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Downwasting glacier surface AD 1945-1978


Zone of rock-slope failure,. 1972.


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Figure 17.4 A) Cross-section through a rockslide initiated by recent downwastage of the Maud glacier in New Zealand (adapted from McSaveney, 1993). B) Cross-section through a rock slope failure initiated by downwastage of an outlet glacier of the Myrdalsjokull ice cap, southern Iceland (adapted from Sigurdsson and Williams, 1991).

assumptions the overall probability of failure within a given area diminishes exponentially with time elapsed since deglaciation as the number of potential (i.e. unfailed) failure sites is progressively reduced. For glacially steepened overdip slopes in the Canadian Rockies, they calculated that the average pre-failure lifespan of individual sites is c. 5,700 years, implying that glacially conditioned rock-slope failure may occur many millennia after deglaciation in response to progressive stress release and consequent ramification of the joint network causing reduction of rock-mass strength. Some support for their proposal comes from Yosemite Valley, California, where the volume of rock-slope debris deposited at the foot of granite rockwalls over the period AD 1851—1992 implies an average accumulation rate that is less than half the postglacial average (Wieczorek and Jäger, 1996), consistent with a progressive slowing in the rate of rockslope failure as predicted by the exhaustion model.

17.3.2 Rock-Slope Deformation

Stress release due to deglacial unloading and debuttressing may also result in slow paraglacial rock-slope deformation, also referred to as rock-mass creep. Such deformation represents wholesale failure of large rock masses, but occurs (at least initially) without catastrophic runout of debris. Landforms characteristic of rock-slope deformation include ridge-top graben, crevasse-like tension cracks, antiscarps and convex 'bulging' slopes (Fig. 17.5). Such deformation may reduce slopes to a state of conditional stability, and sometimes precedes catastrophic failure (Chigira, 1992).

A close relationship between debuttressing of rock slopes during deglaciation and initiation of rock slope deformation has been observed by several authors (Tabor, 1971; Radbruch-Hall, 1978). In New Zealand, thinning of the Tasman Glacier has resulted in widespread deformation of the adjacent bedrock slope, producing tension cracks, flexural topples and antiscarps (Blair, 1994). Similarly, the slopes overlooking the downwasting Affliction Glacier in British Columbia exhibit widespread evidence of recent rock-slope deformation, including fractures, antiscarps, elongate graben and collapse pits, collectively indicative of tensional deformation in the near-surface zone. A survey of rock-slope movement at this site by Bovis (1990) indicates that glacier thinning has

Paraglacial Mass Movement
Figure 17.5 Antiscarps and bulging slopes due to paraglacial rock mass deformation on Beinn Fhada, Scotland.

triggered gravitational movement of ~3 X 107 m3 of rock, with surface velocities of a few millimetres to a few centimetres per year (Fig. 17.6). The widespread occurrence of similar indications of rock-slope deformation in mountainous areas that now lack glacier ice strongly suggests that paraglacial rock-slope deformation commonly accompanied deglaciation in Late Pleistocene or Early Holocene times (e.g. Tabor, 1971; Radbruch-Hall et al., 1976; Mahr, 1977; Radbruch-Hall, 1978; Jarman and Ballantyne, 2002).

17.3.3 Paraglacial Rockfall and Talus Accumulation

The third type of response of glacially steepened rockwalls to deglaciation is through enhanced rockfall activity that results in the development of paraglacial talus accumulations (Augustinus, 1995a). Several authors have noted that the large volumes of talus at the foot of rockwalls deglaciated at the end of the Pleistocene are inconsistent with presently modest rockfall activity, and have concluded that the rockfall inputs were very much greater immediately after deglaciation (e.g. Luckman, 1981; Gardner, 1982; Johnson, 1984, 1995; Marion et al., 1995). Luckman and Fiske (1995) found that rockfall input over the past 300 years at a site in the Canadian Rockies has been roughly an order of magnitude too low to have produced the volume of talus now present. Similarly, Hinchliffe and Ballantyne (1999) found that at a site in the Scottish Highlands, roughly 80 per cent of talus accumulation took place within the first 6,000 years after deglaciation, and only about 20 per cent in the ensuing 11,000 years, implying a marked reduction in rockfall rate through time.

Though enhanced rockfall immediately after deglaciation may partly reflect freeze-thaw activity under former periglacial conditions, there is evidence that paraglacial stress-release and associated

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Figure 17.6 Cross-section depicting vectors of paraglacial rock slope deformation triggered by downwastage of Affliction Glacier, British Columbia. The vectors represent 4 years of movement. The position of the groundwater table (GWT) and the underlying shear plane are schematic. (Adapted from Bovis (1990).)

Tension cracks, graben and antiscarps

Tension cracks, graben and antiscarps

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rock-slope instability has exerted the dominant control on postglacial rockfall rates. High rockfall inputs from densely jointed rockwalls in New Zealand have resulted in rapid talus accumulation near the snouts of retreating glaciers (Augustinus, 1995a), and the recent retreat of glaciers on Mexican stratovolcanoes has similarly been accompanied by rapid rockwall degradation and talus accumulation (Palacios and de Marcos, 1998), yet neither location currently experiences severe periglacial conditions. André (1997) has shown that on Svalbard, rates of rockwall retreat due to paraglacial stress release are roughly an order of magnitude higher that those due to freeze-thaw effects. If such findings are representative, they imply that intrinsic paraglacial effects have been much more important than freeze-thaw cycling in promoting rockfall and talus accumulation in formerly glaciated areas, and that many supposedly 'periglacial' talus accumulations reflect a strong element of paraglacial inheritence.

17.3.4 Landforms and Deposits

Paraglacial rock-slope adjustment produces a wide range of landforms and deposits. Mountain slopes and summits are characterized by tension cracks and crevasse-like fractures, split summit ridges, toppled blocks, antiscarps and bulging slopes seamed with nested antiscarp arrays (Bovis, 1990; Chigira, 1992; Jarman and Ballantyne, 2002; Figs 17.5 and 17.6). Sliding failure along slope-parallel stress-release joints is evident in the formation of crescentic overhangs, and large-scale catastrophic rock-slope failures (rock avalanches and deep-seated rotational slides) produce deep arcuate failure scars. Large-scale catastrophic failures may accumulate as valley-side cones of large angular boulders, or extend well beyond the slope-foot as an excess-runout flowslide (Sturzstrom) deposit. Such excess-runout failures may extend several kilometres down valley floors (Evans and Clague, 1994), and are characterized by steep, well-defined lateral margins or levées of very coarse debris and concentric flow ridges. Paraglacial rockfall produces talus cones, coalescing talus cones and talus sheets along the lower slopes of cirque headwalls and glacial troughs. Such talus accumulations generally consist of smaller boulders than the products of large-scale catastrophic failure, and consist of clast-supported diamicts overlain by a surface layer of openwork boulders that exhibit a general increase in size downslope (fall-sorting). They often overlie valley-side morainic deposits and exhibit evidence for surface reworking by debris flows and snow avalanches. In permafrost environments, rockfall talus accumulations may terminate downslope in protalus (lobate) rock glaciers (Ballantyne, 2002b).

17.3.5 Wider Significance

The evidence outlined above suggests that paraglacial rock-slope adjustment represents an important if rarely acknowledged component of postglacial landscape evolution, particularly in alpine environments. It is worth noting two important implications. First, not only do paraglacial rock-slope failure, rock-slope deformation and rockfall alter the form of rockwalls during interglacials, but also rock-mass weakening due to paraglacial stress release may determine the foci of glacial erosion during later periods of glacial advance, raising the possibility that archetypal glacial landforms such as cirques and glacial troughs owe their present form as much to successive episodes of paraglacial adjustment as to successive periods of glacial quarrying and abrasion (Augustinus, 1995b). Second, because paraglacial rock-slope failures, slope deformations and talus accumulations provide an abundant source of readily entrainable debris, much of the sediment transported during the initial stages of a later glacial advance may ultimately be of paraglacial origin, rather than eroded from intact bedrock. If so, it is likely that the initial stages of renewed glaciation are marked by enhanced rates of glacial sediment transport; when glacial recycling of paraglacial debris is complete, glaciers transport only sediment supplied by 'normal' entrainment processes such as subglacial erosion and direct debris delivery from valley-side slopes. A full understanding of the sources of glacigenic sediment therefore requires not only appreciation of the mechanics of glacial erosion, but also the role of earlier paraglacial rock-slope adjustment in providing an abundant source of readily-entrainable debris (Ballantyne, 2002b). Just as paraglacial processes are by definition glacially-conditioned, so some aspects of glacial processes may to some extent reflect paraglacial conditioning or inheritance.

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