the contribution of landslides to landscape evolution in europe

a 9

~-.

bstract

After a review of some analytical and conceptual models of slope evolution with the intervention of mass movements, ti proposal is made of a series of indicators for the quantitative description of landscape evolution in relation to mass movements. The ratio between landy!ide mobilisation rate and downwearing rate is proposed as a quantitative measure of the significance of mass movements in landscape evolution.

A subdivision of European regicns with respect to their susceptibility to landsliding ds a process of landscape evolution is proposed. The role of landsliding for the evolution of landscape in some regions is described. The data presented show t.lat landsliding may in some cases be the main con:ribuaor to landscape change.

I. Introduction

The main geomorphologicai exyression of landscape evolution is the transfer of mass from denudation to accumulation areas; that is, from mountains and highlands in general to fluvi,al and coastal lowlands.

Denudation takes place as a result of a series of processes in which wind, water or ice are active erosive agents, and also as a consequence of slides and other mass movements. These latter phenomena are able to transfer considerable volumes of slope materials downslope under the action of grav- ity, playing an important role in the shaping of landscape features, especially in certain regions, where they can be the main morphogenetic pro-

* Corresponding, author. Triggering factors which determi;z.e the temporal

Elsevier Science B.V. SSDI 0169-555X(95)00070-4

cesses. They include a wide variety of mass movements, suzil as faiis, topples, slida, spreads and _flor~s (Cruden and Varnes, 1993), which may in- vohu volumes of rock, debris or earth ranging from a tew m3 up to several ~zJ~, on slopes with a gradient between 90 and less than one degree.

Cmditionirtg and triggering factors of mass movements are well known. The main condirionirtg .factors include geology and structure (rock and regolith stltingth; regolith tiiickness; number, type, size, spacing, relative position and continuity of rock partings; position of partings with respect to topographic slope) afid topography (relief, slope gradient an3 height). Long-term processes which can be considered as conditioning factors for slope instability, include weathering, tectonic activity and regional uplift, climate and sea-level changes. Human infiu- ence (land-use change, devegetation, interference with the water cycle) coin also affect slope stability.

occurrence of landslides include earthquakes to which permanent and transitional ground deformation, frac- turing, oriented acceleration, pore pressure increase and liquefaction can be related (Radbruch-Hall and Varrres, 1976; Solonenko, 1977), volcanic eruptions, meteorological events (such as heavy rainfall events, inducing water infiltration and pore pressure increase and air temperature warming, inducing glacier or ground ice melting), and human activity (excavation, overloading, concentrated water infiltration). Basal erosion, also a triggering factor, can be produced by channel deepening, lateral displacement of channels and by sea-cliff erosion; all of them, in turn, depending on geology, relief and climate.

Natural or non-human triggering factors should provide an explanation for the temporal occurrence of landslides during the early Holocene; roughly until the Atlantic “climatic optimum”. Human influence or artificial triggering factors must have played an increasingly important role in many parts of Europe during the second half of the Holocene.

As indicated above, the main expression of landscape evolution is the transfer of mass by denudation of upland areas and sedimentation in valley floors and sedimentation basins. Therefore a possible procedure to characterise the role of landslides in landscape evolution is to identify a series of quantitative descriptors or indicators which can express that kind of transfer, directly or indirectly and in relation to other landscape processes.

Deep-seated gravitational slope dcformatians (Zischinsky , 1969; Nemcok, ! 972; Ter-S tepanian, 1977; Sorriso-Valvo, 1984; Pasuto and Soldati, 1990; Dramis and Sorriso-Valve, 1994) are large-scale phenomena ( sackung, lateral spreading) which affect high relief slopes, generating peculiar landforms such as large depressions, scarplets and trenches. Under particular conditions, these phenomena can evolve into huge landslides of different types (Dramis and Sorriso-Valvo, 1994; Dramis et al., 1995).

A variety of models, both conceptual and mathematical, have been developed to describe slope evolution, with or without the intervention of landslides; however, there is a lack of quantitative data on absolute chronology and rates of landsliding, so that few attempts have been made at determining the contribution of landslides to landscape change in particular regions.

The consideration of the role of landslides in lan,iscape evolution can be approached in two different ways: establishment of the mechanisms and processes operating at the level of individual landscape features (usually a slope profile); determination of the quantitative contribution of mass movements to landscape evolution at the scale of wide area (a valley, a region). In the latter case suitable quantitative indicators must be defined to allow comparison with other processes of landscape change.

2. Modells of slope ev

The usual procedure when analysing slope types and slope behaviour is to use the “profile approach”, describing slopes by means of two-dimensional models, even though it is generally recognised that variations along the contour line can be important (Carson and Kirkby, 1972; Carson, i 9%; Selby, 1982; Fran- cis, 1987; Freeze, 1987; Kirkby, 1987; Richards and Lorriman, 1987).

Two broad types of slopes are generally recognised (Selby, 1982): (a) weathering-controlled slopes, in which the rate of weathering is less than the potential rate of regolith removal. These slopes are usually bare an.1 ?:ey reflect the type and structure of the outcropping bedrock; (b) transport-limited slopes, in which transport processes cannot keep up with the rate of regolith production; regolith accumu- lates on the slope and slope form is determined by the properties of the regolith which, in turn, also determine the kind and intensity of processes acting on the slope. Transportational slopes are an interme- diate case and they correspond to those where an equilibrium between rates of weathering and removal occurs, thus producing profiles without gain nor loss of material. Obviously, it is in the second type that landslides play a most significant role for landscape evolution even though rock falls and deep-seated gravitational slope deformations can also develop in the first type.

Models of slope evolution in the case of weathering-limited removal (Fisher. 1866; Bakker and Le Heux, 1946, Bakker and Le Heux, 1952; Kirkby, 1987) consider the retreat of cliffs and the conse- quent accumulation of talus deposits.

When slope evolution is transport-limited, it is

te and the actual se

e role of basal erosion in the

the slopes with active basal erosion will depend on the interplay between weathering, slope movements and basal undercutting. In the case of slow basal erosion, as in many river valleys, it has been ob- served that local asymmetry of valleys occurs when one slope is undercut by fluvial action and the other is protected by a floodplain or colluvial footslope. Strahler ( 1950) noted a correspondence between basal erosion and slope angle, and Carson ( 197 1) proposed that this reflected the difference in friction angle between densely packed and loose states, corresponding to the relatively unweathered regolith on basally eroded slopes and weathered material on protected ones.

The relationship between rate of channel downcutting and slope development in fluvial valleys has been discussed by Arnett (197 l), Richards (1977) and Richards and Anderson (1978). These authors showed that in river basins mean and maximum slope angles increase downstream to a peak value and then decrease, and that the frequency of landslides matches this pattern. According to the former authors this indicates that threshold slopes are limited to those sectors of the basin where depth and rate of valley deepening have reached a certain value but floodplain development has not yet protected slopes from basal erosion.

Rapid river erosion into glacial sediments has been shown to produce planar or rotational deep- seated failures associated with steep bluffs (Statham, 1975). The collapsed till looses shear resistance, experiences an increase in pore-water pressure and is then subject to a series of translational slides which reduce slope angle. Brunsden (1974) has described a

A know! dge of geo-

slope development. Stability models can be applied to slopes for

which good quality data on slope geometry, bedruck, hydrological conditions are available.

on different parameters can be assumed and the corresponding variations in the factors of safety calculated for a series of slices. The possibl tion of the slope through landsliding can predicted. It appears from the above that the role of landsliding and other mass movements for slope evolution will depend heavily on weathering, which introduces important changes in both soil and rock physical properties and in the position of the soil/rock interface, thus affecting instability (Carson and Betley, 1970; Kirkby, 1973; Statham, 1974; Francis, 1984).

The variations in chemical composition and physical properties of soils and their influence on the value of maximum stable slope angles, as well as the role of hydrological regime on the onset of slope instability have been discussed by Francis (1987). That author states that short-term, site-specific studies indicate that in some cases the form of steep slopes is controlled by mass movement, but that evolutionary interpretations using this approach become more speculative for longer time scales (and, indeed, larger spatial scales).

Models of slope development taking into account climatic, hydrologic, hydrogeologic and slope stability factors have been proposed by Freeze (1987). An idealised quantitative model of slope instability is presented by that author, which predicts slope angle for different geologic and climatic environments.

me different mechanisms which control dow material transport in transport-limited slopes and therefore, their evolution, are: slow mass movements, rapid mass movements, erosion by surface runoff, piping and subsurface runoff, solution. The model discussed is restricted to the analysis of rapid mass movements, and it considers four main factors: climate, hydrology, hydrogeology and slope instability.

Using different combinations of values for parameters (such as average annual rainfall intensity, number of storm events per year, average duration of each event, saturated hydraulic conductivity, soil porosity, width of hillslope, height of hillslope, cohe- sion, angle of internal friction, specific weight of soil), solutions can be obtained for a great variety of climatic, hydrogeological, topographic and soil situa- tions. The possible occurrence of conditions leading to slope instability can thus be determined and the role of mass movements in slope evolution can be assessed. The model makes it possible to consider the climati&onditions of the past, so that some consequences witn respect to the influence of climate changes on landslides at different periods can be drawn.

Although the model has a clear potential for understanding the range of slope angles in a region, it can by no mean:) explain the variability of slope gradients to be found in any one area (Rouse, 1975; Rouse and Farhan, 1976; Freeze, 1980; Moser and Wohensinn, 1983; Rulon and Freeze, 1985). Parame- ters such as satutated hydraulic conductivity, cohe- sion or angle of internal friction present considerable differences over short distances both along and across slope profiles.

One particularly interesting consequence of this model is that the mechanism of slope evolution depends on the relative values of rainfall intensity and saturated hydraulic conductivity of the regolith. If the former is higher, overland flow will develop and surface erosion will be the main process. If the latter is higher, subsurface pore pressures will be the main agent, leading to slope instability. Climk!te changes can, of course, have produced in the past or produce in the future variations in the relative value of these parameters. Slope angle may thus be used as an hdiCatiOn of past climate in regions with low sensitivity of soil types with respect to instability.

A departure from the usual profile approach for the explanation of slope evolution has been proposed by Dalrymple et al. (1968) and Conacher and Dal- rymple (197%. Recognising that changes along contour lines can be just as important as those along a profile, these authors developed a land surface model based on the identification and mapping of a series of units which can be found in any slope, from channel to divide. The evolution of the slope can thus be described in terms of the processes dominant in each unit. It is also possible to develop mathematical models of slope evolution for each unit, although the linkage between the expressions for the different slope segments and the modeN+; of the evolution of the whole slope is not clear. Ot ler models, such as the one by Ahnert (1976) can incude landsliding and flow processes.

One of the qualitative consequences which can be obtained from these types of models is that slopes evolving through slides tend to show parallel retreat and basal debris accumulation. If the rate of channel incision is greater than the rate of transport along the slope, downcutting of the latter takes place and then it collapses by landslides.

From the above it can be seen that mathematical models have a limited value for explaining the long- term evolution of landscape in a region and the role that different processes - including landslides - play in it. Mathematical models are extremely useful in the sense that they provide the means to isolate the effects of different factors, to make predictions and to test them. However, they do not yet provide a precise representation of the complexities and het- erogeneities of actual slopes, except at the profile level and in relatively simple slopes. Moreover, for the application of models it is necessary to have good data on the different parameters controlling the evolution of a slope. Thus, the sound and accurate application of a model depends not only on the correct representation of the processes by the corresponding equations, but also on the availability of detailed field measurements over a sufficiently long time. This is difficult enough for specific slope sites, and nearly impossible for a region, even of a limited extent.

Because of these limitations, and as stated by Selby (1982) due to the fact that we cannot define an initial landform to which we could apply the model

r role for landscape evolut

a given region at suitable conditions,

changes which can describe the evolution of a

identification of the cause of failure of landslides, their relative chronology and their potential instability within the general context of landscape evolution. The model for landslide systems proposed by those authors has 24 dependent and 7 independent variables. If the temporal and spatial scale of the analysis are sufficiently large, some independent variables controlling landslide occurrence can be disregarded. Temperature, precipitation and evapotranspiration are part of climate. River discharge and groundwater inflow depend upon climate and tectonic changes. Rock properties can be ignored assuming that materials prone to failure do exist in the region. Thus, the three fundamental independent variables controlling the occurrence of landslides in a region, over the time span (Jf millennia or tens of millennia, are: climate, earthquakes and tectonism.

In a region, slopes with gradients near the threshold value may fail as a consequence of a triggering event, such as an earthquake, an intense rainfall event or the oversteepening of the slope through channel incision or artificial excavation (Schumm, :l973; Carson, 1976; Coates, 1977).

The operation of the incision landslide model is represented in Fig. 1. In narrow valleys landslide deposits produced as a result of slope oversteepening are continuously removed and they become relict and inetastable upon isolation from basal stream erosion. In wide valleys, deposits become relict when they xe protected from further incision by terraces, but they can be destabilised if lateral channel migra- tion 3rodes those terraces.

The climatic landslide model is shown in Fig. 2. Climate changes towards greater rainfall or lower

Fig. 1. Phases in the incision model for landslide development. F: failure plane, S: scar, D: landslide debris, Tp: paired terrace, Tn: non-paired terrace (after Palmquist and Bible, 1980).

temperature favour increased infiltration, wi increasing pore pressures and decreasing shear strength. Moreover, greater available moisture can cause an increase in stream discharge and channel incision, which in turn increases gradients and also favours instability. As climate becomes warmer or drier, the reduction in pore pressure increases shear strength and stability, and landslides become relict and stable.

Earthquake shaking can increase normal pore pressures, with a corresponding decrease in shear strength; the slope may thus fail. This model is particularly important for areas of high relief and

ACTIVE

\ ,’

RELICT .

Fig. 2. Phases in the climatic model for landslide development, with stable baselevel and with climate-induced incision. W: water table, F: failure plane, S: scar, D: landslide deposit, Tp: paired terrace, Tn: non-paired terrace (after Palmquist and Bible, 1980).

weak materials. The temporal distribution of landslides in this model shows peaks which coincide with periods of maximum seismic activity, followed by a gradual decrease in landslide frequency. The spatial distribution is that of an ellipse with its long axis roughly following the fault which generated the earthquake (Wilson and Keefer, 1985).

The identification of the landslide model operating in a region is based on the chronology of landslides and their position on the slopes. Climatically triggered movements appear on hillsides, Iglder and younger terraces, valley floors and older landslide deposits. Incision landslides never rest on paired terraces nor hillslopes; they appear at the foot of ‘the slope, on valley floors or at the reactivated toe of older slides. In the case of seismically induced landslides the distribution is similar to those produced by climatic causes.

Of course, the different models can operate con- currently. In this case the temporal and spatial distribution of landslides will be more complicated and their role in landscape evolution will be qualitatively and quantitatively greater.

According to Palmquist and Bible ( 1980), the recognition of the type of landslide model operating in a region, or the cause of individual landslides, not only provides the means to establish how landsliding has taken place within the overall evolution of landscape, but it also provides clues concerning their present stability. Climatically induced older landslides should be more stable than seismically induced, as the former took place under climatic conditions for which threshold slope gradients were lower than present ones.

3. Towards a characterisation of the role of mass movements in landscape change

3. I. Indicators

The work summarised above focuses on the analysis of the evolution of slope types, normally using models which make a series of simplifying assumptions, difficult to apply when it comes to describe the evolution of large areas (river or stream basins, morphodynamic regions, etc.). In this case it is nor-

ma1 to find a considerable heterogeneity of materials, landforms and relief and a variety of processes are usually operating in any individual slope, both in space, along and across the profile, and in time, as a consequence of changes in climate and/or other triggering factors. For this kind of assessment an empirical approach, based on the determination of the spatial distribution. magnitude, frequency and rates of m;lterial removal by mass movements seems to be better. Unfortunately, analyses of this kind which require the determination of absolute ages for movements covering a fairly long period (at least thousands of years) in a given area are not frequent.

‘4 possible approach to estimate the contribution of landslides and other mass movements to general landscape evolution is to consider a series of areas characterised by different combinations of conditioning factors, in different morphodynamic environments (climate, tectonic activity, human influence), trying to determine how landslide activity has evolved in them during a certain period, say, the Quatemary or the Holocene.

Landslide activity directly creates a series of landforms and deposits and has an indirect influence on the development of other landscape elements, such as fluvial-lacustrine deposits produced by the damming-up of rivers, soils, vegetation or even the distribution of human settlements. Accordingly, other indicators which could be useful are those related to the persistence in the landscape of landslide- originated elements, to the time needed for certain elements to develop or to “heal” after a landslide event, or the degree of correlation between landslide deposits and forms and some human features.

Dossible indicators which could be used are: 1.’ 2.

3.

4.

5.

6.

number of identifiable movements; area affected by landslides (absolute values, per- centage) total or for a given period of time; volume or mass mobilised (total or for a given period); mobilisation rates (average or for specific periods); geomorphic work (total, rate), for the whole area ~.i- per unit surface; magnitude of individual movements (maximum and average size, in area and/or volume), comparison with other landforms and deposits in the area;

ratio

x-ate;

A. Ccwdrero. F.

isatio

Im

s soil or vegetation, to form again

temporal (occurrence of move relationship with other geom cesses in landscape evolution.

If the main quantitative expression of landscape evolution is the transfer of mass from denudation to sedimentation zones, a

I on transpo~-limited slo s, lower than 1 on weath- s and about 1 on tr~sportatiomal

! I

I - i ‘1 yp. J

0 I (

-___ -_- ___ _- __ ________ ________ ___.___ __P -A-

Fig. 3. Distribution map of landsliding prone terrain units in Eurqe: (1Aa) high relief Alpine chains, <lAba) moderate relief Alpine chains modeiied in hard rock, ( 1 Abb) moderate relief Alpine chains modeiied in soft and moderately hard rock, (1 B) Hercynian high reliefs, ( 1 C) Caiedonian high reliefs, (2A) pm-Alpine erogenic mid-relief areas, (2B) sedimentary structures not (or slightly) affected by orogenesis (basins and pi>tforms), (3A) pre-Alpine erogenic low-relief areas, (3B) alluvial and fluvio-glacial plains and late Quaternary marine deposits, (4) vwicanic reliefs.

198 A. Cendrero, F. Drumis / Gemorphology IS f 1996) 191-211

3.2. Subdivision of European regions with respect to their susceptibility to landsliding as a process of landscape evolution

In order to carry out a preliminary quali Lat:ve assessment of the relative importance of lan&litie~ for landscape evolution in Europe, the continent has been divided into a series of large units, defined on the basis of large-scale morphostructures, relief and lithology (Istituto Geografico De Agostini, 1965; Gaertner and Walther, 1971; Derry, 1980; Embleton, 1984) also considering the typology and frequency of landslide occurrence in different areas (Common, 1966; Zairuba and Mencl, 1969; Tufescu, 1970; Ka- menov et al., 1973; Balteanu, 1974; Panizza, 1973, Panizza, 1974; Gil and Kotarba, 1977; Mahr and Nemcok, 1977; Cotecchia, 1978, Cotecchia, 1986; Vanmaercke-Gottigny, 1978; Nemcok et al., 1982; Coltorti et al., 1984; Embleton, 1984; Dufaure, 1984; Loye-Pilot, 1984; Carton et al., 1987; Canuti et al., 1988, Ayala and Ferrer, 1989; Fialho Rodrigues and Gomes Coelho, 1989; Flageollet, 1989; Kotarba, 1989; Schack Pedersen et al., 1989; Ancochea et al., 1990, Ancochea et al., 1994; Haeberli et al., 1990; Maquaire, 1990; Corominas and Baeza, 199 1; Hutchinson, 199 1; Koukis and Ziourkas, 199 1; Ffister et al., 1993; Casale et al., 1994; Dramis et al., 1995).

These units are represented in Fig. 3 and Table 1 and arc briefly described below.

3.3. Description of terrain units

I. Highlands

1 A Alpine chains - These morphostructures were formed during the Late Tertiary and are still affected by active tectonic deformation and uplift.

1Aa High relief Alpine chains - These units underwent strong uplift and deep valley incision, and were also intensely affected by Pleistocene and present glaciation, which oversteepened valley-slopes. Permafrost is present at the highest elevations.

The landslide susceptibility of these units is high to very high: all the types of rapid landslides (e.g. debris flow, debris avalanche, rock fall, rock avalanche, rock slide), especially in deeply weatb- ered hard rocks or in glacial and periglacial debris deposits, are common; deep-seated gravitational slope deformations (Sackung, laterul spread), sometimes evolving into huge landslides, are also frequent, especially in highly seismic areas and in connection with tector\ic slopes (e.g. active fault scarps, thrust fronts); shallow slides and flows (soil slips) frequently occur in thick soil covers as a consequence of heavy rainfall.

The main slope instability factors are: high relief due to uplift and river incision and glacial valley deepening, oversteepening of slopes due to glacial erosion, rock weathering and jointing, present and residual tectonic stress, mountain permafrost degra-

Table I Landsliding prone rerrain units in Europe

1. Highlands I A. Alpine Chains 1 Aa. High Relief Alpine Chains (Area: 4.4%) 1 Ab. Moderate Relief Alpine Chains I Aba. Modelled in Hard Rock (Area: 11.7%) 1 Abb. Modelled in Soft and Moderately Hard Rock (Area: 7.6%)

IB. Hercynian High Reliefs (Area: 10.9%) IC. Caledonian High Reliefs (Area: 8.1%)

2. Willy Areas

3. Lowlands

2A. Pre-Alpine Orogenic Mid-Relief Areas (Baltic Shield, Caledonian andHercynian Reliefs), (Area: 13.9%) 2B. Sedimentary Structures not (or slightly) affected by Orogenesis (Basins and Platforms) (Area: I I .I %) 3A. Pm-Alpine Orogenic Low-Relief Areas (Baltic Shield, Caledonian and Hercynian Reliefs) (Area: 10.2%, 3B. Alluvial and Fiuvio-Glacial Plains; Late Quatemary Marine Deposits (Area: 16.8%)

4. Volcanic Reliefs (Area: 3.3%)

-lasting rainfall (triggeri reaching slides).

1Ab were le affected by glaciation; Pleistocene periglacial processes (frost shattering, slope wash, solifluction) have played an important morphogenetic role covering the slopes with thick overburden nants of planation surfaces are ofte of the reliefs.

lAba Developed in hard rock (igneous, metamorphic, limestone, sandstone etc.) - Ltxldslide susceptibility is moderate: deep-seated gravitational slope deformations, sometimes evolving into huge landslides, are fairly common specially in highly seismic areas and in connection with tectonic slopes; rapid mass-movements (such as debris flows, debris avalanches, rock falls and rock sliti’es) also occur on steep slopes.

The main slope instability factors are: relief due to uplift and river incision, present and residual tectonic stress, high seismic@ (especially in Mediterranean areas), heavy rainfall (both intense and long lasting events).

1Abb Developed in soft rock (mostly flysch formations) - Landslide suscr?ytibility is high to very high. The most frequent gravitational phenomena are landslides with recurrent activity such as translational slides (in structurally complex, stratified formations); rotational slides (in clayey-shaly bedrock formations, in soil overburden, in thick periglacial solifluction deposits filling small valleys and slope depressions; earthflows and mudflows (ranging in length from some tens or hundreds of meters up to some kilometres), favoured by intense rainfall; lateral spreadings, block slides and huge earthjlows triggered by earthquakes or extreme meteorological events.

The main s!ope instability factors are: ‘ ‘ unstable’ ’ bedrock lithology; past periglacial morphogenesis (especially in northern areas); river valley deepening and lateral erosion: present and residual tectonic stress; heavy rainfall: seismic@ (especially in

Landslide suscept ity is moderate to hi& (in ation to bedrock I logy and valley dissection) e most common vitational phenomena are:

deep-seated gravitational slope deformations sometimes evolving into huge landslides; rapid mass- movements (such as debris flows, debris avalanches, rock falls and rock slides) in deeply weathered rocks or glacial and periglacial deposits; rotational and translational slides (which often evolve into mudflows) affecting marly bedrock materials or soil covers as a consequence of heavy rain.falI.

The main slope instability factors are: refiefi “unstable” bedrock lithology; deep weathering; heavy rainfall.

1C Caledonian high reliefs - These units were built up by Early Palaeozoic orogenesis and underwent a long erosional evolution resulting in widespread surface planation. They were also intensely affected by Pleistocene glaciation which covered the reliefs with a thick mass of ice and underwent strong isostatic uplift after the deglaciation. The present reliefs are generally poorly dissected, showing plateau-like surfaces at their top on which active glaciers are still present at the highest altitudes and latitudes. Past periglacial solifluction deposits are fairly widespread on the slopes; active periglacial processes aft&t the area at the highest latitudes. Coastal morphology is characterised by the presence of high cliffs and fjords.

Landslide susceptibility is Iow to moderate (moderate to high along the coast). Most common gravitational phenomena are: rock slides (sometimes of huge dimensions) on high steep slopes, oversW?p- ened by glacial erosion; debri: /Law and earth$ow movements in loose glacial and neriglacial debris materials.

The main slope instability factors are: high relief due to post-glacial isostatic uplift; glacial oversteepening of slopes; past periglacial morphogenesis; coastal erosion; heavy rainfall.

2. Hi@ areas

2A Pre-Alpine erogenic mid-relief areas (Baltic Shield, palaeozoic erogenic belts)

Landslide susceptibility is 10~9 to moderate (mostly depending on bedrock lithology). Most common gravitational phenomena are: rotational slides and earthflows (generally of small dimensions) affecting marly or clayey bedrock or past periglacial deposits; large-scale rotational-translational slides along coastal cliffs.

Slope instability factors ;ue: “unstable” bedrock lithology; past periglacial morphogenesis; coastal erosion: anthropic activity (e.g. agriculture works, excavations); coastal erosion: heavy rainfall.

2B Sedimentary structures not (or slightly) affected by orogenesis ( pre-Alpine basins and platforms, post-Alpine Plio-P!eistocen&lts)

Landslide susceptibility is moderate to high (in coastal areas and., politic tenains). Most common gravitational phenomena are: rotational slides, trarzs- lational slides and earthjlows in clayey bedrock; rotational slides, eartllflows and mudflows (generally of small dimensions) affecting past periglacia! deposits; large-scale rotational-translational slides along coastal cliffs.

Slope instability factors are: “unstable” bedrock lithology; past p&glacial morphogenesis (in north- em countries), coastal erosion; river erosion; anthropic activity (e.g. farming, aqueduct and waste water infiltration into the ground, excavations), heavy rainfall; seismic activity in Mediterranean areas (post-erogenic Plio-Pleistocene belts).

3. Lowlands

3A Pre-Alpine erogenic low relief and gently rolling areas (Baltic Shield, Palaeozoic erogenic belts)

Landslide susceptibility is very low to 1~ (moderate along entrenched rivers and cliff coasts). Most common gravitational phenomena are: rock falls af-

fecting hard bedrock on tional slides, earthflows bedrock and overburden periglacia! deposits.

river escarpments; rota- and mudflows in clayey materials, including past

Slope instability factors are: deep river erosion; coastal erosion: anthropogenic activity (e.g. agricul- tural works, excavations); heavy rainfall.

3B Alluvial and fluvio-gtacia! plains: late Quater- nary upraised marine deposits

Landslide susceptibility is low (moderate to high along entrenched rivers and cliff coasts). Most common gravitational phenomena: topples affecting scqs in alluvial materials; rotational slides, earthflows and mudflows sometimes reactivating ancient large-scale landslides, triggered by deepening of river erosion during the Pleniglacial regression; rotational slides affecting upraised marine deposits; mudflows affecting glacial and periglacial clay deposits, specially in relation to snow melting.

Slope instability factors are: deep river erosion; coastal erosion: past glacial and periglacial morphogenesis; heavy rainfall: snow melting.

4. Volcanic reliefs These units include active and inactive volcanoes

(volcanic reliefs s.s.) and volcanic reliefs s.1. modelled on bedrock of volcanic origin (lava, tuff, vol- canoclastic materials).

Lundslide susceptibility is moderate to very high in volcanic reliefs S.S. (volcanoes), where lahar phenomena (during eruptions) and huge grauita- tional movements (triggered by eruptiyns or earthquakes shocks) may be common. Landslide susceptibility is normally low to moderate in ?~oLc~~nir reliefs s.l., where debris slides and debris flows may affect pyroclastic materials or topples and falls may occur on lava and pyroclastic deposits, swcially when they overlay a clayey sedimentary bedrock.

Slope instability factors are: vclcanic activity; earthquakes: “unstable” slope materials; heavy rainfull; coastal erosion (in volcanic islands); waste-water infiltration from human settlements.

As already considered in the description of the different terrain units, seismicity and climatic conditions play an important role in determining both the

Fig. 4. Schematic maps of climatic and seismic zones in Europe. (A) Clitnurk ;orws: I. limit of climatic zones (1 = oceanic climate, II = Mediterranean climate, III = continental climate), 2. limit of the last glaciation, 3. limit of older glaciations. 4. mountain permafrost. (B) Seismic hones: I. very high seismicity, 2. high seismicity, 3. moderate seismicity, low seismicity.

As a first approximation, the classification of local conditions with respect to landslide susceptibility and, therefore, their potential role for landscape evolution, can be attempted using an urweightcd aggregation system. Each one of the units identified in the three maps presented can be ranked xcording

Table 2 Area1 percentages of subunits deriving from the combination of terrain units of Table I. climate and seismicity in Europl:; k very hi@ seismicity, II: high seismicity, HI: moderate seismicity, IV: low seismicity _I. . .

nect both rapid and reaching mass movements

terised by frequent rainfall, seem more prone to the triggering of shallow landslides (mostly affecting scil covers) and slow landslides characterised by frequent reactivations.

The superposition of the above described terrain its to large regions which can also be identified in

urope with respect to climate and seismicity (Istituto Geografico De Agostini, 1965; Embleton, 1984; Blundell et al., 1992) (Fig. 4) gives a total of 120 potential subunits deriving from the combinations of relief-bedrock, climate and seismicity. The percent- age corresponding to each one of these subr-lnits is shown in Table 2.

Climatic zones

Seismic zones

IAa 4.4 I Aba I I.7 I Abb 7.6 IB 10.9 IC 8.1 2A 13.9 2B II.1 3A 10.2 3B 18.8 4 3.3

Oceanic

I II

0.1 ,

0.2

0.7 0.8

III

1.4 0.7 !.I 0.3 0.8 0.2

IV

0.1 0.3

1.7 2.3 2.4 1.4 2.2 2.7 0.2

Mediterranean

I 11

0.5 0.5 3.2 0.9 1.2 I.1 0.2

I.1 0.5

0.6 0.1

III

0.5 0.3

0.9 0.5

0.5

IV

I.5 I.1 3.7

0.1 4.2

0.5 0.4

Continental

I II

0.5 1.2 1.3 1.2 0.5 I.1 0.4 0.3

0.8 0.8 0.2

0.1 0.2 1.8

a__- III IV

I.5 2.8 2.1 3.2

5.1 2.6 4. I 3.9 0.1 2.7 4.5 2.9 10.4

to landslide instability. This procedure can be carried scale landslides. Synthetic values of susceptibility to

out with regard to different mass movement typolo- landsliding of the Table 2 subunits can also be gies such as, for example, rapid landslides - RL calculated (Table 3). (including rock falls, rock avalaaaches, rock slides, According to the landslide typologies considered, debris avalanches, debris flows, twth flows), slow the synthetic value of any subunit thus characterised recurrent landslides - SL, and deep-seated gravita- would then vary between 3 and 12. Although this is tional slope deformations DGSD and connected large a very crude “index” it can provide a rough idea of

Table 3 Rank of units according to their susceptibility to mass movements of different typ, s (see Figs. 4 and 5) and synthetic landslide susceptibility unweighted values for the subunits of Table 2 (n”: values for terrain units; nc: values for climatic zones; ns - values for seismic zones).

Climatic zones Oceanic ( 1’) Mediterranean (3’) Continental (2’)

Seismic zones I (4s) II (3s) III (25) IV (IS) I (4s) II (3s) III (2s) IV (IS) 1 (4s) II (3s) 111 (2s) IV (IS)

A. Rapid lundslides (RL)

lAa(5”) 5”3s lC 1 Aba (3”) 1 Abb (4”) 4”3s lC IB (4”) IC (2”) 2A (2”) 2B (2”) 3A (I”) 3B (1”) lU3Slc 4 (2”)“’ 2”3’lc

B. SloW recurrent krndslides (SLJ 1 An (4”) 4”3s3c IAba (2”) IAbb (5”) 5”3s3c IB (3”) IC (2”) 2A (2”) 2B (4”) 3A (I”) 3B (1”) lU3s3c 4(1”) l”3s IC

4’2’1’ 2”2”lC l”2s2c l”2s2c l”2slc l”2slc

3”2S3C 2”2s3’ 2U2S3c 4”2s3c lU2s3c I” 2s;c

s”lslc 3” IS lC

4”lSlC 2Ulslc l”lslc 1” lS2c 1” lSIC l”lslc 2”lSlC

4” ls3c 2” lS3’

3” ls3c 2” lS3c 2” ls3c 4” lS3c I” ls3c 1” ls3c lUISlc

5” 2S3c 3”2’3’ 4” 2’3’

3” ls3c 4” lS3c 4” ls3c

5’4’2’ 5U3S2c 5’2’2’ 3’4’2’ 3U3S2c 3’2’2’ 4’4’2’ 4”3’2’ 4’2’2’ 4”45 2c’ 4’3’2’ 4’2’2’

lU4S2c

2’4’2’

lU3S2c

4”3’2‘

l”2s3c

l”2s2c

l”ls2c l”4’2’ 1” lS2c 1’4’2’

1” lS3c 2” lS3’

2” Is2c 5” lS2’ 3” ls2c

2” ls2c 2”4s 1’ 4” lS2c 4”4s lC

1” ls2c 1” ls3c

I”4s2c

4”4s lC 4”3s IC 4’2’1’ 2”4’1’ 2’3’1’ 2”2slc 5”LES 1 c S”3’IC 5”2’lc 3”4s IC 3”3S IC 3U2slc

l”4s lC

1’3’2’

2U3slc

l”3s lC lU3S2c

2UlS2c I” ls2c 2”1s2c 1” ls2c 1” ls2c

2” IS lC 2UlSlc 4Ulslc l”lslc l”lslc

C. Deep-seu fed gruvitutionul slope c1ejhw~ution.s und connected lurge scale lundslides (DGSD)

1Aa (5”) 5”3s lC 5”lSlC 5”4S3C 5”3S3C 5”2”3C 5U4S2c 5”3=2’ 5U2S2c 1 Aba (4”) 4”lSlC 4U4S3c 4U3S3c 4U2S3c 4” ls3c 4”4s 2’ 4U3S2c 4’2’2’ IAbb (3”) 3”3s lC 3”4s3c 3”3s3c 3” 2s3c 3” lS3c 3’4’2’ 3U3S2c 3’2’2’ IB (4”) 4”2slc 4”lSlC 4”4s3c 4” ls3c 4’4’2’ 4U3S2c 4’2’2’ IC (3”) 3’2’1’ 3”lSlC 3” ls2c 2.4 (I”) l”2slc l”lslc 1”4s3c lU2s3c 1” lS3’ 1’4’2’ 1’3’2’ l”2s2c 1” lS2c 2B 12”)(*’ 2”2slc 2”lSl 2”Y3’ 2”2s3c 2” lS3c 2U4S2c 2”2s2c 2UlS2c) 3A (I”) l”2slc l”lslc l”2s2c 1” ls2c 3B (I”) l”3slc l”2slc l”lslc l”4s3c lU2S3c 1” ls3c 1U4S2c l”3S2c 4(1”)‘“’

l”2s2c 1” lS2c 2’3’1’ 2” Is 1’) 2U4S3c) 2” l”3’ 2U3S2c --

t’)4” in the case of active volcanoes and high relief volcanic hills. (*‘4” in c@astal areas affected by active tectonics.

(3)5u in the case of active volckxs and high relief volcanic hills.

Landslide phenomena including ro falls and avalanches, debris an ea slides, are widespread in the Alps ( and Castiglioni, 1984; Canuti et al., 1988; Casale et al., 1994). These phenomena can reach large-scale dimensions such as the Becca Fomace Aosta (50 million m3 in calcschists rocks) which occurred in July 1564, the rock fall (60 million m3 in calcschists) which dammed the Varaita valley in 139 1, the enormous prehistoric slide (250 millions m3 in ophiolithic rocks) whose reactivaction destroyed the village of Villar (in the same valley) in 1655 (Miiller, 1944). Many of these landslides represent the last evolutionary stage of deep-seated gravitational slope deformations which frequently occur in high mountain reliefs (Mahr and Nemcok, 1977; Crescenti et al., 1994). Catastrophic debris flow and earthflow phenomena are often connected with ex treme rainfall events and torrential floods (Govi, 1983) and earthquakes (Govi, 1977). A noteworthy triggering factor for debris flows and other landslides is also the melting of mountain permafrost due to the recent climatic warming. In the Swiss Alps more than 60% of the debris flows activated during sum- mer 1987 originated from the periglacial belt, at altitudes ranging from 2300 and 2400 m a.s.1. and Haeberli et al. (1990) verified that for five of those mass movements the source area was characterised by permafrost conditions. Also in the western Alps large scale landslides and deep-seated slope deformations mostly originate between 2000 and 2500 m a.s.l. (Mortara and Sorzana, 1987), an altitude belt which corresponds to the lower limit of discontinu- ous perm:.?rost, particularly sensitive to climatic changes. An example of the high incidence of landslides in the Alps is given by the Province of Son- drio (Adda River Basin) where 1339 phenomena have been recognised over an area of 3212 km* (Pozzi and Sfondrini, 1972).

Rotational and translational slides and mudflows often occur in the alpine reliefs modelled in flysch, such as the Flvsch Carpathians (1 Abb) (Tufescu,

enomena are

to 2.55 cases/km*, in relation to bedrock lithology (Vallario and Coppola, 1973).

In Calabria (sou em Apennines; lAba) Ian& slides are also widespread, affecting about 25-35% of the whole territory (Sorriso-Valvo, 1985, Sorriso- Valvo, 1989; Casale logical evolution of has been determined by a rapid rate of uplift, up to 1 mm yr-‘, and by a humid climate with intense chemical weathering and hea rains (present rainfall ranges from 500 to 2,2 mm, with heavy downpours which have return periods of about 25 years). Seismicity also plays a role as a determining factor for landslides.

Rotational and translational slides account for 56.9% of movements affecting all types of rocks. They are more frequent, of course, in soft sediments, 7 i 96, arid less fieqtieiit iti Giittk ids, 48%. Nearly 90% of Holocene slides are reactivations of pre-Ho- locene movements. The largest one mobilised 50 X lo6 m of Miocene clays. Flows represent 33.7% of all movements (10% on rocky slopes and 47% on clays). The minimum threshold slope height for flows to occur appears to be 160 m. One example, the Constantino landslide-flow, with a volume of 16 X lo6 m3, decreased average slope angle by 1” and slope height 15 m. Falls, 9.4% of mass movements, are restricted to gneiss and sandstones, and never produce large movements.

A clear relationship between slope height, angle and lithology on the one hand, and type and size of landslides on other hand, has been found in Calabria (Sorriso-Valve, 1985). It appears that slopes are evolving through extreme meteorological (and seismic) events, which produce landslides but leave the geomorphological “ style’ ’ of slopes unchanged. Some processes, such as rockflow-rockslide-debris flow on metamorphic rocks operate through a ‘ ‘cyclic’ ’ model which determines slope evolution.

In areas subject to seismic influence, both the size of individual landslides and total area affected tend to be somewhat greater, but no significant differ-

ences exist in typology and morphometry between recently earthquake-triggered movements and older presumably climate-triggered ones. This suggests that seismicity does not change the style of landscape evolution, but it probably increases its rate. This is confirmed by the fact that only 5% of tie area affec;ed by iandslides after an earthquake represent first-motion movements. The rest are reactivations of older movements.

The size of individual phenomena in this region can be very considerable. The Fiumara-Buonamico landslide, which occurred in January 1973 as a result of an intense rainstorm, moved 16 X lo6 m3 of debris, releasing 2.16 X lc”3 Joule. The height of the slope deceased by 60 m and the slope by 2.5 de- grees. The adjacent r.ver bed was elevated 40 m at the site of the landslide and as much as 10 m two kilometres downstream. The largest landslide-debris flow in Calabria has a 1.9 km* scar and a 3 km long fan. It represented a total energy release of 4 X 10 I5 Joule. A great number of huge landslides were triggered by the catastrophic earthquake which took place in southern Calabria in 1783. It involved a 1700 m displacement of a slide 180 m thick, dammed-upstreams in the Gioia Tauro plain and created 200 new lakes (Cotecchia, 1978).

Deep-seated gravitational slope deformations and connected large scale landslides are fairly frequent in the Umbria-Marche Apennine (1 Aba). Systematic geomorphologic analysis of the area allowed to recognise about 500 phenomena, with a mean exten- sion of 2 km* (maximum about 15 km*) and a frequency of 0.06 cases per km* (Dramis et al., 1995).

A particularly unstable condition is that of coastal cliffs, specially if developed in marly, shaly or clayey bedrock such as in Normandy, southern Wales, Eng- land and Ireland (Hutchinson, 1973, 1991; Brunsden, 1974 Brunsden and Jones, 1976; Bromhead, 1979; McGreal, 1979; Williams and Davies, 1985; Flageol- let, 1989; Maquaire, 1990; Casale et al., 1994). A spectacular example is that in Calvados in Nor- mandy, where the coastal escarpment, made up of chalk resting on marls, is affected by mass movements of different types. Among these the rotational-translational slide of Le Bouffay (a piece 350 m long and 50 m wide which moved suddenly in August 1991) and the rotational slide of Cricque-

boeuf (east of Trouville) which destroyed several buildings in January 1982 (Maquaire, 1990). Large- scale coastal landslides are also frequent in Bulgaria (Kamenov et al., 1973) and in the periadriatic belt of Central Italy (Coltorti et al., 1984, Esu and Grisolia, 1991) where active tectonics is also a triggering factor (Dramis et al., 1995).

An interesting fact to be pointed out is the similarity in the value of RLR in the Canary Islands (4) and the Cantabrian Cordillera (1 Aba). Despite the great differences in the periods considered, relief, geology, present climate, type and magnitude of individual landslides, this ratio is in both areas approximately between 0.2 and 1. This could be simply a coinci- dence or the result of inaccurate estimates based on insufficient data. Nevertheless, the possibility of a relative constancy of this ratio, for a wide range of morphodynamic conditions, should be considered and investigated.

The contribution of mass movements to landscape evolution in the Cantabrian Cordillera has been discussed by Cendrero et al. (1994) and Gonzalez Diez et al. (1996). They show that the role of these processes varies significantly according to geological and geomorphological environment.

In the Magdalena-Pas valley, on Mesozoic clay- stones, siltstones and sandstones, nearly 50% of the study area has been affected by late Pleistocene to Holocene mass movements. The occurrence of these movements coincided with periods of increasing rainfall, as it would be expected. Mobilisation rates vary betwe..II 0.06 mm yr’ ’ for the Pleistocene to 0.52 mm yr”’ for the first half of the Holocene. In the last 500 years, as a result of human influence, the rate increased to 0.73 mm yr- i . The average rate for the last 100,000 years was about 0.13 mm yr- ’ .

Denudation rates determired in two small basins vary between 0.004 and 06 mm yr-’ for the Holocene. Average denudation rates for the whole valley during the Quatemary range between 0.05 and 0.2 mm yr- ‘. Stream incision rates have been estimated between 0.15 and c).5 mm yr- I. Landslide deposits and landfomrs are significant elements of the landscape, with a maximum size of 2.3 km* and 266 X lo6 m3 (the study area has 150 km*). As it would be expected, the “landscape life” of these features is related to their size. No deposits with an area less than 2000 m* persist for more than 3000

r than mobilisation rates.

Nalon, landslides represent an important control for human features. over half of the settlements in these very rural areas appear on large Pleistocene landslide deposits. These materials provide an adequate sub- stratum for the development of soils and relatively flat areas on the slopes, thus creating the appropriate conditions for forest clearance, agriculture and graz- ing, which were the basis of traditional settlements.

Landslides can represent in some cases the main factor of landscape evolution and relief degradation. Some of the Canary Islands are a good example of this type of situation. The main denudational features

is, about 4% of the total volume of the

to a “ mobi~isation rate’ ’ of 0.076 mm yr - ’ . At least two episodes of large landslides have

been identified in L ma (Fig. 6; Ancochea et al., 1994). The volumes ilised during the last episode can be estimated at 25 km”, corresponding to an average rate of 0.05 I mm yr- ’ for the last 700,rOOO

1 volume of the island is approxi- landslides representing a per-

centage similar to the one in Tenerife, about 5%. In Wierro (FGster et al., 1993) three large features

cover nearly fifty per cent of the island (Fig. 5). The landslide of Las Playas, to the SE, formed less than 0.29 Ma ago and the one of El Golfo, to the I’&

--

n / 7 cl /’ 2

d

D Fig. 5. Large Quatemary landslides in the Canary Islands: I. limit of landslide, 2. surface cracks after 1949 eruption (modified from

Ancochea et al., 1990, Ancochea et al., 1994 and Ftister et al , 1993).

206

between 80,000 and 6000 BP. The third one, in El J&in to the SW, has not been dated. The volumes affected by these movements were 3 km3, 15 km” and 6 km3 respectively. This represents 11% of the total volume of the island, about 225 km3. The mobilisation rate for the last 300,000 years was some 0.26 mm yr- ‘.

These events represented a huge release of energy. The landslide which formed Las Caiiadas re- leased about 5 X lOI* Joule; the smallest one described, Las Playas, represented an energy loss of 7 5 X lOI . Joule .

Events similar to the ones described for these islands took place in Gran Canaria, also in the Quatemary, and during the Miocene-Pliocene in the

islands of Fuerteventura and La Gomera (Cantagrel et al., 1984; Coello et al.. 1992), but they are not so we’ll documented.

Landslide processes in these islands seem to have an episodic character. Volcanic lavas and pyroclas- tics accumulate during periods of 10 to lo* ka, until the stress due to the accumulation of materials (with thicknesses which can be estimated between 500 and 1500 m) exceeds a threshold value and part of the volcanic edifice collapses. These landslides are es- sentially planar, with low angle slip surfaces following weak pyroclastic or sedimentary layers. Volcanic eruptions continue accumulating materials over the collapsed block or in other parts of the island and the process takes place again.

15Ma

07Ma

12Ma

-n---7 -e--b8 09 c&J 10 ?? 11 12

Fig. t:. Subaerial evolution of the island of La Palma. showing two large landslide episodes around 1.2 and 0.7 Ma, the latter followed by smakr movements: (l-6) volcanic formations, (6-7) landslide limit, (9- I I) eruptive centers (after Ancochea et al., 1994).

a. Landslide features re

erosion but also of accumulation of newer volcanic materials. This is the case of the depression of Las CaZadas, partly filled by the ptions of the Teide- Pica Viejo during the last ,000 years, but still clearly recognisable.

Finally, it must be mentioned that the potential destructiveness of this low-frequency events is important, because of the very large volumes mobilised. In fact, only major volcanic eruptions with large pyroclastic emissions would represent a greater hazard.

Destructive episodes similar to the ones presented here have been described in other regions; among others rjawaii, (Duffield et al., 1982) Reunion (Duf- field et al., 1982), Martinique (Vincent et al., 1989), French Central Massif (Cantagrel and Birot, 1990), Volcin Colima, Mexico (Luhr and Prestgaard, 1988), Ecuador (Nieto and Schuster, 1988; Ficcarelli et al., 1992). Sicbert (1984) discusses 98 such avalanches in volcanoes all over the world and Ui et al. (1986) find these kinds of deposits in 49% of Japanese volcanoes.

5. Conclusions

A variety of analytical models exist which explain the evolution of slopes in relation to different determining and conditioning factors. These models allow to make quantitative predictions of the response of slopes to changes in different parameters, but normally only at the scale of specific slopes and under simplifying assumptions which do not correspond to real field conditions. The models can also be applied

e also been proposed to spatial and temporal occurrence of slope

movements in a region, as a consequence of the models explain the mation about “how

much” landslides are contributing to lan

e characterisation of different morphodynamic environments on the basis of large scale features, as well as the series of indicators which have been proposed, repres s a possible approach for the quantification of role of these processes in land- sc:&pe evolution. The elative Landslide Rate is con-

icularly useful for “measuring” the rela- nce of landsliding in different areas.

The analysis of data available for some European re:.ions shows that mass movements are in some cases the main process of landscape evolution. The apparent similarity bf&ween the values of the RLR in very different environments suggests the possibility of a relative constancy of this index in a wide range of geomorphological conditions.

More data should be obtained from different regions, in order to gain a better understanding of the contribution of these processes to landscape evolution and to test the ideas suggested here.

This paper is part of the EC Climatology and Natural Hazards project EPOCH (CT 90.005 CERG) “Temporal occurrence and forecasting of landslides in the European Community”. Publication No. 16. Complementary funding was given by the CICYT, Spain.

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