Mass Movements: Landforms Shaped Under the Force of Gravity

Abstract

“A heavyweight for Einstein – Probing gravity where no one has done it before”. This was the headline of a press release in April 2013 by the Max Planck Institute for Radio Astronomy (Bonn, Germany). John Antoniadis, a student of the institution and his colleagues put Einstein’s general theory of relativity to the test in a cosmic laboratory 7,000 light years from Earth, where two exotic stars – a pulsar or neutron star and a white-dwarf – are circling each other. Einstein’s theory states that objects with mass cause a curvature in space-time, which we perceive as gravity. And the scientists’ results matched Einstein’s prediction perfectly even in the most extreme conditions tested yet. It is almost impossible to imagine that the gravity on the pulsar’s surface is 300 billion times greater than the gravity on Earth. From all four universal forces of nature, gravity is the weakest one and equates 9.81 m/s² close to the surface. You can overcome it easily for example if you drop an iron nail on the floor and pick it up with a small bar magnet from a child’s toy. But gaze out of the window: Gravity is shaping the landscape everywhere around us when it exceeds the cohesive strength of slope materials and friction and as a consequence downslope movements of masses of soil, sediments or rocks occur. Collectively they are termed mass movements. The downslope movement can be very slow and almost imperceptible or very sudden and catastrophic when massive landslides leave a path of destruction in their wake. This chapter shows examples of different mass wasting processes and landforms.

Gravity acts everywhere in the universe and everywhere at the Earth’s surface, pulling everything towards the center of the Earth. But what really causes the force of gravity to take over and initiate slope materials to either fall, slide or flow, and thereby shaping the morphologies of mountain and valley systems, both on the continents and on the ocean floors? To answer that question we have to consider the three most important parameters that influence mass wasting by lowering the resistance to movement.

• 1. The steepness of the slope reflects its state of weathering and fragmentation of the bedrock as well as the geologic-tectonic preconditions. Slopes built up from hard rocks like massive granite or limestone that form sheer vertical cliffs. In contrast, slopes in softer rocks such as shales or volcanic ashes are more prone to weathering and are easily fragmented – they become unstable and slide downhill thereby assume to an equilibrium slope called the angle of repose. For unconsolidated materials, this angle of repose increases with grain size up to about 37°. At any slope that is steeper than the angle of repose and thus unstable, gravity will take over and lead to mass wasting processes that reinstall the stable angle of repose for the specific material.

• 2. Consolidated vs. unconsolidated slope material. Slope materials vary greatly depending on the geology of the local terrain. The stability of a slope depends largely on the rock type or constitution of the material that builds the local terrain. Slopes can be built up from unconsolidated materials (clay, silt, sand and gravel), which are loose and not cemented, or from consolidated materials that may be hard bedrock, or materials that are compacted or cemented by binding mineral cements.

• 3. The amount of water saturation in the slope materials. This parameter is defined by the porosity of the slope material and the amount of pore water, determined by precipitation intensities or other water sources the slope may be exposed to. Both natural and anthropogenic factors may play a role, and an example of the latter is represented by the loss of vegetation cover by deforestation. Water saturation of ground surfaces acts as a lubricant. As a result, the friction between particles or coarser fragments is lowered and movement is more readily initiated – the slope material may start to behave like a fluid and flows under the force of gravity downhill.

As previously mentioned, the precursors to all these mass movement events are weathering processes that weaken the massive bedrock structure along joints, faults, and predisposed bedding planes in sedimentary rocks. Even the pressure of freezing water and in particular repeated freezing and thawing cycles in higher mountain areas, a process called frost wedging, can trigger rockfalls.

Once a slope is unstable, a mass movement is highly likely and all it needs is a trigger to set things in motion. Often mass movements are initiated by ground vibrations such as those that occur during an earthquake or simply by heavy rainfalls during storms or monsoonal weather patterns. Ongoing weathering and erosion processes can gradually steepen slopes and initiate a sudden collapse of the slope flank as well.

Earth scientists classify different types of mass movements according to the nature of the material (coherent rock or unconsolidated material), the velocity of the movement (gradual with a few cm/year or sudden with up to several hundred km/h) and the physical nature of the movement – will the material fall, slide or flow? The latter implies that the moving material behaves as a fluid while sliding masses move more or less as one unit.

However, the initial mass movement type may change into another during the downslope movement and the characteristics of these different phases can vary significantly during one event (e.g., in terms of velocity, depending on the slope gradient and internal composition). Certain mass movements are very dynamic processes and tend to be highly complex, and changes in their physical behavior may occur within one single event.

1 Mass Movement of Hard Rock

Let’s turn to mass movements of hard rock and start to explore the most basic type: rockfalls (Figs. 8.1 and 8.2). During a rockfall, fragmented and detached materials from steep or vertical mountain slopes plummet to the ground and tumble over the slope in free fall motion with high velocity, usually over a rather short distance. Ceasing of the movement depends mostly on friction, i.e. the relation of the particle in motion and the roughness of the slope. Over longer time periods the material accumulates at the foot of the mountain in the form of a talus cone (Fig. 8.3a–c). Their profile often reveals a straight slope segment at less than the angle of repose and a basal concavity towards the valley floor. Talus cones formed by rockfalls exhibit a distinct sorting arrangement of their sediments over a longitudinal transect (Fig. 8.3b). Fine particles generally accumulate at the top, coarser particles at the bottom part. Talus cones are important components of alpine sediment cascades as they act as both sediment sinks for rockfall debris and sources for debris flows and, ultimately, the fluvial systems.

Fig. 8.1

(a) Rockfall of single large boulders along the “White Rim” of Canyonlands National Park in Utah, USA. Large parts of the softer underlying rock layers are covered by block- and boulder-sized rock debris. Their shape and size is predefined by the joint pattern in the overlying carbonate sandstone. Image shows a similar location as in Fig. 8.14b. (b) A larger section of the White Rim in Canyonlands National Park in Utah, USA. Numerous blocks below the White Rim escarpment demonstrate the importance of rockfalls for the back-wearing of the carbonate sandstone and, ultimately, the canyon. (c) Rockfall in Yosemite National Park, California, USA. The mass movement may also be termed as a rock avalanche, since rock volume is relatively high and the rock debris has reached the opposite valley slopes, pointing to a more complex process of downslope movement 38°05′20″N, 119°26′25″E. Traces of an older rockfall or avalanche can be seen in the right (northern) part of the image, where scarce vegetation marks the pathway of the rock debris (Image credit: ©Google earth 2012)

Fig. 8.2

(a, b) A rockfall at the western flank of Half Dome, a famous mountain in the Yosemite National Park. The sequence shows the Yosemite Valley before (a: January 2005) and after (b: July 2012) the event. Large scale weathering (exfoliation) of the granite basement favors the occurrence of rockfalls in the area (Image credit: ©Google earth 2012)

Fig. 8.3

Talus cones form below rockfall scars in the upper rock walls and accumulate against the valley bottom. Adjacent talus cones may coalesce laterally if enough rockfall material is delivered to the foot of the mountain slope. However, many talus cones consist of solid rock in their inner part, and only the surface layers are built by rock debris. In the initial stages of talus accumulation, the lower part of the slope/rock wall is protected from weathering and more importantly, from disintegration. In contrast, the unprotected upper slopes are subject to weathering and erosion for a much longer time. Over geologic time the talus neck grows vertically, thereby protecting higher slope areas from collapse. (a) Talus cones below almost vertical rock walls in the Val Mora (Graubünden, Switzerland, at 46°34′16.12″N, 10°19′36.11″E). When rockfall material in the upper cone parts becomes water-saturated, it is transported further downslope by debris flows, forming elongate scars and lobate deposits in the upper and lower parts of the cones, respectively. (b) Talus cones in the Southern Alps of New Zealand (upper Rakaia Valley, at 43°16′1.22″S, 170°58′41.61″E). The typical sorting pattern with larger blocks in the lower, distal part of the talus cone is visible. (c) Talus cones (right) and lateral moraines (left) south of Bow Lake, Banff National Park, Canadian Rocky Mountains (51°39′55.29″N, 116°27′29.96″W). In addition to the rockfall material, debris flows may develop from the talus cones due to water saturation of rock debris sections (Image credit: ©Google earth 2012)

During rockslides, the rock masses slide down a slope and, in some cases, move rather slow (Figs. 8.4, 8.5, 8.6, and 8.7). Large rock units may move as a single unit during rockslides. Internal structural weakness of the bedrock (e.g. along slope parallel bedding planes in sedimentary rocks) is the main reason for rockslides to occur in many cases.

Fig. 8.4

(a) Rockfalls and slides along the walls of a volcanic crater in the Galápagos. The material derived from the crater rim lies on top of relatively fresh and black lava. In contrast to several steep talus cones at the foot of the crater walls, the bright lobate deposit formed during one or several slides. The scar of the larger rockslide is clearly visible on the image. (b) Randa rockslide in Mattertal, Switzerland, with 30 million m³ of rock in 1991 (Image credit: ©Google earth 2012)

Fig. 8.5

(a) The dark vegetated area in the Ötztal valley (Austria) marks the Köfels rockslide, the largest slide in crystalline rocks in the European Alps. The deposit covers an area of at least 11.5 km². See text for more explanation. (b) Typical Toma landscape at Fern Pass, showing steep and conical Toma-hills, funnel- or basin-shaped depressions transformed into shallow lakes. The Toma landscape (lower image center) was formed after large rockslides ~4,100 years ago; the deep rockslide scar of the northern slide is visible in the upper left part of the image (Image credit: ©Google earth 2012)

Fig. 8.6

The mass movement events (rockslides, sometimes also referred to as rock avalanches) at Tschirgant Mountain occurred along the Inntal fault system, related to internal rock deformation and fracturing. The Toma landscape of the debris is clearly visible (image center). While new chronological information suggests that the events occurred during the 4th millennium before present, interaction of the rockslide with the late-glacial Ötztal glacier was also proposed (Image credit: ©Google earth 2012)

Fig. 8.7

(a, b) Sequence of images before and after the Maierato rockslide on February 15th, 2010, in Calabria, southern Italy. The slide was induced by heavy rainfall during the days before the slide, which affected the contact layer between the silt- to clay-rich and evaporitic late Tertiary sedimentary rocks (Image credit: ©Google earth 2012)

The two major slide types are rotational slides (also called slumps) and translational slides (these types also occur in slides of rock debris, see below). Both types may behave very differently, including long-term creep, catastrophic movement that is preceded by long-term creep and sudden catastrophic movement with no creep phase. Rotational slides or slumps are slides in which the downward and outward movement of a mass takes place on top of a concave upward failure surface; they usually move slower than translational slides. A translational slide is a mass that slides downward and outward on top of an inclined planar surface. In particular on steep slopes they can be very fast and destructive, and a major hazard for humans. Rockslides of higher velocity may disaggregate and transform into rock avalanches or debris flows during the process of movement.

Two well-known rockslide examples are the Tschirgant and Köfels events, located in the European Alps. The Köfels rockslide in the valley of the Ötztaler Ache River (Ötztal, Austria) is the largest slide in crystalline rocks in the European Alps (Fig. 8.5a). It is assumed that the friction during the slide generated high temperatures (~1,700 °C!), leading to petrographic changes of the underlying rocks. The slide deposit (dark vegetated area in the valley, Fig. 8.5a) covers an area of at least 11.5 km². It descended from the western slope of the valley (left) damming in transverse the Ötztal Ache River and collided with the opposite side of the valley. The rockslide deposit blocked the mouth of the tributary valley from the east (right), which found a new way of descending as a waterfall into the main valley. Due to the blockage, a 7 km long lake formed, which accumulated 92 m thick of sediments. Simultaneously the river has cut a deep gorge into the slide deposits.

The typical morphological pattern of large rockslide deposits is called Toma landscape (Figs. 8.5 and 8.6) and it is the result of an abrupt stop of the high velocity mass movement. It is characterized by steep and conical Toma-hills (named after a landscape in Switzerland) and is associated to funnel- or basin-shaped depressions, which may have transformed into shallow lake systems between these steep conical hills. A good example is found at the Fern Pass (Austria; Fig. 8.5b), where large rockslides ~4,100 years ago formed the typical Toma landscape.

The Tschirgant Mountain (2,370 m above sea level) is located in the Inn valley of Austria (Fig. 8.6). The massif is composed mostly of dolomites of the Wetterstein Formation and lies at the southern margin of the Northern Limestone Alps. It is separated from the metamorphic Ötztal basement complex by the northeast–southwest-striking Inntal fault system, which causes internal rock deformation and fracturing leading to the event. The Tschirgant rockslide detached from the mountain flank at a height more than 1,400 m (Fig. 8.6, the white rockslide scarp, called “Weißwand”, is still visible in the background). It had an estimated volume of ~230 million m³, spread over an area of 9 km². The maximum runout of the rockslide is 6.2 km into the mouth of the Ötztal Ache River, crossing the tectonic and petrographic boundary between the Northern Limestone Alps and the Crystalline Central Alps. The provenance of the rock fragments within the Toma-hills (i.e., the irregular and hilly landscape of the rockslide mass in the valley) gives testimony to this event as they are composed of limestone and dolomite as well as crystalline rocks. During the slides, the rivers Inn and Ötztal Ache were dammed for a considerable period of time as supported by lake sediments in the higher parts of the valleys. Radiocarbon ages of buried wood fragments at Tschirgant suggest two distinct mass wasting events in the 4th millennium before present, though interactions of the rockslide with the late-glacial Ötztal glacier (i.e., the rock mass accumulated on top of the glacier tongue and was (partly) transported down-valley with the glacier ice) was previously proposed.

In addition to large and rapid-moving rockslides with long runout, rock avalanches are among the most destructive mass movements of rocks (Figs. 8.8 and 8.9) – they can develop from the fall or slide of a large rock body and involve increasing fragmentation of the rock mass that generates rock fragments of different grain size, further decreasing its cohesion. Rock avalanches travel with extreme velocity (up to more than 150 km/h) over long distances down-valley and can transport large volumes of material. The long travel distances of large flow-like rock avalanches may be explained by the fluidization of the debris caused by the incorporation of air, the existence of trapped air below the moving rock mass, water saturation, and/or additional processes of rock fragment interactions such as dynamic rock fragmentation and granular agitation.

Fig. 8.8

Scar of the Val Pola rock avalanche (Valtellina, Italy), which occurred on July 28th, 1987. It destroyed the villages of S. Antonio, Morignone and Piazza. The heavy rainfall and related lateral erosion of debris deposits by the Val Pola River caused a massive slope failure and the motion of more than 35 million m³ rock debris material. Typically, the event underwent several stages: the initial debris avalanche was followed by debris slides, debris flows, and a mud flow, the latter caused by the outflow of a temporary lake, which had formed by damming of the river (Image credit: ©Google earth 2012)

Fig. 8.9

(a) Rock avalanche deposit (light brown, lower right of image) in Alaska (USA) accumulated on top of the Black Rapids glacier (63°27′35.33″N, 146°10′0.76″W). The material derived from the steep wall of the mountain located on the right of the image. The rock avalanche was caused by an earthquake on November 3rd, 2002, occurring along the Denali Fault in Alaska. (b) Another rock avalanche deposit on top of a glacier in the Denali National Park, Alaska (62°59′6.68″N, 150°30′2.34″W). The deposit seems to be built up by two separate rock avalanche debris lobes, most likely caused by two different events. The image shows blocks of more than 20 m in length (Image credit: ©Google earth 2012)

Mass movements such as rock avalanches are typically triggered by earthquakes, but particularly occurred during the transition from the last glacial period to the Holocene, when retreating valley glaciers and ice caps left behind deep valleys with steep slopes. The exposed slopes became unstable under the new ice-free conditions and tended to restore equilibrium conditions of the stable angle of repose. In addition, receding permafrost conditions in the Holocene further destabilized steeper slopes that were bound together by the subsurface ice. Some of the largest alpine rock mass movements have occurred in this period: the Flims rockslide (9–12 km³ rock volume) in the upper Rhine Valley in Switzerland was dated to ~9,500 years before present, and the largest rockslide in the crystalline Alps at Köfels (Ötztal Valley, Austria; Fig. 8.5a) with an exceptional rock volume of more than 2 km³ occurred during the early Holocene, based on radiocarbon dating of buried wood fragments and cosmogenic nuclide dating of surface boulders.

Decaying permafrost during the recent past with warmer than average temperatures is also thought to be a contributing factor to much more recent rockslides such as the Randa rockslides. These rockslides occurred only 3 weeks apart in the Mattertal of the Swiss Alps, involving > 30 million m³ of bedrock (Fig. 8.4b).

2 Mass Movement of Unconsolidated Materials

Mass movements of unconsolidated materials are slower than most rock mass movements, mainly due to the lower slope angles at which materials like sand, clay, silt or fragmented bedrock or various mixtures of these components become unstable. The moving masses may also include vegetation like trees or man-made materials from infrastructure such as parts of houses, vehicles or fences.

As is the case for hard rock movements, the velocity and the nature of the movement (falling, sliding or flowing) also determines the type of unconsolidated mass movements: The slowest type of unconsolidated mass movements is debris creep or soil creep which describes a generally gradual and slow downhill movement with a velocity of 1–20 cm per year depending on the slope angle, water content and vegetation density. Vertical tree growth is often offset by soil creep and is visible as typical “hook” growth features at the bottom of the trunk.

Particularly in colder climates on permafrost or in areas with strong winter frost, a certain type of soil creep called solifluction occurs, causing typical soil lobes and sheets on gentle slopes (for images see Chap. 13). The water in the top soil layer alternately freezes and thaws causing the soil to gradually creep downhill. In permafrost areas, these freeze-thaw cycles are restricted to the “active layer” above the permafrost ground and the movement is restricted to the summer.

Fluid mass movements like earth flows composed of relatively fine-grained material such as soil, weathered shales or clay stones, or coarser debris flows (Figs. 8.10 and 8.11) occur when high rainfall saturates permeable materials in surface layers that may overly less permeable rocks. If saturated by water, the friction between particles or coarser fragments is lowered, and sections of debris cones may become unstable and start to flow downhill. Where slopes are steep enough this results in a debris flow. Debris flows involve large volumes of sediment, but may flow very fluidly and reach high velocities. The typical central ravines with accompanying coarser ridges are formed due to internal differences in the flowing rock mass, the related pore volume and water contents, and, consequently, the friction within the rock debris, existing between the central and outer parts of the debris flow. While the debris material is still flowing in the center, the rock debris is stopped along the outer parts.

Fig. 8.10

(a, b) If saturated by water, the friction between particles or coarser fragments is lowered. Sections of debris cones or debris accumulations on the higher slopes may become unstable and start to flow downhill under the force of gravity, resulting in a debris flow. Though containing larger volumes of sediment, debris flows may flow very fluidly. The images show typical associated ravines with accompanying ridges on alluvial fans in northwest Argentina (23°43′6.40″S, 65°27′2.10″W) where debris flows are characteristic (Image credit: ©Google earth 2012)

Fig. 8.11

(a–f) Large alluvial fans in northern Chile, stretching westwards from the coastal cordillera to the coast. Large parts of the fans are built up by debris flow deposits deriving from the slopes of the Cordillera. Pathways of debris flows with typical ravines and accompanying ridges as well as terminal lobes are particularly visible in (b–f) (Image credit: ©Google earth 2012)

A flowing mass composed of predominantly very fine-grained material (finer than sand) and high water content is termed a mudflow. Mudflows and debris flows or avalanches are imminent natural hazards in volcanic regions where rainfall can soak layers of unconsolidated pyroclastic ashes or erupting lava melts huge amounts of snow and ice. In connection with volcanic debris and eruptions, these destructive mass wasting events are called lahars (Javanese for lava) (Fig. 8.12).

Fig. 8.12

La Casita landslide and lahar pathway in Nicaragua. On October 30th, 1998, hurricane Mitch resulted in intense rainfall in Central America, causing parts of the southern flank of the Casita volcano to slide. The original mass of the slide successively increased by incorporating volcanic material and water and turned into a lahar (from a hyper-concentrated flow and later a debris flow). The La Casita event exemplifies the dynamic complexity of numerous mass movement events, which are often composed of multiples stages with different characteristics of the process of movement. (see also URL: http://volcanoes.usgs.gov/hazards/lahar/casita.php) (Image credit: ©Google earth 2012)

Similar to rockslides, debris slides may travel downhill at lower speeds above one or more failure surfaces (Figs. 8.13 and 8.14). These planar detachments are often related to the existence of colluvial material or rock debris above more competent rocks.

Fig. 8.13

(a) Landslide scar on a steep slope on the Brazilian coast (23°06′45″S, 44°14′59″E). In January 2011, a period of heavy rainfall over several days caused the 2011 landslide disaster in the state of Rio de Janeiro, Brazil. The intense rainfall caused thousands of mass movements, in particular along the Serra do Mar between the states of Espirito Santo and Santa Catarina. The deep weathering of the crystalline basement favors the occurrence of debris slides, generating disintegrated rock debris and saprolite on the slopes directly above the crystalline basement. (b) The municipality of Nova Friburgo was heavily affected by the event. The image shows a rural area with numerous debris slides (22°14′0.33″S, 42°38′7.49″W)

Fig. 8.14

(a, b) The image sequence shows the same rural area in the municipality of Nova Friburgo (Brazil) before and after the January 2011 landslide event (22°15′03″S, 42°34′54″W) (Image credit: ©Google earth 2012)

Related to rock avalanches, debris avalanches are among the most dangerous and devastating mass movements. Here, in contrast to debris flows, the process of debris movement may pass into sliding or falling. They are common in volcanic and humid mountainous regions and can be characterized, similar to debris flows, by a fluidal behavior. Debris avalanches are very destructive as they can reach extreme velocities. Similar to rock avalanches, the process of movement may be explained by the influence of air, water (e.g., where glacier ice is incorporated into the flow) and rock fragment interactions. In 1962 and 1970, catastrophic ice-debris avalanches occurred on Mt. Huascarán in the Cordillera Blanca of the Andes. The 1970 event was induced by an earthquake and travelled down for 17 km at a velocity of 280 km/h burying the villages of Yungay and Ranrahirca together with more than 20,000 people in Peru.