Introduction

The geological environment in the Tibetan Plateau is one of the weakest environments in the world. This area has experienced complicated geological processes including uplift combined with fluvial or glacial downcutting, earthquakes, and precipitation, with the geomorphology characterized by steep hillslopes and deeply incised valleys (Deng et al. 2019). Large-scale landsliding is of great significance in shaping landscapes in the Tibetan Plateau (Dortch et al. 2009). Especially, massive landslide debris could rush into river valleys and form dammed lakes, causing catastrophic floods threatening to human life and property because of the breakage of the dams (Fan et al. 2014; Korup and Tweed 2007; Zhang et al. 2019). For example, the 2014 landslide disaster near Oso in Washington State, USA, caused more than 40 fatalities (Iverson et al. 2015), and the catastrophic Baige landslide occurring on October 10, 2018, in southwest China generated a dammed lake causing a direct economic loss of approximately 1.0 billion USD (Tian et al. 2020). Besides, records show that some of the largest floods on Earth were caused by the rapid release of water from landslide-dammed lakes (O’Connor and Costa 2004; Zhang et al. 2019; Fan et al. 2020).

To date, many large river-damming landslides have been mapped along the rivers in the Tibet Plateau (Wang et al. 2019; Wu et al. 2019, 2020), including Jinsha River (Chen et al. 2013; Kong et al. 2009; Li et al. 2020), Lancang River (Tu and Deng 2020), and Yarlung Tsangpo River (Liu et al. 2015). Among these landslides, the mechanism of ancient ones is particularly complicated. Earthquakes and heavy rainfall are deemed to be the most possible triggers for large-scale ancient river-damming landslides (Dortch et al. 2009). However, in alpine valleys of the Tibetan Plateau, catastrophic slope failure initiated by buckling deformations also plays an important role in shaping landscapes.

Buckling often develops in layered rock slopes where rock slabs are approximately parallel to the hillslope (Cavers 1981). Slab weight, pore water pressure, and earthquake all could be potential triggers of buckling (Zhao et al. 2018). Many methods are used to analyze the modes of buckling, such as the Euler’s method, the limit equilibrium method, and numerical methods. Cavers (1981) developed a formula for different cases of buckling, including flexural buckling, three-hinge buckling of planar slopes and curved slopes. Tommasi et al. (2009) used the distinct element method to model buckling deformation of high natural slopes. Qi et al. (2015) proposed an analytical solution to analyze buckling failure caused by earthquake. Rosales Garzon (2016) used the tangent-modulus theory to study the mode of buckling failure. Zhao et al. (2018) analyzed the buckling mechanism of the Xinmo landslide using the Euler’s method and found that the landslide was triggered by precipitation.

This paper reports a study of a large ancient river-damming landslide related to buckling failure in the Tibetan Plateau. Detailed field mapping and theoretical analysis was conducted to reveal the potential mechanism of the landslide. The failure of this landslide involves large buckles with different deformation phases, which provide a good chance to study the deformation mechanism of layered rock mass under relatively well-known boundary conditions.

Geological setting

The study area is located at the southeast fringe of the Tibetan Plateau, approximately 50 km to the south of the Batang county, Sichuan province (Fig. 1a). The geomorphology is characterized by steep hillslopes with slope angles varying from 30° to 65° and deeply incised valleys. Several large-scale ancient river-damming landslides have been mapped in this region (Chen et al. 2013; Li et al. 2020), which have strongly modified the landscape.

Fig. 1
figure 1

modified from Cao et al. (2015))

Geomorphological and geological characteristics of the study area. (a) Topography. (b) Geological setting (

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Rocks exposed in this region mainly consist of metamorphic rocks from the Lower Permian period, including mica schist, quartz schist, and amphibolite, which are highly fractured due to tectonic movement. Slope stability in this area is mainly affected by two regional faults (Fig. 1b): (1) the Batang fault (F1) and (2) the Xiongsong-Suwalong fault (F2). The Batang fault trends NE30º and is approximately 200 km long. It is a late Holocene active left-lateral strike-slip fault with the slip rate of 2–4 mm/year, which triggered the Batang earthquake (Ms 7.5) in 1870 (Huang et al. 2015). The Xiongsong-Suwalong fault, trending approximately SN with a length of 300 km, is a branch of the Jinshajiang fault belt which experienced a right-lateral strike-slip movement of hundreds of kilometers since the Paleocene (Bao et al. 2020). The Xiongsong-Suwalong fault is also a late Holocene active fault and triggered the Wangdalong earthquake (Ms 6.5) in 1923. Based on historical records of earthquake events (Wu and Cai 1992; Li et al. 2020), seven strong earthquakes (Ms ≥ 6) have occurred in the study area since 1722, as shown in Fig. 1b. Earthquakes cause strong ground vibrations (Castelli et al. 2021), alter the stress conditions in slopes, and generate discontinuities in bedrocks, which strongly affect landslide activity (Ferraro et al. 2009; Valagussa et al. 2019; Li et al. 2020).

This area has a semi-arid climate. The temperature ranges from − 10 °C in December–January to 35 °C in July–August. Rainfall is concentrated in June to September, and the average annual precipitation is 468.3 mm (Chen et al. 2013; Li et al. 2020). Vegetation in this area is poorly developed.

Characteristics of the Xiaguiwa landslide

Landslide description

The Xiaguiwa landslide is located at the east bank of the Jinsha River (Fig. 2). It covers a surface area of 1.7 km2 with a width of 1000 m, a length of 1700 m and an estimated volume of 4.9 × 107 m3. The elevation of the scarp is approximately 3200 m, and the elevation of the bottom of the source area is 2510 m which is 65 m above the river level (2445 m).

Fig. 2
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Location of the Xiaguiwa landslide

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The valley of the Jinsha River in this region is formed by well stratified amphibolite with the interbed composed of mica schist, both in the Lower Permian age (Fig. 1b and Fig. 3). The polarized light microscopy based on thin sections was conducted for mineral identification (Zhu et al. 2011). The results show that the amphibolite mainly composed of hornblende (70 vol %), plagioclase (20%), and quartz (5%), developing a nematoblastic texture and a schistose structure. Besides, the mica schist mainly composed of muscovite (55 vol %), quartz (20%), plagioclase (10%), and biotite (10%), characterized by a lepidoblastic texture and a schistose structure. Field mapping shows that three sets of fractures are developed in the rock mass (Fig. 4): (1) fracture set 1 referring to bedding discontinuities with an average orientation of 245°∠41°, (2) fracture set 2 representing steeply dipping discontinuities with an average orientation of 169°∠78°, and (3)fracture set 3 with an average orientation of 51°∠40°. Bedding discontinuities are highly persistent, relatively planar and usually filled by clay or rock debris. They dip parallelly to the slope surface, with dip angles varying from 24° to 65°, divide amphibolite rocks into slabs (Fig. 3a), and thereby facilitate slope instability. Schist rocks are brittle and easy to be weathered, and the strength is sharply decreased when softened. Besides, the schist rock contains thin sericitization horizons parallel to bedding discontinuities, which are weaker than the bedding material (Fig. 3b).

Fig. 3
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Rocks exposed in the study area and their mineral compositions. Abbreviations: Pl-plagioclase; Qtz-quartz; Hbl-hornblende; Bt-biotite; Ms-muscovite; Ser-sericite

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Fig. 4
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Upper-hemisphere and equal-area projections of fracture orientations mapped from the studied rock slope. Set 1 refers to bedding discontinuities with an average orientation of 245°∠41°; set 2 refers to joints with an average orientation of 169°∠78°; set 3 refers to joints with an average orientation of 51°∠40°

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Detailed field investigations find that the Xiaguiwa landslide can be divided into 3 zones according to geomorphic features and deformation characteristics, as shown in Fig. 2. These zones are separated by small gullies. The upstream part of the landslide (zone 1) is characterized by planar rock surface, tension crack, buckle folds, and residual landslide debris overlaying on planar bedrocks. The downstream part of the landslide (zone 3) is characterized by steep head scarp, buckle folds, successive landslide deposits distributed below the scarp, and residual landslide dam in the river valley. In between (zone 2), the slope surface is strongly disturbed but without slope failure, characterized by a series of tension cracks trending parallelly to bedding discontinuities.

Zone 1

This zone is 0.8 km wide in the NW–SE direction and 1.0 km long (Fig. 5a). In the upper part of zone 1, bared bedrock is observed. There is a narrow platform with a width of about 5 m across the middle of this zone (Fig. 5b), which is assumed to be the hinge of a buckle fold. Below the platform, bared sliding surface is exposed, and three separate rock avalanches are deposited at the elevation of 2510–2750 m (Fig. 5a). The ground surface created by the landslide above the Zhuci road is wrinkled by buckle folds (Fig. 5a,c). The rock mass near the toe of the slope has suffered from gravitational deformations, and the dip angle of the strata are steeper than elsewhere, reaching up to 70°, as shown in Fig. 5d. Below the Zhuci road, most of landslide deposits has been washed away by the river, only a small bit of deposits is distributed at the slope toe (Fig. 5a). As shown in Fig. 5e, a large portion of the slope is subjected to large displacements due to buckling. The exposed slide surface of the ancient landslide has been deformed by gravity after the main slide event. Two small shallow and bowl-shaped rockslide scars have been left above the platform (Fig. 5a).

Fig. 5
figure 5

Characteristics of the slope in zone 1. (a) Geomorphology of zone 1. (b) Narrow platform in the middle of zone 1 caused by buckling. (c) Buckle fold exposed by excavation at the foot of zone 1. (d) Slip surface of the landslide and steeply dipping rock mass at the slope toe suffering from gravitational deformation. (e) Sketch map of profile I-I’

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Zone 2

Zone 2 is 0.4 km wide in the NW–SE direction and 0.8 km long. A series of tension cracks parallel to the bedding strike are exposed on the slope surface (Fig. 6a), indicating that the slope suffers from strongly polyphase gravitational deformations. Bedrocks outcrop around the extensional cracks, with the attitude of N30°W/45°SW (Fig. 6b). Field investigations show that the deformation of the slope in this zone has not reached equilibrium.

Fig. 6
figure 6

Characteristics of the slope in zones 2 and 3. (a) Geomorphology of zones 2 and 3. (b) Bedrock exposed at the foot of zone 2. (c) Buckle fold exposed by excavation at the foot of zone 3. (d) Sketch map of profile II-II’

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Zone 3

This region is 0.85 km long in the NE–SW direction and 0.5 km wide. The elevation of the top and the bottom of this landslide is 2970 m asl and 2510 m asl, respectively. Fan-like landslide deposits are distributed in the river valley at 2445–2520 m asl (Fig. 6a). The deposits once blocked the Jinsha River and formed a dammed lake. The height of the dam is estimated to be over 60 m based on the dam relic. Two successive deposits are observed (Fig. 6a): the landslide dam and a subsequent rockslide deposit overlying on the bedrock. Highway excavation cut through the toe of the landslide and a buckling fold was revealed (Fig. 6c), indicating the landslide is likely to be a toe buckling translational slide (Fig. 6d). Lacustrine sediments overlying on the landslide deposits are observed overlying on the landslide deposit (Fig. 6a). Optically stimulated luminescence (OSL) dating was performed on two samples obtained from the top of the lacustrine sediments in the front of the landslide dam (Fig. 6a). The samples were taken from freshly cleaned profiles using steel tubes. The OSL result shows that the landslide-dammed lake deposits at an age of 2.6 ± 0.2 ka.

Buckling deformations

The modes of buckling deformations can be grouped into two categories (Cavers 1981): three-hinge buckling and flexural buckling. The two kinds of buckling are observed in different regions of the study area in terms of the curvature and thickness of rock slabs.

Three-hinge buckling

Three-hinge buckling is observed at the upstream boundary of the Xiaguiwa landslide (Fig. 7a, b), as well as the platform in the middle part of zone 1 (Fig. 5b). Rock mass was cut into thin layers by a set of bedding discontinuities. The dip angle of bedding discontinuities progressively increases toward downhill, and the slope surface exhibits a gentle convex morphology (Fig. 7a). The pack of layers in three-hinge buckling is 3–5 m thick and the thickness of rock layers varies from 0.3 to 0.8 m.

Fig. 7
figure 7

Buckling deformations in the study area. (a) Three-hinge buckling at the upstream boundary of zone 1. (b) Photo showing the deformation of rock slabs in the threehinge buckling. (c) Flexural buckling at the foot of zone 1. (d) Flexural buckling near the foot of zone 3

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Flexural buckling

Flexural buckling is distributed at the foot of zones 1 and 3, with elevations varying between 2510 and 2535 m, as shown in Fig. 7c, d. The pack of layers in flexural buckling is 10–30 m thick. Rock layers within flexural buckling are 0.1 to 0.5 m thick, which is thinner than those within three-hinge buckling. The axis of buckle fold induced by flexural buckling is approximately parallel to the strike of rock layers and normal to the sliding direction.

Characteristics of lacustrine sediments

When the Xiaguiwa landslide occurred, massive rock avalanche rushed into the Jinsha River, generated a landslide dam over 60 m high below zone 3 and blocked the river (Fig. 6a). In the front of the dam, lacustrine sediments overly on landslide debris (Fig. 6a), indicating that they formed after the landslide. The lacustrine sediments are intermittently distributed along both banks, extending about 6.5 km upstream (Fig. 2). Characteristics of the lacustrine sediments caused by the Xiaguiwa landslide are shown in Fig. 8. The thickness of lacustrine sediments varies from 0.3 to 8 m. A special phenomenon was observed 1.8 km upstream from the landslide dam, as shown in Fig. 8d. Rock avalanche was buried by lacustrine sediments and appears as a locally phacoidal structure. We believe that three small landslides occurred in this area during the lacustrine sedimentation process. The debris of the landslides partially covered on the lacustrine deposits and afterwards was overlaid by subsequent sediments.

Fig. 8
figure 8

Profiles of sediments in the study area along the Jinsha river. (a) Lacustrine sediments resting on rock avalanche. (b) Lacustrine sediments overlying on bedrock. (c) Lacustrine sediments with planar laminations. (d) Photo showing a lacustrine-colluvial depositional system. (e) Rhythmically bedded silty clay composed of regularly alternating light-yellow and gray laminae. (f) Rock avalanche appearing as a phacoidal structure

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Particle gradation tests on two samples (samples 1 and 2) of lacustrine sediments were carried out. Note that sample 1 was obtained from the lacustrine sediments in the front of the landslide dam (Fig. 6a), and sample 2 was obtained from the lacustrine sediments 1.8 km upstream from the landslide dam (Fig. 8d).The results (Fig. 9) show that the lacustrine sediments are composed of silty clay with particle sizes of 0.002–0.25 mm, characterized by horizontally layered beddings (Fig. 8c), which is deemed to be typical lake sedimentary facies (Wu et al. 2019).

Fig. 9
figure 9

Grain size distribution of lacustrine sediments. Samples 1 and 2 were obtained from the lacustrine sediments in the front of the landslide dam, and 1.8 km upstream from the landslide dam, respectively

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Results and discussion

Generally, buckling mechanism requires that the dip of rock slabs is higher than the slope dip (Pereira and Lana 2012). The angle of internal friction, the slope dip and the discontinuity dip are controlling parameters for buckling deformations. Failure by buckling may be caused by the weight of the slab itself, by the increasing pore water pressure due to precipitation, by freezing pressures from ice, or by seismic shaking (Cavers 1981).

According to field investigations, rock formations in the study area are mainly composed of amphibolite in thin layers, characterized by crystalloblastic texture and schistose structure due to complicated tectonic activities. The amphibolite rocks are cut into planar slabs by bedding discontinuities that dip parallelly to the hillslope and have a steeper dip than the slope surface, which are prone to buckling. Thin-stratified mica schist outcrops as interlayers intercalated in the amphibolite formation, which is the weakest strata in the slope and provides a potential slip surface for the landslide. The environmental conditions including rainfall and seismic activities could be potential triggers for the Xiaguiwa landslide but are not well understood.

Potential mechanism of buckling

In this work, the Euler formula (Cavers 1981) is used to analyze the potential mechanism of buckling deformations in the study area. In the method, the critical loading Pcr for buckling without considering pore water pressure is calculated by (Hu and Cruden 1993)

$${P}_{cr}=\frac{{\pi }^{2}EI}{{4{L}_{b}}^{2}}$$
(1)

where E refers to the Young’s modulus, Lb means buckle length, and I is the moment of inertia computed by

$$I=\frac{{t}^{3}}{12}$$
(2)

where t refers to the thickness of buckled rock beds. The buckle force Pd caused by the weight of the upper portion of the rock beds is computed by (Hu and Cruden 1993)

$${P}_{d}=\left(L-{L}_{b}\right)t\left(\gamma\ \mathrm{sin}\ \alpha - \gamma\ \mathrm{cos}\ \alpha\ \mathrm{tan}\ \varphi\ -c/t\right)$$
(3)

where L, α, and γ are the total length, the dip and the unit weight of a rock bed, respectively; φ and c are the internal friction angle and cohesion along the bedding discontinuity, respectively. The factor of safety F is defined as (Hu and Cruden 1993)

$$F=\frac{{\pi }^{2}EI}{{4{L}_{b}}^{2}\left(L-{L}_{b}\right)\ t\ \left(\gamma\ \mathrm{sin}\ \alpha\ - \ \gamma\ \mathrm{cos}\ \alpha\ \mathrm{tan}\ \varphi\ - \ c/t\right)}$$
(4)

Then

$$\frac{\partial F}{\partial {L}_{b}}=\frac{{\pi }^{2}EI\left[2{L}_{b}\left(L-{L}_{b}\right)-{{L}_{b}}^{2}\right]}{4{{L}_{b}}^{4}{\left(L-{L}_{b}\right)}^{2}\ t\ \left(\gamma\ \mathrm{sin}\ \alpha\ - \ \gamma\ \mathrm{cos}\ \alpha\ \mathrm{tan}\ \varphi\ - \ c/t\right)}$$
(5)

Solving Eq. 4 = 0 gives Lb = 2L/3, suggesting that F achieves minimum when Lb = 2L/3. Assuming F = 1, the critical slope length for buckling Lcr can be determined:

$${L}_{cr}=0.25{\left(\frac{{9\pi }^{2}{Et}^{2}}{\left(\gamma\ \mathrm{sin}\ \alpha\ - \ \gamma\ \mathrm{cos}\ \alpha \ \mathrm{tan}\ \varphi\ - \ c/t\right)}\right)}^{1/3}$$
(6)

Parameters used for calculation are obtained from previous studies (Duan 2009; Du 2015) in the study area (Table 1). Field investigations show that the thickness of rock beds ranges from 0.04 to 1.00 m, with an average of 0.23 m, as shown in Fig. 10. The calculated results (Fig. 11a) show that the critical length of rock beds Lcr in profile I-I’ varies from 32 to 278 m corresponding to the bed thickness ranging from 0.04 to 1.00 m, both smaller than the actual length of rock beds (450 m). Therefore, rock beds are long enough to buckle under self-weight. Besides, the length of rock beds in profile II-II’ (405 m) is also longer than the critical length for buckling (40 to 340 m), suggesting that the rock bed in profile II-II’ can buckle under gravity.

Table 1 Parameters used in the analysis
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Fig. 10
figure 10

Distribution of bedding thickness

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Fig. 11
figure 11

Buckling analysis using the Euler’s method

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Assuming c is equal to zero, the relationship between buckle length Lb and bed thickness d under the critical state (F = 1) is represented by

$$d=\frac{{4L}_{b}}{\pi }{\left(\frac{3\gamma \left(L-{L}_{b}\right)\left(\mathrm{sin}\ \alpha\ - \ \mathrm{cos}\ \alpha\ \mathrm{tan}\ \varphi \right)}{E}\right)}^{1/2}$$
(7)

This relation can be delineated by Fig. 11b. The maximum thickness of rock beds for buckling in profiles I-I’ and II-II’ is 2.06 m and 1.30 m, respectively. Since the bed thickness of the slope in the study area varies between 0.04 and 1.00 m, which is smaller than the maximum bed thickness for buckling, we could conclude that the rock beds are thin enough to buckle under self-weight.

Potential causes of slope failure

The study area experienced complicated geological processes including uplift combined with fluvial undercutting, earthquakes, and rainfall. Therefore, determination of ancient landslide triggers in this area is difficult. The most likely triggers of landslides for a dip rock slope include rock mass properties, pore water pressure and earthquakes (Costa and Schuster 1988; Barnard et al. 2001; Cavallaro et al. 2019; Dortch et al. 2009; Tu and Deng 2020).

The information about pore water pressure and possible earthquakes when the Xiaguiwa landslide occurred, however, is unknown. In this work, a highly idealized sliding block model is used to analyze the possible causes of the landslide. The factor of safety F for a dip slope can be calculated by (Dortch et al. 2009).

$$F=\left(1-r\right)\frac{\mathrm{tan}\varphi }{\mathrm{tan}\beta }$$
(8)

where r represents the coefficient of pore- or cleft-water pressure coefficient; β and φ refer to the dip and the internal friction angle of the potential sliding surface, respectively. Cohesion along the sliding surface is assumed to be zero. The value of r ranges from 0 to 0.4 for non-artesian conditions, while the value of φ generally changes between 25° and 45°, relying on the characteristic of discontinuities and lithology (Dortch et al. 2009). Therefore, three stability fields can be defined for a slope, as shown in Fig. 12: (1) slopes falling within stability field 1 are unconditionally unstable; (2) slopes falling within stability field 2 will fail due to the increase of pore-water pressure; and (3) slopes falling within stability field 3 needs additional forces such as earthquakes to fail.

Fig. 12
figure 12

modified from Dortch et al. (2009))

The Xiaguiwa landslide plotted on idealized stability fields (

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The internal friction angles along the beddings of mica schist vary from 24° and 31° (Wang 2014). The dip of rock beds in profiles I-I’ and II-II’ are 40° and 35°, respectively. Thus, the slope in zones 1 and 2 is completely plotted within stability field 1 (Fig. 12), indicating that it can fail by buckling just under gravity. Based on the idealized model, we could conclude that the Xiaguiwa landslide was likely to start as gravitationally induced buckling of rock slabs, which agrees well with the results obtained by the Euler’s formula. More specifically, the landslide may initiate as a translational sliding along the weak interlayer composed of mica schist at the upper part of the slope and then formed buckles by curving amphibolite rock beds near the slope toe. This conclusion is supported by field investigations: (1) residual mica schist with a thickness of 1.2 m, which is strongly weathered, outcrops below the platform of zone 1 (Fig. 13a), which is deemed to be an ideal slip surface for the landslide; (2) strongly weathered mica schist is exposed on the slip surface at the downstream boundary of zone 1(Fig. 13b); and (3) landslide deposits composed of mica schist is exposed by road excavation near the toe of zone 3 (Fig. 13c).

Fig. 13
figure 13

Mica schist outcrops in the slope. (a) Mica schist below the platform of zone 1. (b) Mica schist distributed on the slip surface at the downstream boundary of zone 1. (c) Landslide deposits composed of mica schist exposed near the toe of zone 3

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Earthquake or pore water pressure, however, cannot be excluded as a potential trigger for the Xiaguiwa landslide. Affected by the two late Holocene active faults, the Xiongsong-Suwanglong fault and the Batang fault, the study area has experienced several strong earthquakes (Ms ≥ 6) on centennial to millennial timescales. These earthquakes are responsible for many landslides in the Batang-Suwalong segment of the Jinsha River (Chen et al. 2013; Li et al. 2020). Field investigation shows that the mica schist interbed and rock mass structure are responsible for the slope instability. Toe erosion caused by rapid and continuous downcutting of the Jinsha River forces layered rock mass to intensively buckle. The slope surface after the landslide has been deformed by gravitational deformations, characterized by small landslide scars, tension cracks and buckle folds (Fig. 5 and Fig. 6). Therefore, the evolution of the river valley caused by large-scale landsliding has not achieved equilibrium.

Conclusions

This paper reports a study of a large ancient river-damming landslide related to buckling failure in the Tibetan Plateau. The landslide with an estimated debris volume of 4.9 × 107 m3 is located at the upper Jinsha River, SE Tibetan Plateau. It once formed a landslide dam over 60 m high and blocked the river. Lacustrine sediments, composed of silty clay with particle sizes of 0.002–0.25 mm, are intermittently distributed along both banks, extending about 6.5 km upstream. The OSL dating indicates that the lacustrine sediments have an age of 2.6 ± 0.2 ka. Field investigation shows that the landslide can be divided into three zones based on geomorphic features, and each zone has different characteristics that provide information on the mechanism of the landslide.

Theoretical analysis was performed to investigate the potential cause and mechanism of the landslide. The results suggest that the river-damming landslide is most likely triggered by buckling of planar rock slabs induced by gravity. The intercalation of mica schist interbeds in the amphibolite formation and the structural conditions are responsible for the slope instability. The landslide may start as a translational sliding along the weak interlayer composed of mica schist at the upper part of the slope and then formed buckles by curving amphibolite rock beds near the slope toe. Earthquake or pore water pressure, however, cannot be excluded as a potential trigger for the slope failure, given that the study area is in a seismically active zone and suffers from heavy precipitation during the rainy season. The hillslope has still been affected by deformational deformations, characterized by tension cracks, buckle folds, and small landslide scars distributed on the slope surface, suggesting that the evolution of the river valley caused by large-scale landsliding has not achieved equilibrium.