Landslide susceptibility in cuesta scarps of SW-Germany (Swabian Alb) part of Vignettes:Vignette Collection
Keywords. GIS-based hazard modelling, slope stability, Hirschkopf landslide, landscape evolution, natural hazards. Damage to settlements and infrastructure as well as human casualties caused by landslides are increasing worldwide (Singhroy et al., 2004). In the low mountain areas of Central Europe the expansion of urban and industrial areas into landslide prone terrain increases the potential for slope instability (Damm and Terhorst, 2009; Terlien et al., 1995). In the context of climate change coupled with increasing precipitation rates regionally, investigations of natural hazards in the temperate zone are becoming more pronounced in scientific and public interest. To predict future developments, and to compile susceptibility maps, it is essential to know about past processes. Ancient landslides essentially contribute to landscape evolution and therefore, the retrospective analysis on the base of detailed field survey is a key to improve the quality of GIS studies (Brunsden and Lee, 2004). The study area is located at the steep cuesta scarp of the Jurassic escarpment of the Swabian Alb (SW-Germany, Fig. 1), which has been studied for many years (e.g. Bibus, 1986; Terhorst, 2001). During recent and historic times, large and small landslides have periodically occurred at the steep slopes of the cuesta scarp. The occurrence of the landslides is linked to a combination of causative factors, reflecting general natural settings in the study area. Primarily, the Swabian Jurassic escarpment, which rises 300400 m over its foreland, is characterised by very steep slopes, providing high potential energy for landslides. Furthermore, geological conditions, namely alternation between permeable and impermeable Jurassic rocks, are most important for slope instability. The plasticity of the Middle Jurassic clay (Callovian) and, in particular, the boundary between this clay and the overlying Oxford marls (ox1) is conducive to landslides (Fig. 2). The impermeable Callovian clays perch water and form the main spring horizon in the cuesta slopes. As a consequence, the Oxford marls moisten and the enhanced pore water pressure softens the marl layers. The main type of mass movement can be described as slump-earth flow (Fleming and Varnes, 1991), a combination of sliding and flowing movements (Fig. 2). Typical geomorphological phenomena of the slopes are rotational sliding blocks with an average length between 200300 m and a width between 2050 m. Field studies show that this type of slide mass forms large parts of the cuesta scarp. Inside the mass wasting deposits, small translational slides and flows are present (Terhorst, 2001). Stratigraphical investigations show that large rotational landslides occurred during the Pleistocene period, whereas during the Holocene period, only older mass wasting deposits were reworked and displaced by translational slides and flows (Terhorst, 2007). Pleistocene slides are combined mass movements, as slump-earth flows forming rotational blocks. The displaced material consists of Oxford limestones and Oxford marls. Extensive semi-circular main scarps on the upper slopes of the Jurassic scarp, marginal depressions on the block surfaces and the tongue-like foot of the landslides, are some of the typical features (Fig. 2). During the Holocene period, secondary landslides with high recurrence frequency and limited extent have reworked the palaeo-landslides. This happens at specific vulnerable spots within the Pleistocene landslide areas and is controlled by hydrology and therefore by precipitation rates (see Kraut, 1999). Due to these processes, the stability of the rotational block decreases successively. Even the slightest movement in the rotational blocks can reactivate the Pleistocene slide masses and lead to catastrophic results. The best-known example for a catastrophic event is the Hirschkopf Landslide in 1983 (Fig. 3), where a minor block displacement led to massive movement of detritus and bedrock on the upper slope and devastating flow processes and destruction on the lower slope (Bibus, 1986). Thus, the Holocene period must be regarded as the preparatory phase for the development of new and larger rotational movements. In general, broad areas of the Swabian Alb cuesta scarp must be classified as potential risk areas for mass movements (Neuhäuser and Terhorst, 2007). According to Kallinich (1999) up to 20 % of the slope areas of the W Swabian Alb cuesta scarp are moving actively. Of major importance is the fact that ca. 90% of the recent slope movements are coupled to Pleistocene slide masses (Bibus and Terhorst, 2001). A combination of field data, a comprehensive database on landslides (Kraut, 1999), and statistical analyses forms the base for the compilation of GIS-based landslide susceptibility maps using different methods. Thein (2000) created susceptibility maps on the base of logistic regression. Kreja and Terhorst (2005) generated a landslide susceptibility map for a residential area with house damages based on the Open Source GIS 'SINMAP' (Stability Index Mapping) published by Tarboton (1997) and Goodwin et al. (1999) (Fig. 4). This assessment approach requires detailed hydrological data and high-resolution digital terrain models (minimum resolution 10 m) and therefore, is applied to small areas. For example, the map shows that recently affected positions are directly linked to older slide masses and that the house damages belong to the stability index of "lower threshold" (Fig. 4). Neuhäuser and Terhorst (2007) applied the 'weights-of-evidence-method' to produce a susceptibility map for a large area of the cuesta scarp (Fig. 5). The method was originally developed for mineral exploration (Bonham-Carter, 2002). The main objective was to quantify the susceptibility to landslides, to identify and rank the preparatory (causative) sliding factors. During the last decades, damages due to mass movements have been consistently occurring in the study area. Although large mass displacements like in the Moessingen landslide of 1983 are rare, catastrophic events are in the realm of possibility. In general, the occurrence of recent landslides is closely linked to palaeo-landslides and this fact form an important basis for the development of risk models and the compilation of planning maps for the cuesta scarp. Through an integrative approach the maximum data available for the study area was compiled in order to create susceptibility maps for landslide hazard. Several modelling approaches showed that the evalutation of maps and archives only has not proved satisfactory. Implication of geomorphological field analyses and landscape evolution lead to more differentiated results in hazard modelling. Download SINMAP here: http://hydrology.neng.usu.edu/sinmap/getinfo.html
Slope formation in the Flysch zone of the Vienna Forest (Austria) part of Vignettes:Vignette Collection
Keywords: Landslide, Flysch, landscape evolution, Quaternary sediments Human activities take place in a space, which is influenced mainly by Quaternary sediments and forms. They affect present day geomorphological processes and hazards (e.g. Arnaud-Fassetta et al., 2005). To predict future environmental developments, also concerning landuse and planning management, it is essential to analyse and understand the past processes that led to recent conditions (Damm, 2006). Furthermore, the reconstruction of former landscapes and ecosystems is a prerequisite to evaluate and to measure human modifications in space and time. It is evident that the spatio-temporal susceptibility of morphodynamic processes depends on a long-term development (Neuhäuser and Terhorst, 2007). Thus, in studying natural hazards and/or geo-ecosystems as a core field of applied geomorphology, the chronological aspect plays an essential role for the assessment of landscapes (Birkeland et al., 2003; Trimble, 1975). The Flysch zone of the Vienna forest is an undulating landscape in the central European low mountains (Fig. 1). Flysch is composed of layers of (calcareous) sandstones, marly shales, calcareous marls and clay shists. The interbedded sediments are covered by Pleistocene periglacial cover beds and loess. In general, the Flysch regions are considered to be susceptible to landslides. Both petrography of the bedrock and the soil mechanical properties of the Quaternary sediments control the slope dynamics in the study area. The structure of the Quaternary sediments and occurrence of the dense bedded basal cover bed are responsible for the development of landslides in steep slopes in the northern Vienna Forest. The discrepancy concerning permeability in loess influenced layers and the underlying basal cover bed, consisting mainly of marly and clayey material, is one of the essential controlling factors for the initiation of landslides. The presented work analyses cycles of slope formation with respect to long-term Quaternary processes and current process dynamics. It aims to distinguish steps of slope formation for the Pleistocene and Holocene periods, and to gather information on the degree of landscape changes. Indeed, datable material is rarely available, but chronostratigraphical classification can also be done by analysing the development of soils and sediments (see Birkeland et al., 2003). In this context, the occurrence of loess on the slopes of the Vienna Forest is of major importance. Palaeo-surfaces covered by loess are at least of Pleistocene origin. Furthermore, the distribution of three main Pleistocene periglacial cover beds allows distinguishing Pleistocene forms from Holocene ones (cf. Semmel, 1996; Terhorst, 2007). In general, the cover beds formed by solifluction on top of permafrost in the former active layer. Cover beds in European low mountain areas are subdivided in loess-bearing cover beds, namely the upper and middle periglacial cover bed, and the loess-free basal periglacial cover bed. Concerning slope stability it is important to note that the basal periglacial cover bed in the study area consists of marly and clayey components, mainly. Whereas the middle and basal periglacial cover bed are of variable Pleistocene age, the formation of the upper periglacial cover bed is dated on the base of volcanic tephra to the Younger Dryas (Semmel, 1996), a cold period between 12 700-11 560 cal y BP (Streif, 2007). On the basis of pedological and geomorphological investigations, it is possible to infer five morphodynamic phases of slope formation that partly reoccur in an alternating pattern in the study area. (1) The first initial phase of slope formation is constituted by a completely developed sediment and soil profile (Fig. 2). Periglacial cover beds and loess sediments have been deposited on top of bedrock. Figure 3/1 displays the distribution of the sediments and the geomorphological setting before the onset of the mass movements (see Damm et al., 2009). The Flysch sandstone is covered by the basal periglacial cover bed (Fig. 2, II and C3). In flat positions of the palaeo-surface, the loess has been deposited and was covered afterwards by the upper periglacial cover bed. In places, where the loess deposits are absent, the upper periglacial cover bed is directly situated on top of the basal one. (2) During the second phase, landslides affect the surface. In the course of the process, the permeable loess deposits slide on the top of the basal periglacial cover bed, which mainly consist of impermeable clay and marls. This landslide led to a landscape change in terms of geomorphology, sediments and soils (Figure 3/2). In the middle and lower slope, most Pleistocene sediments were removed. Exclusively the basal periglacial cover bed is partly conserved. (3) Subsequent to the sliding of the loess layers, a transition to the third phase is initiated, which corresponds to erosion and weathering within the sequence. For the calcareous and marly sandstone, the tendency for rapid and profound decomposition in the course of decalcification is a key factor. In the case where the periglacial layers disappeared completely, bedrock is exposed at the surface and affected by decomposition under the influence of percolation water (Figure 4/3). Due to this process, more stable blocks of sandstone are separated inside a predominant sandy matrix. (4) The decomposition process prepares the onset of the fourth phase of slope evolution, with a second sliding phase (Fig. 4/4). Here, the deeply decomposed sandstone is affected by further slides in the lower slope. (5) Afterwards, the phases three and four may reoccur in alternating patterns and/or rockfall start in the decomposed sandstone (Fig. 5/5). At present, relevant slopes are susceptible to rock and debris fall as well as to translational slides. The study indicates that different stages of slope formation as described by the five phases are present in the study area. On the base of last Glacial topography and sediments (phase 1) mass movements, erosion, and disintegration of bedrock are active (phase 2 to 5) since the incision of the creek Hagenbach in the Late Glacial and Holocene. There are only minor areas that show undisturbed conditions, most of the studied slopes are affected by mass movements. In particular, exposed sandstone, due to intense decomposition, is predestined to translational slides as well as to rockfall processes. The specific stages of slope formation cause a distinct spatial distribution of Quaternary sediments and solid rock at the present land surface and affect the recent morphodynamics in the study area sustainably.