite reflect long-lasting weathering under warm and wet conditions (e.g., Whalley et al., 2004), Goodfellow (2007) recently claimed that the presence of these minerals is not necessarily linked to specific climatic conditions. Kaolinite, he argued, might also result from an early diagenetic mineralization phase (Fjellanger and Nystuen, 2007), and gibbsite, which can occur in the early stages of weathering, might have developed during Pleistocene interglacials (e.g., Hall et al., 1989; Bouchard et al., 1995).

Evolution of Slopes in a Cold Climate 8.24.4.4.3

•

•

A unified scheme of blockfield formation: Neogene inheritance and Quaternary development

Surface lowering and dating of tor emergence by cosmogenic nuclides: A recent study has suggested that the two, Quaternary periglacial and Neogene, models of blockfield formation, hitherto considered as incompatible, can be reconciled and unified in one single theoretical evolution framework (Ballantyne, 2010). This study assumes that the Tertiary planation surfaces on which autochthonous blockfields occur, continued to evolve during the Quaternary by surface lowering (Goodfellow, 2007). The key point here is the dating of the emergence of tors above blockfield-mantled plateaus, induced by the contrast in erosion rate between the tor and the adjacent regolith as a consequence of the higher resistance of the tor bedrock due, e.g., to a lower joint density (Andre´, 2004; Hall and Phillips, 2006). Dating of plateau blockfields and tors by terrestrial cosmogenic nuclides recently demonstrated that the plateaus covered by blockfields evolved significantly during the Quaternary (e.g., Marquette et al., 2004; Phillips et al., 2006; Darmody et al., 2008), contradicting the previous picture of surfaces unchanged since the Neogene (Andre´, 2003, 2004; Whalley et al., 2004). Despite signs of glacial alteration (e.g., Andre´, 2004), tors on these plateaus were largely considered as preglacial landforms preserved under Quaternary coldbased glaciers and ice sheets (e.g., Bierman et al., 1999; Stroeven et al., 2002). Coupled 10Be and 26Al measurements revealed the complex history of alternating episodes of surface exposure and shielding beneath ice sheets of these tors, notably showing that the oldest sampled tor surfaces were of early Pleistocene age and that tor emergence in many instances happened only in the middle Pleistocene, between B0.2 and B1 Myr (Phillips et al., 2006; Darmody et al., 2008). As pointed out by Ballantyne (2010), Quaternary lowering rates of the surrounding surfaces may thus be derived from these data and, though the tor’s material is of Tertiary origin, its uncovering and the formation of blockfields are Quaternary in age. For example, the 10Be concentration profile measured for a glacially unaltered tor in Scotland allowed Phillips et al. (2006) to calculate a 35 m lowering of the surrounding surface during the last 1 Ma. Temporal model of autochthonous blockfield formation: Ballantyne (2010) recently proposed a synthesizing model taking into account the new age data and the nonubiquitous presence of Neogene material within the blockfields (Figure 9). He assumed that Quaternary blockfields have developed from a pre-Quaternary regolithic mantle, except when the latter had previously been completely removed, e.g., by glacial erosion. The presence or absence of Tertiary evidence in modern blockfields depends thus not only on the initial thickness of the preQuaternary regolith but also on the rate at which it was removed, in parallel with surface lowering.

Figure 9 provides a time-integrated illustration of Ballantyne’s model. Assuming an original regolith thickness of several meters over a (sub)horizontal Tertiary surface and some

407

surface lowering, the modeled blockfield evolution can be divided into four stages: 1. Pre-Quaternary chemical weathering was first active and produced residual corestones embedded in the saprolite, in particular in its deeper part close to the weathering front (Figure 9(a)). In more weathering-resistant places, these corestones were more numerous and, everywhere they survived across the whole weathering profile and finally emerged at the surface, tors started to form sooner or later. 2. With the climate cooling starting at the onset of the Quaternary, the weathering front ceased to deepen and the saprolitic mantle was gradually removed. Incipient blockfields (‘proto-blockfields’) formed from the concentration of residual corestones either overlying some remaining saprolite layer or embedded within a saprolitic matrix (Figure 9(b)). From this time onwards, the armoring effect of the developing block cover probably reduced the rate of surface lowering (Boelhouwers, 2004). 3. In a transitional phase, continued stripping and thinning of the weathering mantle allowed frost activity to reach the top of the underlying bedrock during cold episodes. As a consequence, frost shattering progressively became the predominant process delivering debris to the blockfield. Although the blockfield retained a part of inherited corestones, the latter became progressively diluted in an increasing proportion of clasts and fine-grained sediments delivered by mechanical weathering processes under periglacial conditions (Figure 9(c)). Other processes, such as vertical frost sorting inducing clast accumulation at the surface, or granular disintegration of exposed blocks, may also have been active at this moment. As the depth of annual freeze–thaw cycles (i.e., of the active layer) now controls the depth of the weathering front, the rates of rock disintegration and surface lowering balance through a feedback mechanism that prevents further thickening of the blockfield. 4. In a late stage of blockfield maturity, every evidence of Tertiary inherited material may have been removed, so that a ‘strictly periglacial’ blockfield is observed. Though modest traces of chemical weathering are perhaps still visible, the blockfield would henceforth be essentially constituted of debris produced by mechanical weathering, within or over fine-grained sediments delivered by granular disintegration (Figure 9(d)). In summary, this time-integrated model of Ballantyne (2010) hypothesizes that all blockfields were originally derived from Tertiary regolithic material, and that many of them are still in the transitional stage (3), whereas others already reached the maturity stage (4).

8.24.5 8.24.5.1

Block Streams Definition and Geographic Distribution

Block streams may be seen as a particular type of blockfield characterized by downslope linear, more or less narrow ribbons of blocks (Figure 10, Van Steijn et al., 2002; Wilson, 2007), bearing witness to a longitudinal, downslope transport component. Confusion sometimes arose in the literature because some block streams are part of larger blockfields (see e.g., Clark,

408

Evolution of Slopes in a Cold Climate

Incipient tor 0

Metres

Saprolite 5 10

WF Unweathered bedrock

Corestones

(a) 15

Tor

0

Metres

Residual saprolite 5 WF 10

Unweathered bedrock

Corestones

(b) 15

Tor

Metres

0

Blockfield of corestones and frost-weathered regolith

5 10

Unweathered bedrock

WF

(c) 15

0

Metres

Tor 5 10

Blockfield of frost-weathered regolith Unweathered bedrock

WF

(d) 15

Figure 9 Time-integrated model of blockfield evolution (and parallel tor emergence). WF corresponds to weathering front: (a) Pre-Quaternary chemical weathering produces residual corestones embedded in the saprolite whereas an incipient tors starts to form. (b) Climate cooling at the onset of the Quaternary results in cessation of weathering front to deepen and progressive removal of the saprolitic mantle. Incipient blockfield (‘proto-blockfield’) starts to form. (c) Continued stripping of the weathering mantle allowed frost activity to reach the top of the underlying bedrock during cold episodes. Consequently, frost shattering progressively became the predominant process delivering debris to the blockfield, although it retains a part of inherited corestones. (d) The (almost) complete removal of any evidence of Tertiary inherited material produces a ‘strictly periglacial’ blockfield (stage of maturity). Reproduced from Ballantyne, C.K., 2010. A general model of autochthonous blockfield evolution. Permafrost and Periglacial Processes 21, 289–300.

1972) and, more generally, various types of block accumulation (blockfields, block streams, ‘block slopes’, rock glaciers) can occur in close spatial association, even grading into one another or merging (Wilson, 2007). Although periglacial block streams may also occur on hill sides (Figures 10(a) and 10(b)), many of them occupy valley bottoms, where they are clearly distinct from river deposits (Figures 10(c) and 10(d), Boelhouwers et al., 2002; Colhoun, 2002; Wilson et al., 2008). Though less frequent than blockfields, block streams have been described worldwide in current or former periglacial areas, including Australia and Tasmania (Caine and Jennings, 1968; Colhoun, 2002; Barrows et al., 2004), China (Harris et al., 1998), Italy (Ginesu, 1990), Russia where they are

termed kurums (Romanovskii and Tyurin, 1986), South Africa (Boelhouwers, 1999; Boelhouwers et al., 2002), United Kingdom (Clark, 1994), and USA (Potter and Moss, 1968; Cremeens et al., 2005). Large-size block streams known as stone runs have also been extensively studied in the Falkland Islands (Figure 10(a) and 10(d)–10(g), Clark, 1972; Clapperton, 1975; Aldiss and Edwards, 1999; Andre´ et al., 2008; Hansom et al., 2008; Wilson et al., 2008).

8.24.5.2

Form Characteristics

The size of block streams may be fairly variable. Small features exhibit widths varying from a few meters to a few tens of meters

Evolution of Slopes in a Cold Climate

409

Figure 10 Block streams: topographic situations, planform patterns, and cross-section: (a) Large hillside block stream (stone run) in the Falkand Islands exceeding 3 km in length. (b) Sinuous hillside block stream (Morgan Run) at Mount Storm in West Virginia (USA). (c) Typical straight valley-bottom block stream in Jasper National Park (Alberta, Canada). (d) Valley-bottom stone run (main trunk) joined by tributary block streams running down hillslopes and forming a dendritic pattern in the Falkand Islands. The arrow indicates downstream direction. (e) Multichannel block stream displaying a braided pattern in the Falkand Islands. (f) Parallel block stripes running downhill very close to each other, this stripped pattern being specific to the Falkland Islands. (g) Cross-section in a Falkland Islands block stream displaying a reverse grading and a lack of fine-grained matrix between the blocks at their surface. The survey pole is 1 m in height. (a) Reproduced from Hansom, J.D., Evans, D.J.A., Sanderson, D.C.W., Bingham, R.G., Bentley, M.J., 2008. Constraining the age and formation of stone runs in the Falkland Islands using Optically Stimulated Luminescence. Geomorphology 94, 117–130; (b) Reproduced from Cremeens, D.L., Darmody, R.G., George, S.E., 2005. Upper slope landforms and age of bedrock exposures in the St. Francois Mountains, Missouri: a comparison to relict periglacial features in the Appalachian Plateau of West Virginia, Geomorphology 70, 71–84; (c) Reproduced from Van Steijn, H., Boelhouwers, J., Harris, S., He´tu, B., 2002. Recent research on the nature, origin and climatic relations of blocky and stratified slope deposits. Progress in Physical Geography 26(4), 551–575; (d) Reproduced from Wilson, P., Bentley, M.J., Schnabel, C., Clark, R., Xu, S., 2008. Stone run (block stream) formation in the Falkland Islands over several cold stages, deduced from cosmogenic isotope (10Be and 26Al) surface exposure dating. Journal of Quaternary Science 23(5), 461–473 (this figure is available in color online at www.interscience.wiley.com/journal/jqs); (e) and (g) Reproduced from Wilson, P., 2007. Block/rock streams. In: Elias, S.A. (Ed.), Encyclopedia of Quaternary Science/Periglacial Landforms, Rock Forms. Elsevier, Amsterdam, pp. 2217–2225, and (f) Reproduced from Andre´, M.-F., Hall, K., Bertran, P., Arocena, J., 2008. Stone runs in the Falkland Islands: periglacial or tropical Geomorphology 95, 524–543.

410

Evolution of Slopes in a Cold Climate

and lengths up to several hundred meters, whereas widths of up to a few hundred meters and lengths up to 4–5 km have been recorded for the largest block streams (Harris et al., 1998; Barrows et al., 2004; Andre´ et al., 2008; Wilson et al., 2008). In the Falkland Islands, where stone runs are particularly welldeveloped, they typically have a width of B100 m and a length of up to 2 km (Figure 10(a), Hansom et al., 2008). Block streams may also be arranged in a variety of planform patterns (Wilson, 2007). The classical morphology corresponds to a single-channeled straight to sinuous block stream (Figure 10(a)–10(c)), commonly confined in a valley bottom. A variation on this simple configuration consists in a main trunk located in the valley axis and joined by tributary block streams running down hillslopes and forming a dendritic pattern (Figure 10(d)). Multichannel planforms displaying a braided pattern tend to occur on more open slopes (Figure 10(e)). A specific pattern observed in the Falkland Islands corresponds to narrow, parallel block stripes running downhill very close to each other (Figure 10(f)). Block streams generally develop on ground sloping less than 151 (i.e., slightly more than other blockfields), sometimes even on slopes of only 1–21 (Clark, 1972; Hansom et al., 2008). However, longitudinal gradients as steep as 301 have been observed for block streams in mountainous areas (Ginesu, 1990; Harris et al., 1998; Barrows et al., 2004). Block streams forming the downslope continuation of blockfields generally display a concave-up profile (see Clark, 1972). Like blockfields, block streams may either start from the top of convex ridges devoid of visible debris source or be clearly related with tors or rock scarps (Romanovskii and Tyurin, 1986; Boelhouwers, 1999). The rock debris spread at the surface of a block stream may vary in size from a few decimeters to several meters (Potter and Moss, 1968; Clark, 1994). In some cases, they hardly show any longitudinal change in size (Potter and Moss, 1968; Boelhouwers et al., 2002) whereas in others their dimensions decrease downstream (Caine and Jennings, 1968; Cremeens et al., 2005). The average block stream thickness is between 1 and 3 m (Wilson, 2007), with maximum observed values up to 9 m (Boelhouwers, 1999) and minimum values down to B15 cm (Harris et al., 1998). The vertical variations in debris size described in relict block streams, especially in the Falkland Islands, reveal some kind of reverse grading (Figure 10(g); Clark, 1972; Clapperton, 1975; Hansom et al., 2008). Moreover, some stone runs, as also Tasmanian block streams (Caine and Jennings, 1968), have no fine-grained matrix between the blocks at their surface (Figure 10(g)).

8.24.5.3

The block production mechanisms are similar to those already evoked for blockfields, thus including frost shattering (macrogelivation) and thermal stresses (e.g., Clark, 1972; Boelhouwers, 1999). As for the reverse grading of the blocks, it has been assigned to frost heave acting during longitudinal movement (Potter and Moss, 1968; Aldiss and Edwards, 1999; Boelhouwers, 1999). The longitudinal displacement processes and the matrix removal deserve a more detailed discussion, which we turn to now. Since Andersson (1906, 1907), solifluction s.l. is the most cited mass-wasting process to explain the downslope motion of block streams (Caine and Jennings, 1968; Potter and Moss, 1968; Clapperton, 1975; Caine, 1983; Boelhouwers et al., 2002). As a consequence, most authors argue that a longitudinal transport over some distance required that the blocks were embedded in a fine-grained matrix, which was moreover favorable to frost heaving and block sorting. However, frost heaving itself, in combination with frost wedging, was also suggested to cause block motion (Wilson, 2007), in particular for block heaps at the foot of cliffs in the absence of evidence for fine-grained material (Caine and Jennings, 1968; Clark, 1972; Boelhouwers, 1999). According to Caine (1983), most of the downslope transport in Tasmanian block streams may be attributed to sliding failures in the underlying impervious and remolded clays. The mechanism would have involved impeded drainage under deep seasonal freezing, followed by artesian pressure buildup and shearing when the fine-grained and icy matrix between the blocks thawed. Additional longitudinal transport might have been promoted by frost creep and frost heaving but apparently affected only the upper zone of the block streams (Caine, 1983; Boelhouwers, 1999). The final stage of block stream evolution is characterized by the washing or flushing out of the fine-grained matrix from the upper part of the block stream when it is no longer active (Caine and Jennings, 1968; Potter and Moss, 1968; Caine, 1983; Clark, 1994; Boelhouwers, 1999; Boelhouwers et al., 2002). This is needed to explain the block openwork at the surface (Figure 10(g)), and has commonly been related to the end of the longitudinal motion, which allowed running water to percolate through the blocks and wash down the fines. This ultimate reworking is supported to some extent by the frequent observation of water flow at the base of block streams and by the presence of water-transported fine deposits at the toe of some features (see Caine, 1983; Clark, 1994). In addition, Aldiss and Edwards (1999) suggested that the substitution of the matrix by ice during freezing might help to effectively separate the fine and coarse elements by progressive ice-coating of the blocks.

Processes

Once subject of considerable debate, the formation of block streams is now recognized to result from multiple processes operating either in combination or in succession (Wilson, 2007). An outline of the four-stage block stream formation and evolution was already foreshadowed in Andersson’s pioneering work (Andersson, 1906, 1907). It begins with block production processes, then brings into play longitudinal transport by mass-wasting processes also involving vertical sorting of the blocks, and may finally include the removal of interstitial fine-grained material from inactive block streams.

8.24.5.4

(Paleo)Climatic Meaning and Dating of Relict Block Streams

In a review study, Harris (1994) established relations between basic climatic parameters, such as the mean annual air temperature (MAAT) and the mean annual precipitation (MAP), and the occurrence of active block streams worldwide (North America, China, Siberia). The latter occur in areas with cold (MAATB 6 to 20 1C) and dry (MAPB50–500 mm) climatic conditions (Figure 11). Colder and wetter climates were thought to favor, respectively, gelifluction and rock glaciers,

Evolution of Slopes in a Cold Climate

411

Mean annual air temperature MAAT (°C)

5

0

–5

–10 Block stream Gelifluction –15 Rock glacier Overlap 0

500

1000

1500

Mean annual precipitation MAP (mm) Figure 11 Relationship between the occurrence of active block streams worldwide (and other landforms such as rock glaciers and solifluction forms) and basic climatic parameters (mean annual temperature (MAAT) and mean annual precipitation (MAP)). Reproduced from Harris, S.A., 1994. Climatic zonality of periglacial landforms in mountain areas. Arctic 47(2), 184–192, with permission from Arctic Institute of North America.

with limited overlap between the ‘block stream zone’ and those of active gelifluction and rock glaciers (Figure 11, Harris, 1994). Different environmental conclusions were drawn from the analysis of relict landforms in South Africa (Boelhouwers, 1999) and elsewhere (Wilson, 2007). These block streams were active under MAAT of B0 1C, lower than those of today by 7–8 1C. Though they were not linked to the presence of permafrost, up to 2-m-deep seasonal freezing was required in association with a sparse insulating snowcover during the autumnal freeze-up (Boelhouwers, 1999). The age of many block streams has been inferred from their occurrence in areas close to the ice sheet’s maximal extent of the LGM (26.5–20 ka), where they would have experienced harsh periglacial conditions at that time (e.g., Potter and Moss, 1968; Clark, 1972, 1994; Boelhouwers, 1999; Boelhouwers et al., 2002). However, these conclusions were recently revised on the basis of absolute dating of block streams, notably by optically stimulated luminescence (OSL) and cosmogenic nuclides (Barrows et al., 2004; Cremeens et al., 2005; Hansom et al., 2008; Wilson et al., 2008). Exposure ages from 36Cl concentration measurements obtained in various places of Australia (Barrows et al., 2004) indicated a peak of periglacial activity, notably block stream development, between 23 and 16 ka, with a weighted mean of 21.970.5 ka remarkably coincident with the LGM. OSL dating of sediments buried under stone runs in the Falkland Islands also revealed that a major phase of block stream activity occurred under severe periglacial conditions at 32–27 ka, almost coeval with the LGM (Hansom et al., 2008). However, both studies also highlighted other phases of block stream development that significantly predates the LGM. Hansom et al. (2008) obtained a saturated OSL age of Z54 ka, and some cosmogenic ages in Australia pointed to phases of block stream activity around 55–60 and 80–90 ka, and even during the middle Pleistocene (up to 525750 ka, see Barrows et al.,

2004). These results were fully confirmed by 10Be and 26Al exposure ages obtained for block streams in the Falkland Islands (Wilson et al., 2008). In this study, the use of both nuclides moreover allowed the unraveling of complex exposure histories, which highlighted that the stone runs had been inactive during the LGM but developed at different times between 42 and 731 ka, thus demonstrating that large valleybottom block streams may be as old as 700–800 ka (Wilson et al., 2008). It seems thus that periglacial conditions optimal for block stream activity have existed in some areas during the LGM but might have occurred at different times as well in other regions. Moreover, the recently obtained middle Pleistocene ages of some block streams attest their survival over several glacial cycles, with possible reactivations (Wilson et al., 2008). Andre´ et al. (2008) even suggested that the stone runs of the Falkland Islands might have a pre-Quaternary origin. According to them, they would be polygenetic landforms, with a pre-Quaternary production of debris under subtropical or temperate climates, and a periglacial signature restricted to reworking during the Quaternary cold periods.

8.24.6

Research Perspectives

In this section, we emphasized the major landforms that likely reflect the most characteristic expression of slope evolution under periglacial conditions. Thorough morphological descriptions of debated landforms are numerous but the processes implied in their formation probably require further studies, especially on active forms. More fundamentally, paleoenvironmental interpretations, in particular paleoclimatic inferences drawn from presently inactive features, still remain a very delicate and tricky task. In this respect, recent investigation techniques quantifying the rates at which these

412

Evolution of Slopes in a Cold Climate

periglacial landforms evolve or those establishing a reliable chronological framework for relict forms should receive much more attention in the future. Although geophysical methods (such as GPR, electric resistivity, and seismic refraction) were successfully employed to assess sediment thickness and analyze internal structure of talus slopes, their utilization unfortunately remains much targeted and restricted to very specific settings in cold environments. They represent to our opinion a precious tool to yield quantitative data and should therefore be extended to further scopes in the periglacial field. The potential of mathematical or numerical modeling also appears underrated if one refers to their infrequent (or even almost absent) application regarding slope evolution in cold, unglaciated areas. Quantitative, time-integrated simulations may indeed greatly help deciphering the paleoenvironmental meaning of the studied landforms. An equivalent report can be addressed to the utilization of modern absolute dating techniques to relict slope features. Paleoclimatic interpretations commonly rely on unverified assumptions and definitely suffer from the lack of a robust temporal framework. In comparison to the glacial field where they are extensively used to reconstruct the timing of deglaciation in diverse settings, terrestrial cosmogenic nuclides were barely applied in the periglacial environment. The very few studies dealing with this technique to produce absolute ages for relict forms (almost restricted to block streams) however proved that inherent methodological difficulties can be overcome. Dating techniques based on in-situ-produced cosmogenic nuclides (surface exposure and burial dating) thus represent a helpful and powerful alternative to classic 14C and considerably extend the dating timespan for inactive landforms from the Last Glacial up to the early Pleistocene. The same statement can be held for OSL dating although the interest for this method only remains emerging in the periglacial environment.

References Aldiss, D.T., Edwards, E.J., 1999. The geology of the Falkland Islands. Technical Report (WC/99/10). British Geological Survey, Keyworth. Andersson, J.G., 1906. Solifluction, a component of subaerial denudation. Journal of Geology 14, 91–112. Andersson, J.G., 1907. Contributions to the geology of the Falkland Islands. Wissenschaftliche Ergebnisse der Schwedischen Su¨dpolar-Expedition 1901–1903(3), 1–38. Andre´, M.-F., 1997. Holocene rockwall retreat in Svalbard: a triple-rate evolution. Earth Surface Processes and Landforms 22, 423–440. Andre´, M.-F., 2003. Do periglacial landscapes evolve under periglacial conditions? Geomorphology 52(1–2), 149–164. Andre´, M.-F., 2004. The geomorphic impact of glaciers as indicated by tors in North Sweden (Aurivaara, 681N). Geomorphology 57(3–4), 403–421. Andre´, M.-F., Hall, K., Bertran, P., Arocena, J., 2008. Stone runs in the Falkland Islands: periglacial or tropical? Geomorphology 95, 524–543. Ballantyne, C.K., 1984. The Late Devensian periglaciation of upland Scotland. Quaternary Science Reviews 3, 311–343. Ballantyne, C.K., 1998. Age and significance of mountain-top detritus. Permafrost and Periglacial Processes 9, 327–345. Ballantyne, C.K., 2002. Paraglacial geomorphology. Quaternary Science Reviews 21(18–19), 1935–2017. Ballantyne, C.K., 2010. A general model of autochthonous blockfield evolution. Permafrost and Periglacial Processes 21, 289–300. Ballantyne, C.K., Harris, C., 1994. The Periglaciation of Great-Britain. Cambridge University Press, Cambridge, 330 pp.

Ballantyne, C.K., McCarroll, D., Nesje, A., Dahl, S.O., Stone, J.O., 1998. The last ice sheet in North-West Scotland: reconstruction and implications. Quaternary Science Reviews 17, 1149–1184. Barrows, T.T., Stone, J.O., Fifield, L.K., 2004. Exposure ages for Pleistocene periglacial deposits in Australia. Quaternary Science Reviews 23, 697–708. Bastin, B., Guillien, Y., 1971. Approche palynologique des gre`zes lite´es de Sonneville et d’E´choisy (Charentes). Comptes rendus de l’Acade´mie des Sciences (Paris) 273, 2063–2066. Bertran, P., 1992. Micromorphologie de gre`zes lite´es des Charentes et du Chaˆtillonnais (France). Quaternaire 3(1), 4–15. Bertran, P., Coutard, J.-P., Francou, B., Ozouf, J.-C., Texier, J.-P., 1992. Donne´es nouvelles sur l’origine du litage des gre`zes: implications pale´oclimatiques. Ge´ographie Physique et Quaternaire 46(1), 97–112. Bertran, P., Francou, B., Pech, P., 1993. Stratogene`se associe´e a` la dynamique des coule´es a` front pierreux en milieu alpin, La Mortice, Alpes Me´ridionales, France. Ge´ographie Physique et Quaternaire 47, 93–100. Bierman, P.R., Marsella, K.A., Patterson, C., Thompson Davis, P., Caffee, M., 1999. Mid-Pleistocene cosmogenic minimum-age limits for pre-Wisconsinan glacial surfaces in southwestern Minnesota and southern Baffin Island: a multiple nuclide approach. Geomorphology 27, 25–39. Blikra, L.H., Nemec, W., 1998. Postglacial colluvium in western Norway: depositional processes, facies and palaeoclimatic record. Sedimentology 45, 909–959. Boardman, J., 1978. Gre`zes lite´es near Keswick, Cumbria. Biuletyn Peryglacjalny 27, 23–34. Boelhouwers, J., 1999. Relict periglacial slope deposits in the Hex River Mountains, South Africa: observations and Palaeoenvironmental Implications. Geomorphology 30(3), 245–258. Boelhouwers, J., 2004. New perspectives in autochtonous blockfield development. Polar Geography 28(2), 133–146. Boelhouwers, J., Holness, S., Meiklejohn, K.I., Sumner, P.D., 2002. Observations on a blockstream in the vicinity of Sani Pass, Lesotho highlands, Southern Africa. Permafrost and Periglacial Processes 13(4), 251–257. Boelhouwers, J., Holness, S., Sumner, P.D., 2003. The maritime Subantarctic: a distinct periglacial environment. Geomorphology 52, 39–55. Bouchard, S., Jolicoeur, S., Guillaume, P., 1995. Characteristics and significance of two pre-late Wisconsinan weathering profiles (Adirondacks, USA and Miramichi Highlands, Canada). Geomorphology 12, 75–89. Caine, N., 1983. The Mountains of Northeastern Tasmania. A Study of Alpine Geomorphology. Balkema, Rotterdam, 200 pp. Caine, N., Jennings, J., 1968. Some blockstreams of the Toolong Range, Kosciusko State Park, New South Wales. Journal and Proceedings of the Royal Society of New South Wales 101, 93–103. Church, M., Stock, R.F., Ryder, J.M., 1979. Contemporary sedimentary environments on Baffin Island, N.W.T., Canada: Debris slope accumulations. Arctic and Alpine Research 11, 371–402. Clapperton, C.M., 1975. Further observations on the stone runs of the Falkland Islands. Biuletyn Peryglacjalny 24, 211–217. Clark, R., 1972. Periglacial landforms and landscapes in the Falkland Islands. Biuletyn Peryglacjalny 21, 33–50. Clark, R., 1994. Tors, rock platforms and debris slopes at Stiperstones, Shropshire, England. Field Studies 8, 451–472. Colhoun, E.A., 2002. Periglacial landforms and deposits of Tasmania. South African Journal of Science 98, 55–63. Coltorti, M., Dramis, F., 1995. The chronology of upper Pleistocene stratified slopewaste deposits in central Italy. Permafrost and Periglacial Processes 6, 235–242. Coltorti, M., Dramis, F., Pambianchi, G., 1983. Stratified slope-waste deposits in the Esino River basin (Umbria-Marche Apennines, Central Italy). Polarforschung 53(2), 59–66. Cremeens, D.L., Darmody, R.G., George, S.E., 2005. Upper Slope Landforms and Age of Bedrock Exposures in the St. Francois Mountains, Missouri: A Comparison to Relict Periglacial Features in the Appalachian Plateau of West Virginia. Geomorphology 70, 71–84. Curry, A.M., Black, R., 2003. Structure, sedimentology and evolution of rockfall talus, Mynydd Du, south Wales. Proceedings Geologists’ Association 114, 49–64. Curry, A.M., Morris, C.J., 2004. Lateglacial and Holocene talus slope development and rockwall retreat on Mynydd Du, UK. Geomorphology 58, 85–106. Czudek, T., 1993. Pleistocene periglacial structures and landforms in Western Czechoslovakia. Permafrost and Periglacial Processes 4, 65–75. Czudek, T., 1995. Cryoplanation terraces: a brief review and some remarks. Geografiska Annaler A 77(1–2), 95–105.

Evolution of Slopes in a Cold Climate

Czudek, T., Demek, J., 1973. The valley cryopediments in Eastern Siberia. Biuletyn Peryglacjalny 22, 117–130. Dahl, E., 1987. The nunatak theory reconsidered. Ecological Bulletins 38, 77–94. Dahl, R., 1966. Block fields, weathering pits and tor-like forms in the Narvik Mountains, Nordland, Norway. Geografiska Annaler A 48, 55–85. Darmody, R.G., Thorn, C.E., Seppa¨la¨, M., Campbell, S.W., Li, Y.K., Harbor, J., 2008. Age and weathering status of granite tors in arctic Finland (B681 N). Geomorphology 94, 10–23. Demek, J., 1969. Cryogene processes and the development of cryoplanation terraces. Biuletyn Peryglacjalny 18, 115–125. De Wolf, Y., 1988. Stratified slope deposits. In: Clark, M.J. (Ed.), Advances in Periglacial Geomorphology. Wiley, New York, pp. 113–149. Dixon, J.C., Thorn, C.E., 2005. Chemical weathering and landscape development in mid-latitude alpine environments. Geomorphology 67(1–2), 127–145. Dredge, L.A., 1992. Breakup of limestone bedrock by frost shattering and chemical weathering, Eastern Canadian Arctic. Arctic and Alpine Research 24(4), 314–323. Dredge, L.A., 2000. Age and origin of upland block fields on Melville Peninsula, eastern Canadian Arctic. Geografiska Annaler A 82, 443–454. Dyke, A.S., 1993. Landscapes of cold-centered late Wisconsin ice caps, Arctic Canada. Progress in Physical Geography 17, 223–247. Dylik, J., 1957. Tentative comparisons of planation surfaces occurring under warm and under cold semiarid climatic conditions. Biuletyn Peryglacjalny 5, 175–186. Eaton, L.J., Morgan, B.A., Kochel, R.C., Howard, A.D., 2003. Quaternary deposits and landscape evolution of the central Blue Ridge of Virginia. Geomorphology 56, 139–154. Fjellanger, J., Nystuen, J.P., 2007. Diagenesis and weathering of quartzite at the paleic surface on the Varanger Peninsula, northern Norway. Norwegian Journal of Geology 87, 133–145. Fjellanger, J., Sørbel, L., Linge, H., Brook, E.J., Raisbeck, G.M., Yiou, F., 2006. Glacial survival of blockfields on the Varanger Peninsula, northern Norway. Geomorphology 82(3–4), 255–272. Francou, B., 1988. Eboulis stratifie´s dans les Hautes Andes Centrales du Pe´rou. Zeitschrift fu¨r Geomorphologie (Neue Folge) 32, 47–76. Francou, B., 1989. La stratoge´ne`se dans les formations de pente soumises a` l’action du gel. Une nouvelle conception du proble`me. Bulletin de l’Association franc- aise pour l’e´tude du Quaternaire 26(4), 185–199. Francou, B., 1990. Stratification mechanisms in slope deposits in high subequatorial mountains. Permafrost and Periglacial Processes 1, 249–263. Francou, B., Mante´, C., 1990. Analysis of the segmentation in the profile of alpine talus slopes. Permafrost and Periglacial Processes 1(1), 53–60. French, H.M., 2000. Does Lozinski’s periglacial realm exist today? A discussion relevant to modern usage of the term ‘periglacial’. Permafrost and Periglacial Processes 11, 35–42. French, H.M., 2007. The Periglacial Environment, third ed. John Wiley & Sons, Chichester, UK, 458 pp. French, H.M., Harry, D.G., 1992. Pediments and cold-climate conditions, Barn Mountains, unglaciated northern Yukon, Canada. Geografiska Annaler A 74, 145–157. Garcia-Ruiz, J.M., Valero-Garce´s, B., Gonza´lez-Sampe´riz, P., Lorente, A., MartiBono, C., Begueria, S., Edwards, L., 2001. Stratified scree in the central Spanish Pyrenees: palaeoenvironmental implications. Permafrost and Periglacial Processes 12, 233–242. Gardner, J.S., 1980. Frequency, magnitude and spatial distribution of mountain rockfalls and rockslides in the Highwood Pass area, Alberta, Canada. In: Coates, D.R., Vitek, J. (Eds.), Thresholds in Geomorphology. Allen and Unwin, London, pp. 267–295. Gardner, T.W., Ritter, J.B., Shuman, C.A., Bell, J.C., Sasowsky, K.C., Pinter, N., 1991. A periglacial stratified slope deposit in the Valley and Ridge province of central Pennsylvania, USA: sedimentology, stratigraphy, and geomorphic evolution. Permafrost and Periglacial Processes 2, 141–162. Gengnian, L., Zhijiu, C., Daokai, G., Yongqiu, W., 1999. The Stratified slope deposits at Kunlunshan Pass, Tibet Plateau, China. Permafrost and Periglacial Processes 10, 369–375. Ginesu, S., 1990. Periglacial deposits in Sardinia: the blockstreams near Pranu Mannu. Geografia Fisica e Dinamica Quaternaria 13, 179–181. Goddard, A., 1989. Les vestiges des manteaux d’alte´ration sur les socles des hautes latitutes? Indentification, signification. Zeitschrift fu¨r Geomorphologie (Supplementband) 72, 1–20. Goodfellow, B.W., 2007. Relict non-glacial surfaces in formerly glaciated landscapes. Earth-Science Reviews 80, 47–73. Goodfellow, B.W., Stroeven, A.P., Ha¨ttestrand, C., Kleman, J., Jansson, K.N., 2008. Deciphering a non-glacial/glacial landscape mosaic in the northern Swedish mountains. Geomorphology 93(3–4), 213–232.

413

Gorbunov, A.P., Marchenko, S.S., Seversky, E.V., 2004. The termal environment of blocky materials in the mountains of central Asia. Permafrost and Periglacial Processes 15, 95–98. Grosso, S.A., Corte, A.E., 1991. Cryoplanation surfaces in the central Andes at latitude 35 1S. Permafrost and Periglacial Processes 2, 49–58. Guillien, Y., 1951. Les gre`zes lite´es de Charente. Revue de Ge´ographie des Pyrene´es et du Sud-Ouest 22, 155–162. Guillien, Y., 1964. Gre`zes lite´es et bancs de neige. Geologie en Mijnbouw 43, 103–112. Guillien, Y., Lautridou, J.P., 1970. Recherches de ge´lifraction expe´rimentale du Centre de Ge´omorphologie. I – Calcaires des charentes, Essais au gel de quelques calcaires charentais. Bulletin no. 5 du Centre de Ge´omorphologie du CNRS, Caen, 49 p. Hales, T.C., Roering, J.J., 2005. Climate-controlled variations in scree production, Southern Alps, New Zealand. Geology 33, 701–704. Hall, A.M., Mellor, A., Wilson, M.J., 1989. The clay mineralogy and age of deeplyweathered rock in north-east Scotland. Zeitschrift fu¨r Geomorphologie (Supplementband) 72, 47–63. Hall, A.M., Phillips, W.M., 2006. Glacial modification of granite tors in the Cairngorms, Scotland. Journal of Quaternary Science 21, 811–830. Hall, K.J., 1998. Nivation or cryoplanation: different terms, same features? Polar Geography 22(1), 1–16. Hall, K.J., 1999. The role of thermal stress fatigue in the breakdown of rock in cold regions. Geomorphology 31, 47–63. Hall, K.J., Andre´, M.-F., 2003. Rock thermal data at the grain scale: applicability to granular disintegration in cold environments. Earth Surface Processes and Landforms 28, 823–836. Hall, K.J., Andre´, M.-F., 2010. Some further observations regarding ‘‘cryoplanation terraces’’ on Alexander Island. Antarctic Science 22(2), 175–183. Hall, K.J., Thorn, C.E., Matsuoka, N., Prick, A., 2002. Weathering in cold regions: some thoughts and perspectives. Progress in Physical Geography 26, 577–603. Hansom, J.D., Evans, D.J.A., Sanderson, D.C.W., Bingham, R.G., Bentley, M.J., 2008. Constraining the age and formation of stone runs in the Falkland Islands using optically stimulated luminescence. Geomorphology 94, 117–130. Hanvey, P.M., Lewis, C.A., 1991. Sedimentology and Genesis of Slope Deposits at Sonskyn, Eastern Cape Drakensberg, South Africa. Permafrost and Periglacial Processes 2, 31–38. Harris, C., 2007. Slope Deposits and Forms. In: Elias, S.A. (Ed.), Encyclopedia of Quaternary Science/Periglacial Landforms, Slope Deposits and Forms. Elsevier, Amsterdam, pp. 2207–2217. Harris, S.A., 1994. Climatic zonality of periglacial landforms in mountain areas. Arctic 47(2), 184–192. Harris, S.A., Cheng, G., Zhao, X., Yongqin, D., 1998. Nature and dynamics of an active block stream, Kunlun Pass, Qinghai Province, People’s Republic of China. Geografiska Annaler A 80(2), 123–133. Harris, S.A., Pedersen, D.E., 1998. Thermal regimes beneath coarse blocky materials. Permafrost and Periglacial Processes 9, 107–120. Harris, S.A., Prick, A., 2000. Conditions of formation of stratified screes, Slims River Valley, Yukon Territory: a possible analogue with some deposits from Belgium. Earth Surface Processes and Landforms 25(5), 463–481. Ha¨ttestrand, C., 1994. Boulder depressions in central Sweden – Remnants of a preLate Weichselian landscape? Geografiska Annaler A 76(3), 153–160. Ha¨ttestrand, C., Stroeven, A.P., 2002. A relict landscape in the centre of Fennoscandian glaciation; geomorphological evidence of minimal Quaternary glacial erosion. Geomorphology 44(1–2), 127–143. He´tu, B., 1991. Eboulis stratifie´s actifs pre`s de Manche-d’Epe´e, Gaspe´sie (Que´bec, Canada). Zeitschrift fu¨r Geomorphologie (Neue Folge) 35, 439–461. He´tu, B., Gray, J.T., 2000. Effects of environmental change on scree slope development throughout the postglacial period in the Chic-Choc Mountains in the northern Gaspe´ Peninsula, Que´bec. Geomorphology 32, 135–155. He´tu, B., Van Steijn, H., Vandelac, P., 1994. Les coule´es de pierres glace´es: un nouveau type de coule´es de pierraille sur les talus d’e´boulis. Ge´ographie Physique et Quaternaire 48, 3–22. Hinchliffe, S., 1999. Timing and significance of talus slope reworking, Trotternish, Skye, northwest Scotland. The Holocene 9, 483–494. Hinchliffe, S., Ballantyne, C.K., 2009. Talus structure and evolution on sandstone mountains in NW Scotland. The Holocene 19(3), 477–486. Hinchliffe, S., Ballantyne, C.K., Walden, J., 1998. The structure and sedimentology of relict talus, Trotternish, northern Skye, Scotland. Earth Surface Processes and Landforms 23, 545–560. Hoffmann, T., Schrott, L., 2002. Modelling sediment thickness and rockwall retreat in an Alpine valley using 2D-seismic refraction (Reintal, Bavarian Alps). Zeitschrift fu¨r Geomorphologie 127, 153–173.

414

Evolution of Slopes in a Cold Climate

Ives, J.D., 1958. Mountain-top detritus and the extent of the last glaciation in northeastern Labrador-Ungava. The Canadian Geographer 12, 25–31. Jomelli, V., Francou, B., 2000. Comparing the characteristics of rockfall talus and snow avalanche landforms in an Alpine environment using a new methodological approach: Massif des Ecrins, French Alps. Geomorphology 35, 181–192. Journaux, A., 1976. Les gre`zes lite´es du Chaˆtillonais. Bulletin de l’Association francaise pour l’e´tude du Quaternaire 3–4, 123–138. Juliussen, H., Humlum, O., 2008. Thermal regime of openwork blockfields on the mountains Elga˚hogna and Sølen, central western Norway. Permafrost and Periglacial Processes 19, 1–18. Kleman, J., Borgstro¨m, I., 1990. The boulder fields of Mt. Fulufjallet, west-central Sweden. Geografiska Annaler A 72, 63–78. Lauriol, B.M., Godbout, L., 1988. Les terrasses de cryoplanation dans le nord du Yukon: distribution, gene`se et aˆge. Ge´ographie Physique et Quaternaire 42, 303–314. Lauriol, B.M., Lalonde, A.E., Dewez, V., 1997. Weathering of quartzite on a cryoplanation terrace in Northern Yukon, Canada. Permafrost and Periglacial Processes 8, 147–153. Lautridou, J.P., Seppa¨la¨, M., 1986. Experimental frost shattering of some Precambrian rocks, Finland. Geografiska Annaler A 68(1–2), 89–100. Luckman, B.H., 1978. Geomorphic work of snow avalanches in the Canadian Rockies. Arctic and Alpine Research 10, 261–276. Luckman, B.H., 1988. Debris accumulation patterns on talus slopes in Surprise Valley, Alberta. Ge´ographie Physique et Quaternaire 42(3), 247–278. Luckman, B.H., 1992. Debris flow and snow avalanche landforms in the Lairig Ghru, Cairngorm Mountains, Scotland. Geografiska Annaler A 74, 109–121. Luckman, B.H., 2007. Talus Slopes. In: Elias, S.A. (Ed.), Encyclopedia of Quaternary Science/Periglacial Landforms, Rock Forms. Elsevier, Amsterdam, pp. 2242–2249. Luckman, B.H., Fiske, C.J., 1997. Holocene development of coarse debris landforms in the Canadian Rockies. In: Matthews, J., Brunsden, D., Frenzel, B., Weiss, M. (Eds.), Rapid Mass Movement and Climatic Variation during the Holocene, Pala¨oklimaforschung, pp. 283–297. Lundqvist, G., 1944. De svenska fja¨llens natur. Svensk Turistfo¨rening Fo¨rlag, Stockholm. Marquette, G.C., Gray, J.T., Gosse, J.C., Courschene, F., Stockli, L., Macpherson, G., Finkel, R., 2004. Felsenmeer, persistence under non-erosive ice in the Torngat and Kaumajet mountains, Quebec and Labrador, as determined by soil weathering and cosmogenic nuclide exposure dating. Canadian Journal of Earth Sciences 41, 19–38. Matsuoka, N., 2001. Microgelivation versus macrogelivation: towards bridging the gap between laboratory and field frost weathering. Permafrost and Periglacial Processes 12(3), 299–313. Matsuoka, N., 2008. Frost weathering and rockwall erosion in the southeastern Swiss Alps: long-term observations. Geomorphology 99, 353–368. McCarroll, D., Ballantyne, C.K., Nesje, A., Dahl, S., 1995. Nunataks of the last ice sheet in north-west Scotland. Boreas 24, 305–323. McCarroll, D., Shakesby, R.A., Matthews, J.A., 1998. Spatial and temporal patterns of Holocene rockfall activity on a Norwegian talus slope estimated by lichenometry. Arctic and Alpine Research 30, 51–60. Mellor, A., Wilson, M.J., 1989. Origin and significance of gibbsitic montane soils in Scotland, UK. Arctic and Alpine Research 21, 417–424. Murton, J.B., Peterson, R., Ozouf, J.-C., 2006. Bedrock fracture by ice segregation in cold regions. Science 314, 1127–1129. Nelson, F.E., 1989. Cryoplanation terraces: periglacial cirque analogues. Geografiska Annaler A 68, 207–211. Nelson, F.E., 1998. Cryoplanation terrace orientation in Alaska. Geografiska Annaler A 80, 135–151. Nemec, W., Kazanci, N., 1999. Quaternary colluvium in west-central Anatolia: sedimentary facies and palaeoclimatic significance. Sedimentology 46, 139–170. Nesje, A., 1989. The geographical and altitudinal distribution of blockfields in southern Norway and its significance to the Pleistocene ice sheets. Zeitschrift fu¨r Geomorphologie (Supplementband) 72, 41–53. Nesje, A., Dahl, S.O., Anda, E., Rye, N., 1988. Blockfields in southern Norway: significance for the Late Weichselian ice sheet. Norsk Geologisk Tidsskrift 68, 149–169. Obanawa, H., Matsukura, Y., 2006. Mathematical modeling of talus development. Computers and Geosciences 32, 1461–1478. Otto, J.C., Sass, O., 2006. Comparing geophysical methods for talus slope investigations in the Turtmann valley (Swiss Alps). Geomorphology 76(3–4), 257–272. Paasche, Ø., Strømsøe, J.R., Dahl, S.E., Linge, H., 2006. Weathering characteristics of arctic islands in northern Norway. Geomorphology 82(3–4), 430–452.

Pappalardo, M., 1999. Observations on stratified slope deposits, Gesso Valley, Italian Maritime Alps. Permafrost and Periglacial Processes 10, 107–111. Perez, F.L., 1993. Talus movement in the high equatorial Andes: a synthesis of ten years of data. Permafrost and Periglacial Processes 4, 199–215. Phillips, W.M., Hall, A.M., Mottram, R., Fifield, L.K., Sudgen, D.E., 2006. Cosmogenic 10Be and 26Al exposure ages of tors and erratics, Cairngorm Mountains, Scotland: timescales for the development of a classic landscape of selective linear glacial erosion. Geomorphology 73(3–4), 224–245. Pieracci, K., Lambiel, C., Reynard, E., 2008. La re´partition du perge´lisol dans trois e´boulis alpins du massif de la Dent de Morcles (Valais, Alpes suisses). Ge´omorphologie, Relief, Processus, Environnement 2, 87–98. Piirola, J., 1969. Frost-sorted block concentrations in western Inari, Finnish Lapland. Fennia 99(2), 1–35. Pissart, A., 1995. L’Ardenne sous le joug du froid. Le modele´ pe´riglaciaire du massif ardennais. In: Demoulin, A. (Ed.), L’Ardenne. Essai de Ge´ographie Physique. Universite´ de Lie`ge, Lie`ge, pp. 136–154. Potter, N., Moss, J.H., 1968. Origin of the Blue Rocks block field and adjacent deposits, Berks County, Pennsylvania. Geological Society of America Bulletin 79, 255–262. Potts, A.S., 1970. Frost action in rocks: some experimental data. Institute of British Geographers Transactions 49, 109–124. Priesnitz, K., 1981. FuXfla¨chen und Ta¨ler in der Arktis NW-Kanadas und Alaskas. Polarforschung 51(2), 145–159. Rapp, A., 1958. Talus slopes and Mountain Walls at Templefjorden, Spitzbergen. Norsk Polarinstitutt Skrifter 119, 96. Rapp, A., 1960. Recent development of mountain slopes in Ka¨rkevagge and surroundings. Geografiska Annaler A 42(2/3), 65–200. Rapp, A., Rudberg, S., 1960. Recent periglacial phenomena in Sweden. Biuletyn Peryglacjalny 8(3), 143–154. Rea, B.R., 2007. Blockfields (Felsenmeer). In: Elias, S.A. (Ed.), Encyclopedia of Quaternary Science/Periglacial Landforms, Rock Forms, Elsevier, Amsterdam, pp. 2225–2236. Rea, B.R., Whalley, W.B., Porter, E.M., 1996b. Rock weathering and the formation of summit blockfields slopes in Norway: Examples and implications. In: Anderson, E., Brooks (Eds.), Advances in Hillslope Processes. John Wiley and Sons Ltd., Chichester, pp. 1257–1275. Rea, B.R., Whalley, W.B., Rainey, M.M., Gordon, J.E., 1996a. Blockfields, old or new? Evidence and implications from some plateaus in northern Norway. Geomorphology 15, 109–121. Reger, R.D., Pewe, T.L., 1976. Cryoplanation terraces: indicators of a permafrost environment. Quaternary Research 6, 99–109. Roaldset, E., Pettersen, E., Longva, O., Mangerud, J., 1982. Remnants of preglacial weathering in western Norway. Norsk Geologisk Tidsskrift 62(3), 169–178. Romanovskii, N.N., Tyurin, A.I., 1986. Kurums. Biuletyn Peryglacjalny 31, 249–259. Sass, O., 2006. Determination of the internal structure of alpine talus deposits using different geophysical methods (Lechtaler Alps, Austria). Geomorphology 80, 45–58. Sass, O., 2007. Bedrock detection and talus thickness assessment in the European Alps using geophysical methods. Journal of Applied Geophysics 62(3), 254–269. Sass, O., Krautblatter, M., 2007. Debris flow-dominated and rockfall-dominated talus slopes: genetic models derived from GPR measurements. Geomorphology 86(1–2), 176–192. Sass, O., Wollny, K., 2001. Investigations regarding Alpine talus slopes using ground-penetrating radar (GPR) in the Bavarian Alps, Germany. Earth Surface Processes and Landforms 26(10), 1071–1086. Sauchyn, D.J., 1986. Particle size and shape variation on alpine debris fans, Canadian Rocky Mountains. Physical Geography 7, 191–217. Smith, B.J., Wright, J.S., Whalley, W.B., 2002. Sources of non-glacial, loess-size quartz silt and the origins of ‘‘desert loess’’. Earth-Science Reviews 59, 1–26. Statham, I., 1976. A scree slope rockfall model. Earth Surface Processes and Landforms 1(1), 43–62. Stroeven, A.P., Fabel, D., Ha¨ttestrand, C., Harbor, J., 2002. A relict landscape in the centre of Fennoscandian glaciation: cosmogenic radionuclide evidence of tors preserved through multiple glacial cycles. Geomorphology 44(1–2), 145–154. Sudgen, D.E., Watts, S.H., 1977. Tors, felsenmeer, and glaciations in northern Cumberland Peninsula, Baffin Island. Canadian Journal of Earth Sciences 14, 2817–2823. Sumner, P.D., Meiklejohn, K.I., 2004. On the development of autochthonous blockfields in the grey basalts of sub-Antarctic Marion Island. Polar Geography 28(2), 120–132.

Evolution of Slopes in a Cold Climate

Sumner, P.D., Meiklejohn, K.I., Nel, W., Hedding, D.W., 2004. Thermal attributes of rock weathering: zonal or azonal? A comparison of rock temperatures in different environments. Polar Geography 28(2), 79–92. Thorn, C.E., Hall, K., 2002. Nivation and cryoplanation: the case for scrutiny and integration. Progress in Physical Geography 26(4), 533–550. Tricart, J., Cailleux, A., 1967. Le modele´ des re´gions pe´riglaciaires. Socie´te´ d’e´dition d’enseignement supe´rieur, Paris, 512 pp. Vandenberghe, J., Czudek, T., 2008. Pleistocene cryopediments on variable terrain. Permafrost and Periglacial Processes 19, 71–83. Van Steijn, H., Bertran, P., Francou, B., He´tu, B., Texier, J.-P., 1995. Models for the genetic and environmental interpretation of stratified slope deposits: review. Permafrost and Periglacial Processes 6, 125–146. Van Steijn, H., Boelhouwers, J., Harris, S., He´tu, B., 2002. Recent research on the nature, origin and climatic relations of blocky and stratified slope deposits. Progress in Physical Geography 26(4), 551–575. Van Steijn, H., De Ruig, J., Hoozemans, F., 1988. Morphological and mechanical aspects of debris flows in parts of the French Alps. Zeitschrift fu¨r Geomorphologie (Neue Folge) 32, 143–161. Van Steijn, H., He´tu, B., 1997. Rain-generated overland flow as a factor in the development of some stratified slope deposits: a case study from the Pays du Bue¨ch (Pre´alpes, France). Ge´ographie Physique et Quaternaire 51, 3–15. Van Steijn, H., Van Brederode, L., Goedheer, J.C., 1984. Stratified slope deposits of the gre`zes-lite´es type in the Arde`che region in the south of France. Geografiska Annaler A 66, 295–305. Walder, J.S., Hallet, B., 1985. A theoretical model of the fracture of rock during freezing. Geological Society of America Bulletin 96(3), 336–346.

415

Walder, J.S., Hallet, B., 1986. The physical basis of frost weathering: toward a more unified and fundamental theory. Arctic and Alpine Research 18(1), 27–32. Washburn, A.L., 1979. Geocryology. A Survey of Periglacial Processes and Environment. Edward Arnold, London, 406 pp. Wasson, R.J., 1979. Stratified debris deposits in the Hindu Kush, Pakistan. Zeitschrift fu¨r Geomorphologie (Neue Folge) 23(3), 301–320. Watson, E., 1965. Gre`zes lite´es ou e´boulis tardiglaciaires dans la re´gion d’Aberystwyth. Bulletin de l’Association des Ge´ographes Franc- ais 38(9), 16–25. Whalley, W.B., McGreevy, J.P., 1987. Weathering. Progress in Physical Geography 11(3), 357–369. Whalley, W.B., Rea, B.R., Rainey, M.M., 2004. Weathering, blockfields and fracture systems and the implications for long-term landscape formation: some evidence from Lyngen and O¨ksfjordjokelen areas in north Norway. Polar Geography 28, 93–119. Whalley, W.B., Rea, B.R., Rainey, M.M., McAlister, J.J., 1997. In: Widdowson, M. (Ed.), Rock Weathering in Blockfields: Some Preliminary Data from Mountain Plateaus in North Norway. Geological Society of London, Special Publication, vol. 120, pp. 133–145. Wilson, P., 2007. Block/rock streams. In: Elias, S.A. (Ed.), Encyclopedia of Quaternary Science/Periglacial Landforms, Rock Forms. Elsevier, Amsterdam, pp. 2217–2225. Wilson, P., Bentley, M.J., Schnabel, C., Clark, R., Xu, S., 2008. Stone run (block stream) formation in the Falkland Islands over several cold stages, deduced from cosmogenic isotope (10Be and 26Al) surface exposure dating. Journal of Quaternary Science 23(5), 461–473. Wiman, S., 1963. A preliminary study of experimental frost weathering. Geografiska Annaler A 45(2–3), 113–121.

Biographical Sketch Born in 1982 in Belgium, Gilles Rixhon obtained his diploma degree in Geography in 2005 at the University of Lie´ge. He completed his Ph.D. in 2010, cosupervised between the Universities of Lie`ge and Aix en Provence – Marseille III (CEREGE). His dissertation was related to the reconstruction of the Quaternary river incision in NW Europe, using terrestrial cosmogenic nuclides as dating method. He works now as research assistant in the Department of Geography in the University of Cologne (with Prof. Bru¨ckner). His research topics mostly concern the application of cosmogenic nuclides in the coastal and fluvial environments.

Born in 1957, Alain Demoulin obtained a PhD in 1984 and the title Agre´ge´ de Faculte´ in 2006 from the University of Lie`ge (Belgium). He published about 90 scientific articles and two books. He is currently Senior Research Associate of the FSR-FNRS in Belgium and works in the Department of Physical Geography and Quaternary of the University of Lie`ge. His main research interests are in the field of neotectonics and morphotectonics of intraplate areas. Since the 1990s, he has carried out several geodetic studies of crustal movements in W Europe. He also conducts researches about the drainage network response to tectonic perturbations in the frame of the Quaternary climatic variations.