- Email: [email protected]
Controls on basalt terrace formation in the eastern Lesotho highlands
Geomorphology 67 (2005) 473 – 485 www.elsevier.com/locate/geomorph
Controls on basalt terrace formation in the eastern Lesotho highlands Stefan Grab*, Craig van Zyl, Nicholas Mulder School of Geography, Archaeology, and Environmental Studies, University of the Witwatersrand, P/Bag 3, Wits 2050, South Africa Received 3 April 2004; received in revised form 3 November 2004; accepted 3 November 2004 Available online 18 December 2004
Abstract The objective of this study is to ascertain the relative importance of lithological controls and geomorphological processes in the development of Drakensberg basalt terraces. Various hypotheses for terrace formation are considered, including geological controls, macroscale geomorphology, and climatic–geomorphological controls. The variations in strength and relative age differences for scarp surface exposures on two slopes of varying aspect are determined, so that a comparison can be made between various scarp outcrops and relative rates of weathering. Scarp outcrops were measured for their rock mass strength, rock surface roughness, and percentage lichen cover. The ethylene glycol test was performed on prepared rock samples to determine susceptibility to tensional breakup. A satellite image depicting the distribution of late-lying contemporary snow was used to assist in the construction of topo-climatic linkages with scarp terrace localities. Findings show that terraces are most common on southeast-facing aspects, and coincide with the major joint strike direction. Lithological factors are thus considered the primary control to such terrace development. It is suggested that a different set of geomorphological processes operates on various slope-altitudinal and slope-orientational positions at any given time. D 2004 Elsevier B.V. All rights reserved. Keywords: Roughness; Lichenometry; Schmidt hammer; Geological and climatic controls; Weathering; Cryoplanation
1. Introduction The study of terraced/stepped topography spans from microscale (cm/m) terracettes, bounded by small earth risers (e.g., Vincent and Clarke, 1976; Bielecki and Mueller, 2002), to macroscale (km) planation surfaces bounded by escarpments (e.g., King, 1976; Peulvast and Claudino Sales, 2004). This paper focuses * Corresponding author. Fax: +27 11 403 7281. E-mail address: [email protected] (S. Grab). 0169-555X/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.geomorph.2004.11.010
on rock scarps and associated terraces, which are of intermediate scale (up to 102 m). Terrace pediments are fundamentally erosional slope forms (Plakht et al., 2000), produced through one or more possible processes including fluvial (e.g., Veldkamp, 1992), glacial (e.g., Colman, 1976), periglacial (e.g., Czudek, 1995), and changing sea level (e.g., Cinque et al., 1995) origins. Terrace pediments have also been ascribed to basal weathering (detchplainsT) (e.g., Bu¨del, 1970) and geological processes such as slope displacement by normal faulting (e.g., Bentivenga et al., 2004).
474
S. Grab et al. / Geomorphology 67 (2005) 473–485
Slope profiles depicting alternating bands of scarps and terraces are typical phenomena to flood basalt outcrops such as those known in Iceland, the Antrim basalts (Northern Ireland), Golan Heights (Israel), Simien Mountains (Ethiopia), Western Ghats (India), and the high ranges of Lesotho. Despite the widespread and conspicuous basalt terraces throughout much of the eastern Lesotho highlands, these have received limited scientific study. The dbasalt terracesT are referred to as such, owing to the step-like appearance of basaltic scarp outcrops above terrace pediments that contour hillslopes. The earliest suggestion was that the Lesotho dbasalt stepsT originate from the Pleistocene cold stages when enhanced frost wedging and nivation processes dattackedT the cross-jointed basalt outcrops (Harper, 1969). Meiklejohn (1992) revisited these ideas, suggesting that the terraces resemble cryoplanation terraces and may be associated with cryogenic processes. An alternative suggestion is that the dbedrock ledgesT are the product of river incision, associated with the earliest erosional cycle on the southern African continent, the post-Gondwana planation (Van Rooy and Van Schalkwyk, 1993). The dstepped topographyT has also been attributed to variations in strength between individual basalt layers (Boelhouwers, 1994). Yet, Mitchell et al. (1996) describe the Lesotho basalt as geochemically relatively uniform. An issue that emerges from the brief and mostly descriptive literature is the need to identify the relative importance of geological origins and that of past and present geomorphological processes in such basalt terrace formation. A similar problem has emerged within the international dcryoplanationT literature (e.g., Czudek, 1995; Hall, 1997, 1998; Nelson, 1998), which is, as yet, also unresolved. A recent conclusion stemming from cryoplanation research suggests that future work should differentiate terraces of varying age and identify the relations of individual terrace locations with lithology, structural factors, snow accumulation patterns, and geomorphic processes (Nelson, 1998). Given that the terraces examined in this study morphologically resemble features typically referred to as dcryoplanation terraces,T we aim to follow-up on some of the suggestions made by Nelson (1998). Our main objective is to determine whether geological attributes or geomorphic processes (such as those associated with dcryoplanationT and fluvial erosion) are the primary factor for scarp terrace
development over time. Further, we investigate whether the terraces have a preferred occurrence on any given slope aspect, and, if so, whether this is geologically or climato-geomorphically controlled. Satellite images depicting the distribution of late-lying contemporary snow patches are used to assist in the discussion on topo-climatic linkages with scarp terrace orientation. Whilst it is not possible to accurately test the diverse and synergistic set of geomorphic processes that have operated since the origin of the Drakensberg basalt, an attempt is made to examine the variations in strength and relative surface age differences for basalt exposures along two slope profiles of varying altitudinal positions and aspect (north and west facing), so as to differentiate and debate relative rates of recent scarp weathering and recession.
2. Geological setting of the Lesotho Highlands The Drakensberg volcanic group is the product of a phase of discrete volcanic centers, which ejected material, followed by a phase of basaltic lava extrusions associated with dyke intrusions (Dingle et al., 1983). This later stage of mobile lava flows is estimated to have occurred about 182F2 Ma ago (Hooper et al., 1993). The Kraai River Formation, which typically displays columnar jointing and pillow lavas, has only been recognized in the eastern Cape regions (Dingle et al., 1983), while pillow lavas have been reported in the Indedema Valley in the Cathedral Peak region (McCarthy, 1970). The 800-m section of exposed basalt stratigraphy of the Sani Pass area has been divided into a lower 175-m section with four compositionally distinct stratigraphic units, and an overlying 625-m-thick succession of homogeneous Lesotho basalt (Mitchell et al., 1996). Each distinctive flow unit maintains its thickness for great distances (King, 1963) and averages 6 m (OSC, 1986), although the median value is 3.2 m (Galliers et al., 1991). Given that the terraces are found within the Lesotho basalt type, it may be assumed that the geochemistry for each terrace is similar (cf. Van Rooy and Nixon, 1990). The similarity in bulk chemical composition thus reduces the likelihood that chemical composition is an important determining factor in terrace development. Rather, the degree of alteration of primary minerals to secondary smectite minerals
S. Grab et al. / Geomorphology 67 (2005) 473–485
for each scarp should be used. The alteration of the basalt could play an important role in terrace development by affecting the depth of weathering (Van Rooy and Van Schalkwyk, 1993). An important observation that should be considered in the formation of basalt terraces is the exposure of an amygdaloidal zone at the base of each scarp, and in some cases near the top as well. A distinct zonation can be identified for each lava flow: a very thin aphanitic basalt at the base, followed by pipe amygdales or infilled vesicles, and then a massive central zone of weakly vesicular basalt and an upper layer of strongly vesicular basalt (cf. King, 1963; Nixon, 1973; Van Rooy and Van Schalkwyk, 1993) (Fig. 1). This zonation was observed in almost all scarp faces investigated, suggesting that each scarp face represents a single lava flow.
3. Methods Two terraced slope transects (at bearings 2808 and 3508) were investigated near Sani Pass, eastern Lesotho (Fig. 2). At this locality, terraces do not occur between bearings 508 and 2708. Slopes extend from 2960 to 3060 m asl, over which six distinct scarp outcrop (labeled as A–F) and associated terrace
475
features are identified (Fig. 3). The height of scarp outcrops varies between 3 and 8 m, whilst terrace tread diameters are between 29 and 126 m. A classification for assessing the strength of intact bedrock was first devised by Selby (1980) and later refined by Moon (1984). Typically, seven strength parameters are used to test rock mass strength (RMS) on particular slope gradients for the determination of strength equilibrium on slopes (Selby, 1982a; Moon and Selby, 1983). The RMS technique has been employed on a variety of rock lithologies including basalt, dolerite, quartzite, conglomerates, and shale (cf. Selby, 1980; Moon, 1991). For parallel scarp retreat to occur, a uniform RMS would be expected throughout the profile (Selby, 1982a,b). Near-vertical scarp faces of the Drakensberg basalt are apparently in equilibrium with contemporary processes, thus suggesting that current slope retreat is not dependent on rapid or catastrophic mass movement events (Moon and Selby, 1983). Perhaps, the most important parameter of the RMS classification is that of intact strength (hardness), which is measured using a Schmidt hammer (Day, 1980; McCarroll, 1991). The Schmidt hammer provides a rebound reading (R value ranging from 0 to 100), which is proportional to the hardness of the rock. The Schmidt hammer has been used for a number of rock weathering studies, including the role of late-lying
Fig. 1. Typical basalt structural zonation along a scarp face. Camera lens cap diameter=6 cm.
476
S. Grab et al. / Geomorphology 67 (2005) 473–485
Fig. 2. The site locality of the north-facing and west-facing terraced slopes in the Sani Pass area in eastern Lesotho. The bold line indicates terraces that parallel the course of the Pitsaneng River to the south of the map.
snow patches on bedrock weathering (Hall, 1993), its application to relative age dating (e.g., McCarroll, 1989; Nicholas and Butler, 1996; Sumner et al., 2002), and its effect on tafoni weathering (Matsukura and Tanaka, 2000). For the purpose of our study, an N-type
Schmidt hammer was used to assess the compressive strength characteristics (R values) per scarp face, so as to help establish relative rock surface age differences between the various scarps. Impact sites were cleaned using a carborundum and mean values were calculated
S. Grab et al. / Geomorphology 67 (2005) 473–485
477
Fig. 3. Sketch profiles indicating the position of scarp faces on the north-facing and west-facing slopes.
from 100 readings per site. In addition, the overall RMS rating for individual scarp outcrops was determined according to the various lithological parameters specified by Selby (1980). Rock surface roughness was measured to ascertain the relative difference in extent of surface weathering between individual scarps and to identify possible relations with R values. The roughness profile was recorded by using a profile gauge and calculated as a roughness index according to methods used by Nesje et al. (1994) and McCarroll and Nesje (1996). The use of lichenometry is well established as a relative age-dating technique (cf. Innes, 1985a,b; Nicholas and Butler, 1996), despite considerable geoecological complicating factors (Midgley, 1992). Measuring percentage lichen cover on a scarp face, as a relative age indicator for the exposed rock surface, is more suitable than measuring size–age or shape–age relationships (Innes, 1988). The underlying assumption is that the older, longer-exposed scarp face will have a higher percentage of lichen cover than a surface that has either been subjected to more rapid weathering and erosion, or only recently been exposed. We thus test this hypothesis against the other variables measured for each scarp outcrop. Tracing the lichen on the rock surface and transferring these onto graph paper obtained percentage lichen cover. Lichen were measured along 15-m belt transects per scarp face. Rock properties need to be considered in developing an understanding of cliff behavior (Douglas et al., 1991). Rock samples obtained from the field were cut into a standard size of 10–15 mm thickness, with a
volume/surface area ratio of between 4 and 5. An angle grinder fitted with a cement-cutting disk was used for the ethylene glycol test. Such tests have been used to determine the durability of Drakensberg basalts in geotechnical studies for the Lesotho Highlands Water Project (OSC, 1986; Van Rooy and Nixon, 1990; Van Rooy and Van Schalkwyk, 1993). The basalt specimens are immersed in undiluted ethylene glycol, which is an organic polar liquid. The alteration products of the basalts include montmorillonite, nontronite, and sapronite, with intergrowths of these clay minerals occurring and collectively accounting for ca. 30–40% of the total rock mineralogy (Van Rooy and Nixon, 1990). An important characteristic of these smectite group minerals is their capacity to absorb water molecules between the sheets, thus producing considerable expansion of the structure (Klein and Hurlbut, 1999). It is commonly found that the expansion of these clays upon wetting causes tensional breakup of the rock (Brekke, 1965; Douglas et al., 1991). This effect can be simulated by means of wetting and drying cycles (McGreevy, 1982; Hall and Hall, 1996), although it requires a large number of cycles before reliable results are obtained. The ethylene glycol test offers a simpler and quicker method, with results obtained within 30 days (cf. Van Rooy and Nixon, 1990) owing to the ethylene glycol substituting water molecules in the clay lattice to cause swelling (Gillespie et al., 1994). Such tests have been cautioned, as they are severe and do not necessarily equate to an accurate simulation of natural rock behaviour (Davies, 1989). However, for the purpose of this study, the ethylene glycol test results
478
S. Grab et al. / Geomorphology 67 (2005) 473–485
Fig. 4. Sketches indicating various components investigated in the rock samples following the ethylene glycol test.
classification of a composite of TM bands 3–5 was performed.
are used simply as a source of comparison between various scarp outcrops and considered together with the results from other tests performed. To quantify the relative changes, the samples were observed under 2 magnification and sketches drawn before immersion, after 15 days and again after 30 days of immersion (Fig. 4). Any changes in fractures and swelling of clays are then easily identified, and a qualitative method was employed for the characterization of the detected changes (Table 1). The terrace orientation frequency distribution was determined for an area extending 3 km west of the Great Escarpment and between 29820VS and 29830VS. Scarp tread orientation was measured as the outwardfacing normal-to-scarp lineation, as suggested by Nelson (1998). The same region was assessed for late-lying snow patch orientation frequency distribution using a Landsat 5 satellite image dated August 3, 1990. The previous cold front had passed through 5 days earlier, and so the image depicts snow patches that had survived at least 5 days. A fine unsupervised
4. Cycles of erosion and implications for terrace formation Most hypotheses for slope terrace formation have failed to consider the initial valley-forming processes. Many of the high-lying bevels in Lesotho are classified according to the first cycle of erosion (the Gondwana planation cycle) after the cessation of the lava outpourings, which was followed by the rifting of the Gondwanaland assemblage during the late Jurassic. This initiated the post-Gondwana cycle of erosion, which lies ca. 330 m below the Gondwana summit level (King, 1963). According to Partridge and Maud (1987), the Great Escarpment had receded by ca. 100 km by the mid-Cretaceous, and a further 20 km by the end of the Cretaceous. The Post-African I and II cycles had limited erosive impact on the Lesotho Highlands,
Table 1 Characterization and differentiation of rock alterations after 30 days of immersion in undiluted ethylene glycol (adopted from OSC, 1986) Swelling of clay spots Fracturing along existing joints Small fractures throughout Extensive fracturing Complete breakdown North-facing transect
=visible expansion of clay minerals on the surface =relative to unnaturally induced fractures =especially if radiating from amygdales and expanded clay spots =a high continuity of fractures =crumbling or flaking of rock fragments Scarp locality A
Swelling of clay spots Fracturing along existing joints Small fractures throughout Extensive fracturing Complete breakdown
x x x
West-facing transect
Scarp locality
B
C
D
E
F
A
B
C
D
E
F
x
x x x
x x
x x
x
x x x x
x
x x
x x
x x x
x x
x
x
x
S. Grab et al. / Geomorphology 67 (2005) 473–485
resulting in different surface expressions for the interior, compared to those for coastal regions (Nixon, 1973; Partridge and Maud, 1987). Continued continental margin uplift caused baselevel readjustments with enhanced river incision, supporting the suggestion that the Lesotho terraces are the consequence of fluvial action during erosional cycles (Van Rooy and Van Schalkwyk, 1993). An example of this is found along the Pitsaneng River where the terraces parallel the course of the thalweg (Fig. 2). It has also been suggested that the Lesotho slope hollows have been truncated by African cycle river incision (Dyer and Marker, 1979). However, it is evident from aerial photographic interpretation and ground truthing that the terraces continue unabated across these hollows. Therefore, the terraces cannot be attributed to river erosion or a period of erosion represented by the African cycle, which truncates all erosive cycles. This implies that the terraces do not owe their origin to an erosional cycle.
5. Geomorphological controls It has been suggested that the basalt terraces are the product of cryoplanation processes (Meiklejohn, 1992). Features generally known as cryoplanation terraces have also previously been referred to as dgoletzT terraces, altiplanation terraces, nivation terraces, and equiplanation terraces (see French and Harry, 1992, p. 146). Reger and Pe´ we´ (1976) define cryoplanation terraces as dminor step-like or table-like landforms consisting of nearly horizontal bedrock surfaces covered by a thin veneer of rock debris and bounded by ascending or descending bedrock scarps or bothT (p99). Cryoplanation terraces often occur din a tiered series culminating in a summit flat, but rarely encircle an entire mountain or hill continuouslyT (Nelson, 1989, p. 31). The Lesotho terraces perfectly fit the morphological description of cryoplanation terraces given here. In addition, certain rocks such as basalt are said to be particularly suitable for cryoplanation processes owing to suitable lithological attributes (Priesnitz, 1988). The central tenet to cryoplanation entails the occurrence of long-lasting snow banks against the scarp face, providing moisture for frost wedging within the exposed scarp face and subsequent solifluction (removal of weathered prod-
479
ucts) over the terrace tread (Demek, 1969; Reger and Pe´we´, 1976; Priesnitz, 1988; Czudek, 1995; Grosso and Corte, 1991; Thorn and Hall, 2002), all of which occur on the south-facing slope aspects of eastern Lesotho during contemporary winters (Fig. 5). However, the concept of dcryoplanation,T implying a particular dformT and dgenesisT (Thorn and Hall, 2002), remains subjected to critical debate. Thorn and Hall (2002) conclude that dto deny the fairly widespread existence of features commonly called cryoplanation benches, terraces, and pediments would be foolish: to claim that there is anything approaching an adequate explanation of their origin(s) would be even more foolishT (p. 548). There is also uncertainty regarding the genetic differences, if any, between cryoplanation benches, terraces, and pediments (Thorn and Hall, 2002). Whilst we refer to the Lesotho features as terraces, we do not imply any particular process origin. Periglacial research in the high Drakensberg has produced evidence for both contemporary frost wedging along scarp faces and solifluction over terrace treads (Grab, 1999, 2000). However, the role of freeze– thaw as a weathering process in landform development has raised more questions than answers (e.g., Thorn, 1979; Hall, 1991, 1995; Matsuoka, 1990, 2001). It is generally accepted that even if cryogenic processes are active, terrace formation frequently occurs along structural benches (French and Harry, 1992), or along zones of high joint density (Czudek, 1995). Work in Antarctica has shown that bench initiation may arise from a preferential suite of processes at sites of extensive dilatation jointing (Hall, 1997). The bench
Fig. 5. During colder months, long-lasting snow banks occur below south-facing scarp faces. Solifluction mantles are frequently visible on south-facing terrace treads.
480
S. Grab et al. / Geomorphology 67 (2005) 473–485
forms are extended and maintained through ongoing weathering, mass wasting, and fluvial action (Hall, 1997). The preferential occurrence of Lesotho terraces
Fig. 7. Wedges of ice protruding out of a highly jointed south-facing scarp face (top photograph). The lower photograph shows a typical terrace scarp and adjoining tread. Rockfall debris mantles sometimes occur below scarps where there is a high joint density and ground seepage.
Fig. 6. A comparison of basalt terrace frequency distribution with snow-covered area for various slope aspects (orientations). The test area extends 3 km west of the Great Escarpment, between latitudes 29820 VS and 29830 VS. Snow cover data were obtained from a satellite image dated August 3, 1990.
between 808 and 2008 coincides with about 89% of Lesotho basalt joints striking between 808 and 1408 (Nixon, 1973) (Fig. 6). Joints dipping out of terrace free faces are commonly associated with increased water seepage and sapping at the scarp base. The highly jointed and weathered free faces may eventually fail owing to high pore water pressure (summer) or pressures induced by wedges of ice (winter), which should not only be identified as an agent of weathering, but also as an erosion mechanism (Fig. 7). The highly angular rock fall debris is a product of jointing sets, which commonly have a 608 intersection angle. Data indicate that the highest snow covered area is between bearings 1408 and 2008, with a peak at 1808 (Fig. 6). The longevity of snow patches is limited on north-facing aspects due to the high angle of incidence of insolation. In contrast, the higher-frequency distribution of basalt terraces on south-facing slopes overlaps with a higher-frequency distribution of longer-
S. Grab et al. / Geomorphology 67 (2005) 473–485
Fig. 8. Rock mass strength results for the surveyed north-facing and west-facing scarps in the Sani Pass area. The strength equilibrium envelope is after Selby (1980).
481
slopes absent of scarp outcrops. It is therefore slope orientation, rather than the presence of scarp outcrops, that is the primary control determining the distribution of longer-lasting snow patches. The occurrence of wedges of ice within scarp joints, long-lasting snow banks at the scarp terrace knickpoint, and evidence of solifluction on the terraces indicates that geomorphic attributes and processes commonly associated with dcryoplanationT are present at sites in eastern Lesotho. The extent to which past and contemporary cryogenic processes are contributing to the extension and maintenance of the scarp terrace morphology remains unclear.
6. Results and discussion lasting snow patches (Fig. 6). Scarp outcrops promote shadow and consequential cooling effects at the knickpoint with terrace treads, thus assisting in snow patch preservation. However, the trend in snow cover distribution displayed in Fig. 6 is also characteristic for
Fig. 9. Distribution of rock hardness (Schmidt hammer rebound) values for the surveyed north-facing and west-facing scarps in the Sani Pass area (n=100 per scarp face).
Most scarp faces fall outside the strength equilibrium envelope, displaying high slope angles relative to their mass strength (Fig. 8). This may be attributed to the scarp outcrops being undercut at the point of contact with terrace treads. Such undercutting has been attributed to the headward extension of talus and regolith-covered slopes below the scarps (Moon and Selby, 1983), thus undermining the lowest joint blocks and leading to high-angle outcrops. RMS increases from the upper to the lower scarp outcrops (Fig. 8), mainly influenced by increasing joint spacing and higher rock hardness (rebound) values in the downslope direction (Fig. 9). Similarly, northfacing scarps have lower mass strengths than corresponding west-facing scarps. The west-facing scarps reflect somewhat higher rock hardness values than corresponding north-facing scarps (38 vs. 33; n=600 per transect) (Figs. 8 and 9). It could thus be argued that the extent of weathering or chemical alteration at the surface generally increases upslope and from west to north. The ethylene glycol test results support the downslope trend in R values (Table 1). Although there is little difference between the north-facing and westfacing rock surface results, there is a marked difference between the results for the highest (A) and lowest (F) scarp outcrops (Table 1). Samples from the uppermost scarp outcrop underwent extensive fracturing and breakdown, whilst those from the lowermost outcrop only show swelling of clay spots and fracturing along existing joints. Intermediate scarp
482
S. Grab et al. / Geomorphology 67 (2005) 473–485
outcrops (B–E) show variable alteration and R values. It is thus apparent that as the chemical alteration of the rock mass increases, more secondary minerals are produced and R values therefore decrease. These trends demonstrate that the rock surfaces near the summit scarp have been exposed for much longer than those near the base of the slope. The rock surface roughness profile results (Fig. 10) indicate a progressive decrease in roughness in the downslope direction, which confirms the trends in different surface weathering expressions discussed. Rock surface roughness may provide a relative indication of the length of surface exposure (McCarroll and Nesje, 1996) and thus confirms that the length of rock surface exposure generally increases from lower to upper scarp outcrops. The north-facing scarps have a higher mean roughness index than the corresponding west-facing scarps, which may be attributed to higher chemical alteration (weathering) on the north-facing slope (Fig. 10). Although no clear trend in lichen cover is noticeable on the west-facing slope, there is a general increase in coverage from the upper (17%)
Fig. 11. Percentage lichen cover on the surveyed north-facing and west-facing scarps in the Sani Pass area.
to the lower (37%) scarp outcrops on the northfacing slope (Fig. 11). Variations in microclimate, solar radiation, rock mineralogy, and weathering would complicate the trends displayed. It is nevertheless interesting to note that the general trend in lichen coverage on the north-facing slope closely resembles that for RMS and R values (Figs. 8 and 9). It appears that lichen coverage decreases where rocks are undergoing greater chemical alteration and more intense weathering, and therefore does not offer a suitable rock surface relative age indication in this region. Rock surfaces with enhanced lichen growth may sustain greater mass strength and granular stability, as lichen are known to insulate rock and intercept moisture that would otherwise contribute to weathering (Benedict, 1993). The relative contribution of biologically induced (e.g., lichen) rock weathering processes to basalt disintegration is still unknown to this region.
7. Conclusions
Fig. 10. Different scales of variance of rock surface roughness for the surveyed north-facing and west-facing scarps in the Sani Pass area (n=50 profiles per scarp face).
Although some of the Lesotho terraces may have been accentuated through phases of fluvial incision, they do not owe their origin to the so-called dcycles of erosion.T Geological properties of individual scarp outcrops forming distinct slope terraces in eastern Lesotho correspond with individual lava flows. Despite terraces occurring on all slope aspects, they are most common on southeast-facing positions, thus coinciding with the majority of joint strike directions. Lithological factors are therefore considered the primary control to such terrace development.
S. Grab et al. / Geomorphology 67 (2005) 473–485
Given that the dominant aspect position of late-lying snow packs and occurrence of ice in scarp joints is similar to the major joint strike directions, it is difficult to advocate topo-climatic geomorphological factors as the primary control for terrace development. Nevertheless, given that contemporary snow patches, ice, and frozen ground last much longer (several weeks) on south-facing slopes than on other aspects, it is likely that processes associated with cryoplanation have contributed to the modification and extension of south-facing basalt terraces in eastern Lesotho. Terraces located on north-facing aspects are less likely to be affected by cryogenic processes, but are subjected to higher temperatures and greater thermal and moisture fluxes than southfacing slopes (personal measurements). Scarp outcrops are likely to be impacted by a variety of mechanical, chemical, and biological processes, including ground seepage through fractures. Whilst it can be recognized that the suite of processes extending basalt terraces varies with aspect, it is still uncertain which set of contemporary processes operating on a particular slope aspect is most effective in extending the terraces. Results have shown a general increase in weathering extent and decrease in RMS and R values from west to north, and from lower to higher scarp outcrop localities. It would appear that the higher scarp outcrop localities, most particularly on northfacing aspects, display longer exposed rock surfaces than those at lower elevations along a given slope profile. The higher scarp outcrop surfaces demonstrate exfoliation and granular disintegration, whilst the smoother, less weathered surfaces of lower scarp outcrops may be associated with a more active slope retreat. Such lower scarp outcrops display considerable rockfall debris mantles at zones of ground seepage (Fig. 7). The upper north-facing scarp outcrops are less lichen-covered than on lower slope outcrops, which may be owing to increased granular disintegration and exfoliation, and greater exposure to insolation and desiccation as one advances altitudinally. It is most probable that a different set of geomorphological processes operates on various slopealtitudinal and slope-orientational positions at any given time. Small variations in lithology and geomorphological processes may account for differences
483
in weathering and mass movement surface expressions between the various scarp outcrops and slope terraces. The suite of contemporary geomorphological processes extending the dbasalt terracesT is both spatially and temporally complex, and requires further study.
Acknowledgements We appreciate the financial assistance provided by the Wits University Research Committee Grant and the IAG Grant to attend the IAG International Symposium in Addis Ababa, December 2002. Wendy Job is thanked for her assistance in producing the figures. Our sincere thanks to the two referees who provided valuable input and helped improve an earlier version of the paper. This paper has benefited from valuable discussions and correspondence with Cliff Ollier.
References Benedict, J.B., 1993. Influence of snow upon rates of granodiorite weathering, Colorado Front Range, USA. Boreas 22, 87 – 92. Bentivenga, M., Coltorti, M., Prosser, G., Tavarnelli, E., 2004. A new interpretation of terraces in the Taranto Gulf: the role of extensional faulting. Geomorphology 60, 383 – 402. Bielecki, A.E., Mueller, K.J., 2002. Origin of terraced hillslopes on active folds in the southern San Joaquin Valley, California. Geomorphology 42, 131 – 152. Boelhouwers, J.C., 1994. Periglacial landforms at Giant’s Castle, Natal Drakensberg, South Africa. Permafr. Periglac. Process. 5, 129 – 136. Brekke, T.L., 1965. On the measurement of the relative, potential swellability of hydrothermal montmorillonite clay from joints and faults in Pre-Cambrian and Paleozoic rocks in Norway. Int. J. Rock Mech. Min. Sci. 2, 155 – 165. Bqdel, J., 1970. Pedimente, Rumpffl7chen und RqcklandSteilh7nge; deren active und passive Ruckverlegung in verschiedenen Klimaten. Z. Geomorphol. 14, 1 – 57. Cinque, A., De Pippo, T., Romano, P., 1995. Coastal slope terracing and relative sea-level changes: deductions based on computer simulations. Earth Surf. Processes and Landf. 20, 87 – 103. Colman, S.M., 1976. Inherent factors in the flow of valley glaciers as a possible influence in the formation of stepped glacial valleys. Z. Geomorphol. 20, 297 – 307. Czudek, T., 1995. Cryoplanation terraces—a brief review and some remarks. Geogr. Ann. 77A, 95 – 105. Davies, G., 1989. Materials testing for Lesotho water scheme. S. Afr. J. Sci. 85, 483.
484
S. Grab et al. / Geomorphology 67 (2005) 473–485
Day, M.J., 1980. Rock hardness: field assessment of geomorphic importance. Prof. Geogr. 32, 72 – 81. Demek, J., 1969. Cryogene processes and the development of cryoplanation terraces. Biul. Peryglac. 18, 115 – 125. Dingle, R.V., Siesser, W.G., Newton, A.R., 1983. Mesozoic and Tertiary Geology of Southern Africa. Balkema, Rotterdam. Douglas, G.R., Whalley, W.B., McGreevy, J.P., 1991. Rock properties as controls on freeface debris fall activity. Permafr. Periglac. Process. 2, 311 – 319. Dyer, T.G.J., Marker, M.E., 1979. On some aspects of Lesotho hollows. Z. Geomorphol. 23, 256 – 270. French, H.M., Harry, D.G., 1992. Pediments and cold-climate conditions, Barn Mountains, unglaciated northern Yukon, Canada. Geogr. Ann. 74A, 145 – 157. Galliers, R.M., Brackley, I.J.A., Lephoma, M., Sympson, A.B., 1991. Expected engineering geology of the transfer tunnel. Proceedings of the 10th Regional Conference for Africa on Soil Mechanics and Foundational Engineering, Maseru, Lesotho, pp. 351 – 358. Gillespie, R.J., Eaton, D.R., Humphreys, D.A., Robinson, E.A., 1994. Atoms, Molecules and Reactions. Prentice Hall, New Jersey. Grab, S.W., 1999. Block and debris deposits in the high Drakensberg, Lesotho, southern Africa: implications for high altitude slope processes. Geogr. Ann. 81A, 1 – 16. Grab, S.W., 2000. Stone-banked lobes and environmental implications, high Drakensberg, southern Africa. Permafr. Periglac. Process. 11, 177 – 187. Grosso, S.A., Corte, A.E., 1991. Cryoplanation surfaces in the Central Andes at latitude 35 degrees S. Permafr. Periglac. Process. 2, 49 – 58. Hall, K., 1991. The allocation of the freeze–thaw weathering mechanism in geocryological studies: a critical comment. S. Afr. Geogr. J. 73, 10 – 13. Hall, K., 1993. Enhanced bedrock weathering in association with late-lying snowpatches: evidence from Livingston Island, Antarctica. Earth Surf. Processes and Landf. 18, 121 – 129. Hall, K., 1995. Freeze–thaw weathering: the cold region bPanaceaQ. Polar Geogr. Geol. 19, 79 – 87. Hall, K., 1997. Observations on bcryoplanationQ benches in Antarctica. Antarct. Sci. 9, 181 – 187. Hall, K., 1998. Nivation or cryoplanation: different terms, same features? Polar Geogr. 22, 1 – 16. Hall, K., Hall, A., 1996. Weathering by wetting and drying: some experimental results. Earth Surf. Processes and Landf. 21, 365 – 376. Harper, G., 1969. Periglacial evidence in southern Africa during the Pleistocene epoch. Palaeoecol. Afr. Surround. Isl. Antarct. 4, 71 – 101. Hooper, P.R., Duncan, R.A., Marsh, J.S., Duncan, A.R., 1993. The basalts of Lesotho, Karoo Province, southern Africa. Eos 74, 553. Innes, J.L., 1985a. Lichenometric dating of debris-flow deposits on alpine colluvial fans in southwest Norway. Earth Surf. Processes and Landf. 10, 519 – 528. Innes, J.L., 1985b. Lichenometry. Prog. Phys. Geogr. 9, 187 – 254. Innes, J.L., 1988. The use of lichens in dating. In: Galun, M. (Ed.), CRC Handbook of Lichenology. CRC Press, Boca Raton.
King, L.C., 1963. South African Scenery, 3rd ed. Oliver and Boyd, London. King, L.C., 1976. Planation remnants upon high lands. Z. Geomorphol. 20, 133 – 148. Klein, C., Hurlbut, C.S., 1999. Manual of Mineralogy. John Wiley & Sons, New York. Matsukura, Y., Tanaka, Y., 2000. Effect of rock hardness and moisture content on tafoni weathering in the granite of Mount Doeg-sung, Korea. Geogr. Ann. 82A, 59 – 67. Matsuoka, N., 1990. The rate of bedrock weathering by frost action: field measurements and a predictive model. Earth Surf. Processes and Landf. 15, 73 – 90. Matsuoka, N., 2001. Direct observation of frost wedging in alpine bedrock. Earth Surf. Processes and Landf. 26, 601 – 614. McCarroll, D., 1989. Potential and limitations of the Schmidt hammer for relative-age dating: field tests on Neoglacial moraines, Jotunheimen, southern Norway. Arct. Alp. Res. 21, 268 – 275. McCarroll, D., 1991. The Schmidt hammer, weathering and rock surface roughness. Earth Surf. Processes and Landf. 16, 477 – 480. McCarroll, D., Nesje, A., 1996. Rock surface roughness as an indicator of degree of rock surface weathering. Earth Surf. Processes and Landf. 21, 963 – 977. McCarthy, M.J., 1970. An occurrence of pillow lava in a basal flow of the Drakensberg volcanic stage, Ndedema Valley, Natal. Second Gondwana Symposium: Proceedings and Papers. Council for Scientific and Industrial Research, Pretoria, pp. 433 – 439. McGreevy, J.P., 1982. Hydrothermal alteration and earth surface rock weathering: a basalt example. Earth Surf. Processes and Landf. 7, 189 – 195. Meiklejohn, K.I., 1992. Some aspects of valley asymmetry in the high Drakensberg. S. Afr. Geogr. J. 74, 49 – 53. Midgley, R., 1992. The Visual Dictionary of Plants. Dorling Kindersley, London. Mitchell, A.A., Ramluckan, V.R., Dunlevey, J.N., Eglington, B.M., 1996. The basalt stratigraphy of the Sani Pass, KwaZulu/Natal Drakensberg. S. Afr. J. Geol. 99, 251 – 262. Moon, B.P., 1984. Refinement of a technique for determining rock mass strength for geomorphological purposes. Earth Surf. Processes and Landf. 9, 189 – 193. Moon, B.P., 1991. The significance of rectilinear bedrock slopes. S. Afr. J. Sci. 87, 208 – 210. Moon, B.P., Selby, M.J., 1983. Rock mass strength and scarp forms in southern Africa. Geogr. Ann. 65A, 135 – 145. Nelson, F.E., 1989. Cryoplanation terraces: periglacial cirque analogs. Geogr. Ann. 71A, 31 – 41. Nelson, F.E., 1998. Cryoplanation terrace orientation in Alaska. Geogr. Ann. 80A, 135 – 151. Nesje, A., McCarroll, D., Dahl, S.-O., 1994. Degree of rock surface weathering as an indicator of ice-sheet thickness along an east– west transect across southern Norway. J. Quat. Sci. 9, 337 – 348. Nicholas, J.W., Butler, D.R., 1996. Application of relative-age dating techniques on rock glaciers of the La Sal Mountains, Utah: an interpretation of Holocene palaeoclimates. Geogr. Ann. 78A, 1 – 18.
S. Grab et al. / Geomorphology 67 (2005) 473–485 Nixon, P.H., 1973. Lesotho Kimberlites. Lesotho National Development, Maseru. Olivier Shand and Lahmeyer MacDonald Consortiums (OSC), 1986. Lesotho Highlands Water Project, Feasibility Study, April 1986. Partridge, T.C., Maud, R.R., 1987. Geomorphic evolution of southern Africa since the Mesozoic. S. Afr. J. Geol. 90, 179 – 208. Peulvast, J.-P., Claudino Sales, V., 2004. Stepped surfaces and palaeolandforms in the northern Brazilian bNordesteQ: constraints on models of morphotectonic evolution. Geomorphology 62, 89 – 122. Plakht, J., Patyk-Kara, N., Gorelikova, N., 2000. Terrace pediments in Makhtesh Ramon, central Negev, Israel. Earth Surf. Processes and Landf. 25, 29 – 39. Priesnitz, K., 1988. Cryoplanation. In: Clark, M.J. (Ed.), Advances in Periglacial Geomorphology. Wiley, Chichester, pp. 49 – 67. Reger, R.D., Pe´we´, T.L., 1976. Cryoplanation terraces: indicators of a permafrost environment. Quat. Res. 6, 99 – 109. Selby, M.J., 1980. A rock mass strength classification for geomorphic purposes: with tests from Antarctica and New Zealand. Z. Geomorphol. 24, 31 – 51. Selby, M.J., 1982a. Controls on the stability and inclinations of hillslopes formed on hard rock. Earth Surf. Processes and Landf. 7, 449 – 467.
485
Selby, M.J., 1982b. Rock mass strength and the form of inselbergs in the central Namib Desert. Earth Surf. Processes and Landf. 7, 489 – 497. Sumner, P., Nel, W., Holness, S., Boelhouwers, J., 2002. Rock weathering characteristics as relative-age indicators for glacial and post-glacial landforms on Marion Island. S. Afr. Geogr. J. 84, 153 – 157. Thorn, C.E., 1979. Bedrock freeze thaw weathering regime in an alpine environment, Colorado Front Range. Earth Surf. Processes and Landf. 4, 211 – 228. Thorn, C.E., Hall, K., 2002. Nivation and cryoplanation: the case for scrutiny and integration. Prog. Phys. Geogr. 26, 533 – 550. Van Rooy, J.L., Nixon, N., 1990. Mineralogical alteration and durability of Drakensberg basalts. S. Afr. J. Geol. 93, 729 – 737. Van Rooy, J.L., Van Schalkwyk, A., 1993. The geology of the Lesotho Highlands Water Project with special reference to the durability of construction materials. J. Afr. Earth Sci. 16, 181 – 192. Veldkamp, A., 1992. A 3-D model of Quaternary terrace development, simulations of terrace stratigraphy and valley asymmetry: a case study for the Allier terraces (Limagne, France). Earth Surf. Processes and Landf. 17, 487 – 500. Vincent, P.J., Clarke, J.V., 1976. The terracette enigma—a review. Biul. Peryglac. 25, 65 – 77.
Controls on basalt terrace formation in the eastern Lesotho highlands Stefan Grab*, Craig van Zyl, Nicholas Mulder School of Geography, Archaeology, and Environmental Studies, University of the Witwatersrand, P/Bag 3, Wits 2050, South Africa Received 3 April 2004; received in revised form 3 November 2004; accepted 3 November 2004 Available online 18 December 2004
Abstract The objective of this study is to ascertain the relative importance of lithological controls and geomorphological processes in the development of Drakensberg basalt terraces. Various hypotheses for terrace formation are considered, including geological controls, macroscale geomorphology, and climatic–geomorphological controls. The variations in strength and relative age differences for scarp surface exposures on two slopes of varying aspect are determined, so that a comparison can be made between various scarp outcrops and relative rates of weathering. Scarp outcrops were measured for their rock mass strength, rock surface roughness, and percentage lichen cover. The ethylene glycol test was performed on prepared rock samples to determine susceptibility to tensional breakup. A satellite image depicting the distribution of late-lying contemporary snow was used to assist in the construction of topo-climatic linkages with scarp terrace localities. Findings show that terraces are most common on southeast-facing aspects, and coincide with the major joint strike direction. Lithological factors are thus considered the primary control to such terrace development. It is suggested that a different set of geomorphological processes operates on various slope-altitudinal and slope-orientational positions at any given time. D 2004 Elsevier B.V. All rights reserved. Keywords: Roughness; Lichenometry; Schmidt hammer; Geological and climatic controls; Weathering; Cryoplanation
1. Introduction The study of terraced/stepped topography spans from microscale (cm/m) terracettes, bounded by small earth risers (e.g., Vincent and Clarke, 1976; Bielecki and Mueller, 2002), to macroscale (km) planation surfaces bounded by escarpments (e.g., King, 1976; Peulvast and Claudino Sales, 2004). This paper focuses * Corresponding author. Fax: +27 11 403 7281. E-mail address: [email protected] (S. Grab). 0169-555X/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.geomorph.2004.11.010
on rock scarps and associated terraces, which are of intermediate scale (up to 102 m). Terrace pediments are fundamentally erosional slope forms (Plakht et al., 2000), produced through one or more possible processes including fluvial (e.g., Veldkamp, 1992), glacial (e.g., Colman, 1976), periglacial (e.g., Czudek, 1995), and changing sea level (e.g., Cinque et al., 1995) origins. Terrace pediments have also been ascribed to basal weathering (detchplainsT) (e.g., Bu¨del, 1970) and geological processes such as slope displacement by normal faulting (e.g., Bentivenga et al., 2004).
474
S. Grab et al. / Geomorphology 67 (2005) 473–485
Slope profiles depicting alternating bands of scarps and terraces are typical phenomena to flood basalt outcrops such as those known in Iceland, the Antrim basalts (Northern Ireland), Golan Heights (Israel), Simien Mountains (Ethiopia), Western Ghats (India), and the high ranges of Lesotho. Despite the widespread and conspicuous basalt terraces throughout much of the eastern Lesotho highlands, these have received limited scientific study. The dbasalt terracesT are referred to as such, owing to the step-like appearance of basaltic scarp outcrops above terrace pediments that contour hillslopes. The earliest suggestion was that the Lesotho dbasalt stepsT originate from the Pleistocene cold stages when enhanced frost wedging and nivation processes dattackedT the cross-jointed basalt outcrops (Harper, 1969). Meiklejohn (1992) revisited these ideas, suggesting that the terraces resemble cryoplanation terraces and may be associated with cryogenic processes. An alternative suggestion is that the dbedrock ledgesT are the product of river incision, associated with the earliest erosional cycle on the southern African continent, the post-Gondwana planation (Van Rooy and Van Schalkwyk, 1993). The dstepped topographyT has also been attributed to variations in strength between individual basalt layers (Boelhouwers, 1994). Yet, Mitchell et al. (1996) describe the Lesotho basalt as geochemically relatively uniform. An issue that emerges from the brief and mostly descriptive literature is the need to identify the relative importance of geological origins and that of past and present geomorphological processes in such basalt terrace formation. A similar problem has emerged within the international dcryoplanationT literature (e.g., Czudek, 1995; Hall, 1997, 1998; Nelson, 1998), which is, as yet, also unresolved. A recent conclusion stemming from cryoplanation research suggests that future work should differentiate terraces of varying age and identify the relations of individual terrace locations with lithology, structural factors, snow accumulation patterns, and geomorphic processes (Nelson, 1998). Given that the terraces examined in this study morphologically resemble features typically referred to as dcryoplanation terraces,T we aim to follow-up on some of the suggestions made by Nelson (1998). Our main objective is to determine whether geological attributes or geomorphic processes (such as those associated with dcryoplanationT and fluvial erosion) are the primary factor for scarp terrace
development over time. Further, we investigate whether the terraces have a preferred occurrence on any given slope aspect, and, if so, whether this is geologically or climato-geomorphically controlled. Satellite images depicting the distribution of late-lying contemporary snow patches are used to assist in the discussion on topo-climatic linkages with scarp terrace orientation. Whilst it is not possible to accurately test the diverse and synergistic set of geomorphic processes that have operated since the origin of the Drakensberg basalt, an attempt is made to examine the variations in strength and relative surface age differences for basalt exposures along two slope profiles of varying altitudinal positions and aspect (north and west facing), so as to differentiate and debate relative rates of recent scarp weathering and recession.
2. Geological setting of the Lesotho Highlands The Drakensberg volcanic group is the product of a phase of discrete volcanic centers, which ejected material, followed by a phase of basaltic lava extrusions associated with dyke intrusions (Dingle et al., 1983). This later stage of mobile lava flows is estimated to have occurred about 182F2 Ma ago (Hooper et al., 1993). The Kraai River Formation, which typically displays columnar jointing and pillow lavas, has only been recognized in the eastern Cape regions (Dingle et al., 1983), while pillow lavas have been reported in the Indedema Valley in the Cathedral Peak region (McCarthy, 1970). The 800-m section of exposed basalt stratigraphy of the Sani Pass area has been divided into a lower 175-m section with four compositionally distinct stratigraphic units, and an overlying 625-m-thick succession of homogeneous Lesotho basalt (Mitchell et al., 1996). Each distinctive flow unit maintains its thickness for great distances (King, 1963) and averages 6 m (OSC, 1986), although the median value is 3.2 m (Galliers et al., 1991). Given that the terraces are found within the Lesotho basalt type, it may be assumed that the geochemistry for each terrace is similar (cf. Van Rooy and Nixon, 1990). The similarity in bulk chemical composition thus reduces the likelihood that chemical composition is an important determining factor in terrace development. Rather, the degree of alteration of primary minerals to secondary smectite minerals
S. Grab et al. / Geomorphology 67 (2005) 473–485
for each scarp should be used. The alteration of the basalt could play an important role in terrace development by affecting the depth of weathering (Van Rooy and Van Schalkwyk, 1993). An important observation that should be considered in the formation of basalt terraces is the exposure of an amygdaloidal zone at the base of each scarp, and in some cases near the top as well. A distinct zonation can be identified for each lava flow: a very thin aphanitic basalt at the base, followed by pipe amygdales or infilled vesicles, and then a massive central zone of weakly vesicular basalt and an upper layer of strongly vesicular basalt (cf. King, 1963; Nixon, 1973; Van Rooy and Van Schalkwyk, 1993) (Fig. 1). This zonation was observed in almost all scarp faces investigated, suggesting that each scarp face represents a single lava flow.
3. Methods Two terraced slope transects (at bearings 2808 and 3508) were investigated near Sani Pass, eastern Lesotho (Fig. 2). At this locality, terraces do not occur between bearings 508 and 2708. Slopes extend from 2960 to 3060 m asl, over which six distinct scarp outcrop (labeled as A–F) and associated terrace
475
features are identified (Fig. 3). The height of scarp outcrops varies between 3 and 8 m, whilst terrace tread diameters are between 29 and 126 m. A classification for assessing the strength of intact bedrock was first devised by Selby (1980) and later refined by Moon (1984). Typically, seven strength parameters are used to test rock mass strength (RMS) on particular slope gradients for the determination of strength equilibrium on slopes (Selby, 1982a; Moon and Selby, 1983). The RMS technique has been employed on a variety of rock lithologies including basalt, dolerite, quartzite, conglomerates, and shale (cf. Selby, 1980; Moon, 1991). For parallel scarp retreat to occur, a uniform RMS would be expected throughout the profile (Selby, 1982a,b). Near-vertical scarp faces of the Drakensberg basalt are apparently in equilibrium with contemporary processes, thus suggesting that current slope retreat is not dependent on rapid or catastrophic mass movement events (Moon and Selby, 1983). Perhaps, the most important parameter of the RMS classification is that of intact strength (hardness), which is measured using a Schmidt hammer (Day, 1980; McCarroll, 1991). The Schmidt hammer provides a rebound reading (R value ranging from 0 to 100), which is proportional to the hardness of the rock. The Schmidt hammer has been used for a number of rock weathering studies, including the role of late-lying
Fig. 1. Typical basalt structural zonation along a scarp face. Camera lens cap diameter=6 cm.
476
S. Grab et al. / Geomorphology 67 (2005) 473–485
Fig. 2. The site locality of the north-facing and west-facing terraced slopes in the Sani Pass area in eastern Lesotho. The bold line indicates terraces that parallel the course of the Pitsaneng River to the south of the map.
snow patches on bedrock weathering (Hall, 1993), its application to relative age dating (e.g., McCarroll, 1989; Nicholas and Butler, 1996; Sumner et al., 2002), and its effect on tafoni weathering (Matsukura and Tanaka, 2000). For the purpose of our study, an N-type
Schmidt hammer was used to assess the compressive strength characteristics (R values) per scarp face, so as to help establish relative rock surface age differences between the various scarps. Impact sites were cleaned using a carborundum and mean values were calculated
S. Grab et al. / Geomorphology 67 (2005) 473–485
477
Fig. 3. Sketch profiles indicating the position of scarp faces on the north-facing and west-facing slopes.
from 100 readings per site. In addition, the overall RMS rating for individual scarp outcrops was determined according to the various lithological parameters specified by Selby (1980). Rock surface roughness was measured to ascertain the relative difference in extent of surface weathering between individual scarps and to identify possible relations with R values. The roughness profile was recorded by using a profile gauge and calculated as a roughness index according to methods used by Nesje et al. (1994) and McCarroll and Nesje (1996). The use of lichenometry is well established as a relative age-dating technique (cf. Innes, 1985a,b; Nicholas and Butler, 1996), despite considerable geoecological complicating factors (Midgley, 1992). Measuring percentage lichen cover on a scarp face, as a relative age indicator for the exposed rock surface, is more suitable than measuring size–age or shape–age relationships (Innes, 1988). The underlying assumption is that the older, longer-exposed scarp face will have a higher percentage of lichen cover than a surface that has either been subjected to more rapid weathering and erosion, or only recently been exposed. We thus test this hypothesis against the other variables measured for each scarp outcrop. Tracing the lichen on the rock surface and transferring these onto graph paper obtained percentage lichen cover. Lichen were measured along 15-m belt transects per scarp face. Rock properties need to be considered in developing an understanding of cliff behavior (Douglas et al., 1991). Rock samples obtained from the field were cut into a standard size of 10–15 mm thickness, with a
volume/surface area ratio of between 4 and 5. An angle grinder fitted with a cement-cutting disk was used for the ethylene glycol test. Such tests have been used to determine the durability of Drakensberg basalts in geotechnical studies for the Lesotho Highlands Water Project (OSC, 1986; Van Rooy and Nixon, 1990; Van Rooy and Van Schalkwyk, 1993). The basalt specimens are immersed in undiluted ethylene glycol, which is an organic polar liquid. The alteration products of the basalts include montmorillonite, nontronite, and sapronite, with intergrowths of these clay minerals occurring and collectively accounting for ca. 30–40% of the total rock mineralogy (Van Rooy and Nixon, 1990). An important characteristic of these smectite group minerals is their capacity to absorb water molecules between the sheets, thus producing considerable expansion of the structure (Klein and Hurlbut, 1999). It is commonly found that the expansion of these clays upon wetting causes tensional breakup of the rock (Brekke, 1965; Douglas et al., 1991). This effect can be simulated by means of wetting and drying cycles (McGreevy, 1982; Hall and Hall, 1996), although it requires a large number of cycles before reliable results are obtained. The ethylene glycol test offers a simpler and quicker method, with results obtained within 30 days (cf. Van Rooy and Nixon, 1990) owing to the ethylene glycol substituting water molecules in the clay lattice to cause swelling (Gillespie et al., 1994). Such tests have been cautioned, as they are severe and do not necessarily equate to an accurate simulation of natural rock behaviour (Davies, 1989). However, for the purpose of this study, the ethylene glycol test results
478
S. Grab et al. / Geomorphology 67 (2005) 473–485
Fig. 4. Sketches indicating various components investigated in the rock samples following the ethylene glycol test.
classification of a composite of TM bands 3–5 was performed.
are used simply as a source of comparison between various scarp outcrops and considered together with the results from other tests performed. To quantify the relative changes, the samples were observed under 2 magnification and sketches drawn before immersion, after 15 days and again after 30 days of immersion (Fig. 4). Any changes in fractures and swelling of clays are then easily identified, and a qualitative method was employed for the characterization of the detected changes (Table 1). The terrace orientation frequency distribution was determined for an area extending 3 km west of the Great Escarpment and between 29820VS and 29830VS. Scarp tread orientation was measured as the outwardfacing normal-to-scarp lineation, as suggested by Nelson (1998). The same region was assessed for late-lying snow patch orientation frequency distribution using a Landsat 5 satellite image dated August 3, 1990. The previous cold front had passed through 5 days earlier, and so the image depicts snow patches that had survived at least 5 days. A fine unsupervised
4. Cycles of erosion and implications for terrace formation Most hypotheses for slope terrace formation have failed to consider the initial valley-forming processes. Many of the high-lying bevels in Lesotho are classified according to the first cycle of erosion (the Gondwana planation cycle) after the cessation of the lava outpourings, which was followed by the rifting of the Gondwanaland assemblage during the late Jurassic. This initiated the post-Gondwana cycle of erosion, which lies ca. 330 m below the Gondwana summit level (King, 1963). According to Partridge and Maud (1987), the Great Escarpment had receded by ca. 100 km by the mid-Cretaceous, and a further 20 km by the end of the Cretaceous. The Post-African I and II cycles had limited erosive impact on the Lesotho Highlands,
Table 1 Characterization and differentiation of rock alterations after 30 days of immersion in undiluted ethylene glycol (adopted from OSC, 1986) Swelling of clay spots Fracturing along existing joints Small fractures throughout Extensive fracturing Complete breakdown North-facing transect
=visible expansion of clay minerals on the surface =relative to unnaturally induced fractures =especially if radiating from amygdales and expanded clay spots =a high continuity of fractures =crumbling or flaking of rock fragments Scarp locality A
Swelling of clay spots Fracturing along existing joints Small fractures throughout Extensive fracturing Complete breakdown
x x x
West-facing transect
Scarp locality
B
C
D
E
F
A
B
C
D
E
F
x
x x x
x x
x x
x
x x x x
x
x x
x x
x x x
x x
x
x
x
S. Grab et al. / Geomorphology 67 (2005) 473–485
resulting in different surface expressions for the interior, compared to those for coastal regions (Nixon, 1973; Partridge and Maud, 1987). Continued continental margin uplift caused baselevel readjustments with enhanced river incision, supporting the suggestion that the Lesotho terraces are the consequence of fluvial action during erosional cycles (Van Rooy and Van Schalkwyk, 1993). An example of this is found along the Pitsaneng River where the terraces parallel the course of the thalweg (Fig. 2). It has also been suggested that the Lesotho slope hollows have been truncated by African cycle river incision (Dyer and Marker, 1979). However, it is evident from aerial photographic interpretation and ground truthing that the terraces continue unabated across these hollows. Therefore, the terraces cannot be attributed to river erosion or a period of erosion represented by the African cycle, which truncates all erosive cycles. This implies that the terraces do not owe their origin to an erosional cycle.
5. Geomorphological controls It has been suggested that the basalt terraces are the product of cryoplanation processes (Meiklejohn, 1992). Features generally known as cryoplanation terraces have also previously been referred to as dgoletzT terraces, altiplanation terraces, nivation terraces, and equiplanation terraces (see French and Harry, 1992, p. 146). Reger and Pe´ we´ (1976) define cryoplanation terraces as dminor step-like or table-like landforms consisting of nearly horizontal bedrock surfaces covered by a thin veneer of rock debris and bounded by ascending or descending bedrock scarps or bothT (p99). Cryoplanation terraces often occur din a tiered series culminating in a summit flat, but rarely encircle an entire mountain or hill continuouslyT (Nelson, 1989, p. 31). The Lesotho terraces perfectly fit the morphological description of cryoplanation terraces given here. In addition, certain rocks such as basalt are said to be particularly suitable for cryoplanation processes owing to suitable lithological attributes (Priesnitz, 1988). The central tenet to cryoplanation entails the occurrence of long-lasting snow banks against the scarp face, providing moisture for frost wedging within the exposed scarp face and subsequent solifluction (removal of weathered prod-
479
ucts) over the terrace tread (Demek, 1969; Reger and Pe´we´, 1976; Priesnitz, 1988; Czudek, 1995; Grosso and Corte, 1991; Thorn and Hall, 2002), all of which occur on the south-facing slope aspects of eastern Lesotho during contemporary winters (Fig. 5). However, the concept of dcryoplanation,T implying a particular dformT and dgenesisT (Thorn and Hall, 2002), remains subjected to critical debate. Thorn and Hall (2002) conclude that dto deny the fairly widespread existence of features commonly called cryoplanation benches, terraces, and pediments would be foolish: to claim that there is anything approaching an adequate explanation of their origin(s) would be even more foolishT (p. 548). There is also uncertainty regarding the genetic differences, if any, between cryoplanation benches, terraces, and pediments (Thorn and Hall, 2002). Whilst we refer to the Lesotho features as terraces, we do not imply any particular process origin. Periglacial research in the high Drakensberg has produced evidence for both contemporary frost wedging along scarp faces and solifluction over terrace treads (Grab, 1999, 2000). However, the role of freeze– thaw as a weathering process in landform development has raised more questions than answers (e.g., Thorn, 1979; Hall, 1991, 1995; Matsuoka, 1990, 2001). It is generally accepted that even if cryogenic processes are active, terrace formation frequently occurs along structural benches (French and Harry, 1992), or along zones of high joint density (Czudek, 1995). Work in Antarctica has shown that bench initiation may arise from a preferential suite of processes at sites of extensive dilatation jointing (Hall, 1997). The bench
Fig. 5. During colder months, long-lasting snow banks occur below south-facing scarp faces. Solifluction mantles are frequently visible on south-facing terrace treads.
480
S. Grab et al. / Geomorphology 67 (2005) 473–485
forms are extended and maintained through ongoing weathering, mass wasting, and fluvial action (Hall, 1997). The preferential occurrence of Lesotho terraces
Fig. 7. Wedges of ice protruding out of a highly jointed south-facing scarp face (top photograph). The lower photograph shows a typical terrace scarp and adjoining tread. Rockfall debris mantles sometimes occur below scarps where there is a high joint density and ground seepage.
Fig. 6. A comparison of basalt terrace frequency distribution with snow-covered area for various slope aspects (orientations). The test area extends 3 km west of the Great Escarpment, between latitudes 29820 VS and 29830 VS. Snow cover data were obtained from a satellite image dated August 3, 1990.
between 808 and 2008 coincides with about 89% of Lesotho basalt joints striking between 808 and 1408 (Nixon, 1973) (Fig. 6). Joints dipping out of terrace free faces are commonly associated with increased water seepage and sapping at the scarp base. The highly jointed and weathered free faces may eventually fail owing to high pore water pressure (summer) or pressures induced by wedges of ice (winter), which should not only be identified as an agent of weathering, but also as an erosion mechanism (Fig. 7). The highly angular rock fall debris is a product of jointing sets, which commonly have a 608 intersection angle. Data indicate that the highest snow covered area is between bearings 1408 and 2008, with a peak at 1808 (Fig. 6). The longevity of snow patches is limited on north-facing aspects due to the high angle of incidence of insolation. In contrast, the higher-frequency distribution of basalt terraces on south-facing slopes overlaps with a higher-frequency distribution of longer-
S. Grab et al. / Geomorphology 67 (2005) 473–485
Fig. 8. Rock mass strength results for the surveyed north-facing and west-facing scarps in the Sani Pass area. The strength equilibrium envelope is after Selby (1980).
481
slopes absent of scarp outcrops. It is therefore slope orientation, rather than the presence of scarp outcrops, that is the primary control determining the distribution of longer-lasting snow patches. The occurrence of wedges of ice within scarp joints, long-lasting snow banks at the scarp terrace knickpoint, and evidence of solifluction on the terraces indicates that geomorphic attributes and processes commonly associated with dcryoplanationT are present at sites in eastern Lesotho. The extent to which past and contemporary cryogenic processes are contributing to the extension and maintenance of the scarp terrace morphology remains unclear.
6. Results and discussion lasting snow patches (Fig. 6). Scarp outcrops promote shadow and consequential cooling effects at the knickpoint with terrace treads, thus assisting in snow patch preservation. However, the trend in snow cover distribution displayed in Fig. 6 is also characteristic for
Fig. 9. Distribution of rock hardness (Schmidt hammer rebound) values for the surveyed north-facing and west-facing scarps in the Sani Pass area (n=100 per scarp face).
Most scarp faces fall outside the strength equilibrium envelope, displaying high slope angles relative to their mass strength (Fig. 8). This may be attributed to the scarp outcrops being undercut at the point of contact with terrace treads. Such undercutting has been attributed to the headward extension of talus and regolith-covered slopes below the scarps (Moon and Selby, 1983), thus undermining the lowest joint blocks and leading to high-angle outcrops. RMS increases from the upper to the lower scarp outcrops (Fig. 8), mainly influenced by increasing joint spacing and higher rock hardness (rebound) values in the downslope direction (Fig. 9). Similarly, northfacing scarps have lower mass strengths than corresponding west-facing scarps. The west-facing scarps reflect somewhat higher rock hardness values than corresponding north-facing scarps (38 vs. 33; n=600 per transect) (Figs. 8 and 9). It could thus be argued that the extent of weathering or chemical alteration at the surface generally increases upslope and from west to north. The ethylene glycol test results support the downslope trend in R values (Table 1). Although there is little difference between the north-facing and westfacing rock surface results, there is a marked difference between the results for the highest (A) and lowest (F) scarp outcrops (Table 1). Samples from the uppermost scarp outcrop underwent extensive fracturing and breakdown, whilst those from the lowermost outcrop only show swelling of clay spots and fracturing along existing joints. Intermediate scarp
482
S. Grab et al. / Geomorphology 67 (2005) 473–485
outcrops (B–E) show variable alteration and R values. It is thus apparent that as the chemical alteration of the rock mass increases, more secondary minerals are produced and R values therefore decrease. These trends demonstrate that the rock surfaces near the summit scarp have been exposed for much longer than those near the base of the slope. The rock surface roughness profile results (Fig. 10) indicate a progressive decrease in roughness in the downslope direction, which confirms the trends in different surface weathering expressions discussed. Rock surface roughness may provide a relative indication of the length of surface exposure (McCarroll and Nesje, 1996) and thus confirms that the length of rock surface exposure generally increases from lower to upper scarp outcrops. The north-facing scarps have a higher mean roughness index than the corresponding west-facing scarps, which may be attributed to higher chemical alteration (weathering) on the north-facing slope (Fig. 10). Although no clear trend in lichen cover is noticeable on the west-facing slope, there is a general increase in coverage from the upper (17%)
Fig. 11. Percentage lichen cover on the surveyed north-facing and west-facing scarps in the Sani Pass area.
to the lower (37%) scarp outcrops on the northfacing slope (Fig. 11). Variations in microclimate, solar radiation, rock mineralogy, and weathering would complicate the trends displayed. It is nevertheless interesting to note that the general trend in lichen coverage on the north-facing slope closely resembles that for RMS and R values (Figs. 8 and 9). It appears that lichen coverage decreases where rocks are undergoing greater chemical alteration and more intense weathering, and therefore does not offer a suitable rock surface relative age indication in this region. Rock surfaces with enhanced lichen growth may sustain greater mass strength and granular stability, as lichen are known to insulate rock and intercept moisture that would otherwise contribute to weathering (Benedict, 1993). The relative contribution of biologically induced (e.g., lichen) rock weathering processes to basalt disintegration is still unknown to this region.
7. Conclusions
Fig. 10. Different scales of variance of rock surface roughness for the surveyed north-facing and west-facing scarps in the Sani Pass area (n=50 profiles per scarp face).
Although some of the Lesotho terraces may have been accentuated through phases of fluvial incision, they do not owe their origin to the so-called dcycles of erosion.T Geological properties of individual scarp outcrops forming distinct slope terraces in eastern Lesotho correspond with individual lava flows. Despite terraces occurring on all slope aspects, they are most common on southeast-facing positions, thus coinciding with the majority of joint strike directions. Lithological factors are therefore considered the primary control to such terrace development.
S. Grab et al. / Geomorphology 67 (2005) 473–485
Given that the dominant aspect position of late-lying snow packs and occurrence of ice in scarp joints is similar to the major joint strike directions, it is difficult to advocate topo-climatic geomorphological factors as the primary control for terrace development. Nevertheless, given that contemporary snow patches, ice, and frozen ground last much longer (several weeks) on south-facing slopes than on other aspects, it is likely that processes associated with cryoplanation have contributed to the modification and extension of south-facing basalt terraces in eastern Lesotho. Terraces located on north-facing aspects are less likely to be affected by cryogenic processes, but are subjected to higher temperatures and greater thermal and moisture fluxes than southfacing slopes (personal measurements). Scarp outcrops are likely to be impacted by a variety of mechanical, chemical, and biological processes, including ground seepage through fractures. Whilst it can be recognized that the suite of processes extending basalt terraces varies with aspect, it is still uncertain which set of contemporary processes operating on a particular slope aspect is most effective in extending the terraces. Results have shown a general increase in weathering extent and decrease in RMS and R values from west to north, and from lower to higher scarp outcrop localities. It would appear that the higher scarp outcrop localities, most particularly on northfacing aspects, display longer exposed rock surfaces than those at lower elevations along a given slope profile. The higher scarp outcrop surfaces demonstrate exfoliation and granular disintegration, whilst the smoother, less weathered surfaces of lower scarp outcrops may be associated with a more active slope retreat. Such lower scarp outcrops display considerable rockfall debris mantles at zones of ground seepage (Fig. 7). The upper north-facing scarp outcrops are less lichen-covered than on lower slope outcrops, which may be owing to increased granular disintegration and exfoliation, and greater exposure to insolation and desiccation as one advances altitudinally. It is most probable that a different set of geomorphological processes operates on various slopealtitudinal and slope-orientational positions at any given time. Small variations in lithology and geomorphological processes may account for differences
483
in weathering and mass movement surface expressions between the various scarp outcrops and slope terraces. The suite of contemporary geomorphological processes extending the dbasalt terracesT is both spatially and temporally complex, and requires further study.
Acknowledgements We appreciate the financial assistance provided by the Wits University Research Committee Grant and the IAG Grant to attend the IAG International Symposium in Addis Ababa, December 2002. Wendy Job is thanked for her assistance in producing the figures. Our sincere thanks to the two referees who provided valuable input and helped improve an earlier version of the paper. This paper has benefited from valuable discussions and correspondence with Cliff Ollier.
References Benedict, J.B., 1993. Influence of snow upon rates of granodiorite weathering, Colorado Front Range, USA. Boreas 22, 87 – 92. Bentivenga, M., Coltorti, M., Prosser, G., Tavarnelli, E., 2004. A new interpretation of terraces in the Taranto Gulf: the role of extensional faulting. Geomorphology 60, 383 – 402. Bielecki, A.E., Mueller, K.J., 2002. Origin of terraced hillslopes on active folds in the southern San Joaquin Valley, California. Geomorphology 42, 131 – 152. Boelhouwers, J.C., 1994. Periglacial landforms at Giant’s Castle, Natal Drakensberg, South Africa. Permafr. Periglac. Process. 5, 129 – 136. Brekke, T.L., 1965. On the measurement of the relative, potential swellability of hydrothermal montmorillonite clay from joints and faults in Pre-Cambrian and Paleozoic rocks in Norway. Int. J. Rock Mech. Min. Sci. 2, 155 – 165. Bqdel, J., 1970. Pedimente, Rumpffl7chen und RqcklandSteilh7nge; deren active und passive Ruckverlegung in verschiedenen Klimaten. Z. Geomorphol. 14, 1 – 57. Cinque, A., De Pippo, T., Romano, P., 1995. Coastal slope terracing and relative sea-level changes: deductions based on computer simulations. Earth Surf. Processes and Landf. 20, 87 – 103. Colman, S.M., 1976. Inherent factors in the flow of valley glaciers as a possible influence in the formation of stepped glacial valleys. Z. Geomorphol. 20, 297 – 307. Czudek, T., 1995. Cryoplanation terraces—a brief review and some remarks. Geogr. Ann. 77A, 95 – 105. Davies, G., 1989. Materials testing for Lesotho water scheme. S. Afr. J. Sci. 85, 483.
484
S. Grab et al. / Geomorphology 67 (2005) 473–485
Day, M.J., 1980. Rock hardness: field assessment of geomorphic importance. Prof. Geogr. 32, 72 – 81. Demek, J., 1969. Cryogene processes and the development of cryoplanation terraces. Biul. Peryglac. 18, 115 – 125. Dingle, R.V., Siesser, W.G., Newton, A.R., 1983. Mesozoic and Tertiary Geology of Southern Africa. Balkema, Rotterdam. Douglas, G.R., Whalley, W.B., McGreevy, J.P., 1991. Rock properties as controls on freeface debris fall activity. Permafr. Periglac. Process. 2, 311 – 319. Dyer, T.G.J., Marker, M.E., 1979. On some aspects of Lesotho hollows. Z. Geomorphol. 23, 256 – 270. French, H.M., Harry, D.G., 1992. Pediments and cold-climate conditions, Barn Mountains, unglaciated northern Yukon, Canada. Geogr. Ann. 74A, 145 – 157. Galliers, R.M., Brackley, I.J.A., Lephoma, M., Sympson, A.B., 1991. Expected engineering geology of the transfer tunnel. Proceedings of the 10th Regional Conference for Africa on Soil Mechanics and Foundational Engineering, Maseru, Lesotho, pp. 351 – 358. Gillespie, R.J., Eaton, D.R., Humphreys, D.A., Robinson, E.A., 1994. Atoms, Molecules and Reactions. Prentice Hall, New Jersey. Grab, S.W., 1999. Block and debris deposits in the high Drakensberg, Lesotho, southern Africa: implications for high altitude slope processes. Geogr. Ann. 81A, 1 – 16. Grab, S.W., 2000. Stone-banked lobes and environmental implications, high Drakensberg, southern Africa. Permafr. Periglac. Process. 11, 177 – 187. Grosso, S.A., Corte, A.E., 1991. Cryoplanation surfaces in the Central Andes at latitude 35 degrees S. Permafr. Periglac. Process. 2, 49 – 58. Hall, K., 1991. The allocation of the freeze–thaw weathering mechanism in geocryological studies: a critical comment. S. Afr. Geogr. J. 73, 10 – 13. Hall, K., 1993. Enhanced bedrock weathering in association with late-lying snowpatches: evidence from Livingston Island, Antarctica. Earth Surf. Processes and Landf. 18, 121 – 129. Hall, K., 1995. Freeze–thaw weathering: the cold region bPanaceaQ. Polar Geogr. Geol. 19, 79 – 87. Hall, K., 1997. Observations on bcryoplanationQ benches in Antarctica. Antarct. Sci. 9, 181 – 187. Hall, K., 1998. Nivation or cryoplanation: different terms, same features? Polar Geogr. 22, 1 – 16. Hall, K., Hall, A., 1996. Weathering by wetting and drying: some experimental results. Earth Surf. Processes and Landf. 21, 365 – 376. Harper, G., 1969. Periglacial evidence in southern Africa during the Pleistocene epoch. Palaeoecol. Afr. Surround. Isl. Antarct. 4, 71 – 101. Hooper, P.R., Duncan, R.A., Marsh, J.S., Duncan, A.R., 1993. The basalts of Lesotho, Karoo Province, southern Africa. Eos 74, 553. Innes, J.L., 1985a. Lichenometric dating of debris-flow deposits on alpine colluvial fans in southwest Norway. Earth Surf. Processes and Landf. 10, 519 – 528. Innes, J.L., 1985b. Lichenometry. Prog. Phys. Geogr. 9, 187 – 254. Innes, J.L., 1988. The use of lichens in dating. In: Galun, M. (Ed.), CRC Handbook of Lichenology. CRC Press, Boca Raton.
King, L.C., 1963. South African Scenery, 3rd ed. Oliver and Boyd, London. King, L.C., 1976. Planation remnants upon high lands. Z. Geomorphol. 20, 133 – 148. Klein, C., Hurlbut, C.S., 1999. Manual of Mineralogy. John Wiley & Sons, New York. Matsukura, Y., Tanaka, Y., 2000. Effect of rock hardness and moisture content on tafoni weathering in the granite of Mount Doeg-sung, Korea. Geogr. Ann. 82A, 59 – 67. Matsuoka, N., 1990. The rate of bedrock weathering by frost action: field measurements and a predictive model. Earth Surf. Processes and Landf. 15, 73 – 90. Matsuoka, N., 2001. Direct observation of frost wedging in alpine bedrock. Earth Surf. Processes and Landf. 26, 601 – 614. McCarroll, D., 1989. Potential and limitations of the Schmidt hammer for relative-age dating: field tests on Neoglacial moraines, Jotunheimen, southern Norway. Arct. Alp. Res. 21, 268 – 275. McCarroll, D., 1991. The Schmidt hammer, weathering and rock surface roughness. Earth Surf. Processes and Landf. 16, 477 – 480. McCarroll, D., Nesje, A., 1996. Rock surface roughness as an indicator of degree of rock surface weathering. Earth Surf. Processes and Landf. 21, 963 – 977. McCarthy, M.J., 1970. An occurrence of pillow lava in a basal flow of the Drakensberg volcanic stage, Ndedema Valley, Natal. Second Gondwana Symposium: Proceedings and Papers. Council for Scientific and Industrial Research, Pretoria, pp. 433 – 439. McGreevy, J.P., 1982. Hydrothermal alteration and earth surface rock weathering: a basalt example. Earth Surf. Processes and Landf. 7, 189 – 195. Meiklejohn, K.I., 1992. Some aspects of valley asymmetry in the high Drakensberg. S. Afr. Geogr. J. 74, 49 – 53. Midgley, R., 1992. The Visual Dictionary of Plants. Dorling Kindersley, London. Mitchell, A.A., Ramluckan, V.R., Dunlevey, J.N., Eglington, B.M., 1996. The basalt stratigraphy of the Sani Pass, KwaZulu/Natal Drakensberg. S. Afr. J. Geol. 99, 251 – 262. Moon, B.P., 1984. Refinement of a technique for determining rock mass strength for geomorphological purposes. Earth Surf. Processes and Landf. 9, 189 – 193. Moon, B.P., 1991. The significance of rectilinear bedrock slopes. S. Afr. J. Sci. 87, 208 – 210. Moon, B.P., Selby, M.J., 1983. Rock mass strength and scarp forms in southern Africa. Geogr. Ann. 65A, 135 – 145. Nelson, F.E., 1989. Cryoplanation terraces: periglacial cirque analogs. Geogr. Ann. 71A, 31 – 41. Nelson, F.E., 1998. Cryoplanation terrace orientation in Alaska. Geogr. Ann. 80A, 135 – 151. Nesje, A., McCarroll, D., Dahl, S.-O., 1994. Degree of rock surface weathering as an indicator of ice-sheet thickness along an east– west transect across southern Norway. J. Quat. Sci. 9, 337 – 348. Nicholas, J.W., Butler, D.R., 1996. Application of relative-age dating techniques on rock glaciers of the La Sal Mountains, Utah: an interpretation of Holocene palaeoclimates. Geogr. Ann. 78A, 1 – 18.
S. Grab et al. / Geomorphology 67 (2005) 473–485 Nixon, P.H., 1973. Lesotho Kimberlites. Lesotho National Development, Maseru. Olivier Shand and Lahmeyer MacDonald Consortiums (OSC), 1986. Lesotho Highlands Water Project, Feasibility Study, April 1986. Partridge, T.C., Maud, R.R., 1987. Geomorphic evolution of southern Africa since the Mesozoic. S. Afr. J. Geol. 90, 179 – 208. Peulvast, J.-P., Claudino Sales, V., 2004. Stepped surfaces and palaeolandforms in the northern Brazilian bNordesteQ: constraints on models of morphotectonic evolution. Geomorphology 62, 89 – 122. Plakht, J., Patyk-Kara, N., Gorelikova, N., 2000. Terrace pediments in Makhtesh Ramon, central Negev, Israel. Earth Surf. Processes and Landf. 25, 29 – 39. Priesnitz, K., 1988. Cryoplanation. In: Clark, M.J. (Ed.), Advances in Periglacial Geomorphology. Wiley, Chichester, pp. 49 – 67. Reger, R.D., Pe´we´, T.L., 1976. Cryoplanation terraces: indicators of a permafrost environment. Quat. Res. 6, 99 – 109. Selby, M.J., 1980. A rock mass strength classification for geomorphic purposes: with tests from Antarctica and New Zealand. Z. Geomorphol. 24, 31 – 51. Selby, M.J., 1982a. Controls on the stability and inclinations of hillslopes formed on hard rock. Earth Surf. Processes and Landf. 7, 449 – 467.
485
Selby, M.J., 1982b. Rock mass strength and the form of inselbergs in the central Namib Desert. Earth Surf. Processes and Landf. 7, 489 – 497. Sumner, P., Nel, W., Holness, S., Boelhouwers, J., 2002. Rock weathering characteristics as relative-age indicators for glacial and post-glacial landforms on Marion Island. S. Afr. Geogr. J. 84, 153 – 157. Thorn, C.E., 1979. Bedrock freeze thaw weathering regime in an alpine environment, Colorado Front Range. Earth Surf. Processes and Landf. 4, 211 – 228. Thorn, C.E., Hall, K., 2002. Nivation and cryoplanation: the case for scrutiny and integration. Prog. Phys. Geogr. 26, 533 – 550. Van Rooy, J.L., Nixon, N., 1990. Mineralogical alteration and durability of Drakensberg basalts. S. Afr. J. Geol. 93, 729 – 737. Van Rooy, J.L., Van Schalkwyk, A., 1993. The geology of the Lesotho Highlands Water Project with special reference to the durability of construction materials. J. Afr. Earth Sci. 16, 181 – 192. Veldkamp, A., 1992. A 3-D model of Quaternary terrace development, simulations of terrace stratigraphy and valley asymmetry: a case study for the Allier terraces (Limagne, France). Earth Surf. Processes and Landf. 17, 487 – 500. Vincent, P.J., Clarke, J.V., 1976. The terracette enigma—a review. Biul. Peryglac. 25, 65 – 77.