Failure mechanisms and stability analysis of deep-seated ...

Failure mechanisms and stability analysis of

deep-seated landslides in the northwestern Rift

escarpment, Ethiopia

Dissertation

submitted in partial fulfilment of the requirements

for the degree of Doctor rerum naturalium (Dr. rer. nat.)

to the

Faculty of Geosciences

Ruhr-Universität Bochum

by

Tesfay Kiros Mebrahtu

Supervisors

Prof. Dr. Stefan Wohnlich

Prof. Dr.-Ing. Michael Alber

Bochum 2020

Dedicated to my parents.

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kenntlich gemacht habe. Außerdem versichere ich, dass die vorgelegte

elektronische mit der schriftlichen Version der Dissertation übereinstimmt und

die Abhandlung in dieser oder ähnlicher Form noch nicht anderweitig als

Promotionsleistung vorgelegt und bewertet wurde.

Bochum, 26.10.2020

Ort, Datum Unterschrift

Abstract

Landslides and ground failures are among the common geo-environmental hazards

in many of the tectonically active hilly and mountainous terrains of Ethiopia. The

study area is located in the central-western highlands of Ethiopia forming a

spectacular escarpment along the margins of the southwestern Afar depression

that is tectonically active since the beginning of Cenozoic Era. Owing to this, the

Debre Sina area and its surrounding are a potentially unstable environment in

terms of slope failure leading to landslides and mass movements. Deep-seated

landslides are common in the Debre Sina area where the geo-structural setting

plays a key role in controlling the geometry of the failure surface and its

displacement. Particularly, the Yizaba Wein, Shotel Amba, Nib Amba, and Wanza

Beret localities are repeatedly affected by deep-seated landslides. Despite that,

urban and rural development is currently active in almost the entire area. The

main objective of this study is to understand the processes leading to initiation and

propagation of slope failure, the influencing factors, and the failure mechanisms.

The study was carried out using a multidisciplinary approach based on geological,

morpho-tectonic analysis, geomorphological, hydrogeological, hydrogeochemical,

isotopic, geophysical (seismic refraction and vertical electrical sounding) and

geotechnical (kinematic analysis and numerical modelling) investigations.

According to the results, the deep-seated landslides in the study area are strongly

controlled by geological structures coinciding with the regional trend of the rift

margin faults. The geophysical data indicates that the area is covered by

unconsolidated sediments and highly decomposed and weak volcanic rocks which

are susceptible to sliding when they get moist. The slip surface generally coincides

with the presence of highly fractured and saturated rocks. The results of the

hydrogeochemical and stable isotopes illustrate that the main causes of the

landslide are the steep topography and the pressure formed during precipitation,

which leads to an increased weight of the loose and weathered materials. The

heterogeneity of the geological materials and the presence of impermeable layers

embodied within the highly permeable volcanic rocks can result in the build-up of

hydrostatic pressure at their interface, which can trigger landslides. Intense

fracturing in the tilted basalt and ignimbrite beds can also accelerate infiltration

of water, resulting to the build-up of high hydrostatic pressure causing low

effective normal stress in the rock mass, giving rise to landslides. The numerical

stability analysis indicates that the slope stability of landslide prone hills in the

study area strongly depends on the saturation conditions and the seismic load. In

general, an integrated analysis of all acquired data indicates that the deep-seated

landslides are controlled by different predisposing factors: tectonic uplift;

geological-structural setting; complex morphology of the slope and its high relief

energy; presence of closely spaced normal fault segments with steep slope angles;

and deepening action of the streams (active erosion and gullying). The findings of

this study can be used to understand mechanisms of deep-seated landslides in

similar morphological, geological and tectonic settings.

Kurzfassung

Erdrutsche gehören zu den häufigsten Naturgefahren im tektonisch aktiven

Bergland Äthiopiens. Das Untersuchungsgebiet befindet sich im zentralen

westlichen Hochland Äthiopiens mit spektakulären Steilhängen entlang den

Rändern der südwestlichen Afar-Senke, die seit Beginn des Känozoikums

tektonisch aktiv ist. Das Gebiet rund um Debre Sina ist daher potentiell instabil

und für Erdrutsche anfällig. Besonders tiefsitzende Massenbewegungen sind im

Debre Sina Gebiet häufig, da dort die geo-strukturellen Gegebenheiten eine

Schlüsselrolle in der Geometrie der Rutschfläche und der Hangbewegung spielen.

Gerade die Gebiete Yizaba Wein, Shotel Amba, Nib Amba und Wanza Beret sind

regelmäßig durch tiefsitzende Erdrutsche nach Starkregen betroffen. Dennoch

schreitet die städtische und ländliche Entwicklung im gesamten Gebiet stetig

voran. Das Hauptziel der vorliegenden Arbeit ist das Verständnis der Prozesse, die

zur Auslösung und dem Fortschreiten des Versagensprozesses führen, der

beeinflussenden Faktoren und der Versagensmechanik. Die Studie verwendet

einen interdisziplinären Ansatz bestehend aus einer geologischen und morpho-

tektonischen Analyse, sowie geomorphologischen, hydrogeologischen,

hydrogeochemischen, isotopischen, geophysikalischen (Refraktionsseismik und

Vertikale Elektrische Sondierung) sowie geotechnischen (kinematische Analyse

und numerische Modellierung) Untersuchungen.

Die Ergebnisse der Studie zeigen, dass die tiefsitzenden Erdrutsche im

Untersuchungsgebiet maßgeblich von den geologischen Strukturen geprägt sind,

die dem regionalen Trend der Verwerfungen am westlichen Rand des Rift folgen.

Aus den geologischen und geophysikalischen Untersuchungen lässt sich ableiten,

dass in dem Gebiet von unkonsolidierten Sedimenten und stark zersetzten und

weichen vulkanischen Gesteine vorherrschen, welche unter Wassereinfluss

Rutschbewegungen begünstigen. Allgemein stimmt die Rutschfläche mit der

Anwesenheit von stark geklüfteten und gesättigten Gesteinen überein. Die

Ergebnisse der hydrogeochemischen und stabilen Isotope zeigen, dass die

Hauptursachen für den Erdrutsch die steile Topographie und der während der

Ausfällung gebildete Druck sind, der zu einem erhöhten Gewicht des lockeren und

verwitterten Materials führt. Die Heterogenität der geologischen Materialien und

das Vorhandensein undurchlässiger Schichten, die in den hochdurchlässigen

vulkanischen Gesteinen verankert sind, können zum Aufbau von hydrostatischem

Druck an deren Grenzfläche führen, was Erdrutsche auslösen kann. Intensive

Brüche in den geneigten Basalt- und Ignimbritschichten können auch die

Infiltration von Wasser beschleunigen, was zum Aufbau eines hohen

hydrostatischen Drucks führt, der eine niedrige effektive Normalspannung in der

Gesteinsmasse verursacht, was zu Erdrutschen führt. Die numerische

Stabilitätsanalyse zeigt, dass die Hangstabilität im Untersuchungsgebiet stark

von der Sättigung und der seismischen Belastung abhängt. Ganz allgemein zeigt

eine zusammenfassende Analyse aller gesammelten Daten, dass tiefsitzende

Erdrutsche durch unterschiedliche Vorbedingungen kontrolliert werden:

tektonische Hebung; geologische-struktur Rahmenbedingungen; komplexe

Morphologie der Hänge und ihre hohe Relieffenergie; nahe beieinanderliegende

abschiebende Segmente mit steilen Hangneigungen; Vertiefung der Flüsse durch

aktive Erosion. Somit erweitern die Ergebnisse dieser Studie das Verständnis

tiefsitzender Erdrutsche in ähnlichen morphologischen, geologischen und

tektonischen Gebieten.

Acknowledgements

It is a pleasure to express my gratitude to those who have contributed to the

completion of this thesis. First of all, I am very thankful for the German Academic

Exchange Services (DAAD) for providing me the opportunity to pursue my PhD

study in Germany and its financial assistance.

I would like to express my deepest gratitude to Prof. Dr. Stefan Wohnlich for

providing me an opportunity to conduct doctoral research in his group as well as

for the constant scientific support, encouragement and kind advice he has provided

throughout my time as his student. Throughout my work, his door was always

opened to me when I asked him for advice. So, thank you so much Professor for

your continuous support during the whole journey.

I would also like to extend my gratitude to Prof. Dr.-Ing. Michael Alber for his

continuous scientific contributions to this PhD research work and all his guidance

and kind support throughout the project. Your knowledgeable and constructive

insights, encouragement and your always open door is highly appreciated.

I am thankful to PD Dr. Andre Banning for his valuable scientific reviews of my

research outputs and his continuous support. Thanks a lot, Andre. I am also

grateful to Dr. Thomas Heinze for his good encouragements, discussions, critically

read about my manuscripts and providing helpful comments. It is a joy to work

with this young scientist.

I have special thanks to Dr. Bedru Hussien for sharing me his research skill and

invaluable assistance and helped me during the fieldwork in Ethiopia. My sincere

thanks is extended to Dr. Ermias Hagos for his constant encouragement, generous

reviewing some of the sections in this thesis and insightful comments. Many

thanks also goes to Dr. Tesfaye Asresahagne (General Manager of Geomatrix Plc)

for providing me field logistics during the seismic refraction survey.

I would like to thank the Ruhr University Bochum Research School (RUB-RS), not

only for providing the funding which allowed me to undertake fieldwork, but also

for giving me the opportunity to attend international conferences and meet so

many interesting people. I am also grateful to the kind support and friendly from

Dr. Ursula Justus and Dr. Sarah Gemicioglu and other staffs of the RUB-RS. I

would like to thank the Wilhelm and Günter Esser Foundation for granting me a

scholarship to complete my dissertation.

I am thankful to Dr. Ferdinand Stöckhert, Claudia Brajer, Cedric Solibilda,

Kirsten Bartmann for their technical help and keeping a nice working atmosphere

in the rock mechanics laboratory of the RUB. Many thanks to the technical staffs

of Hydrogeology group, particularly Richard Nicolaus and Oliver Schübbe for their

patience and great effort they put to analyze the chemistry of water samples in the

hydrochemistry laboratory of the RUB. My special thanks to Isodetect

(Environmental Monitoring) in Munich and the Department of Materials and

Earth Sciences at the Technical University of Darmstadt for analysing the isotope

data used in this study.

I would like to thank all my fellow PhD students and staffs of the Applied Geology

Department for keeping a nice and friendly working atmosphere which is a key to

finish the journey. The cheerful and learning time with all other colleagues at the

RUB will stay in my heart forever.

Thanks are due to the Geological Survey of Ethiopia (GSE), National Meteorology

Agency of Ethiopia (NMA), Ethiopian Mapping Agency (EMA), Ethiopian Ministry

of Mines and Petroleum (MoMP) and Ethiopian Ministry of Water and Energy

(MoWE) for their generous provision of secondary data and support letters that are

enormously important for this research project.

Last but not least, a special word of thanks to my family and friends who kept in

touch with me all the time and encouraged me to finish the thesis. I am indebted

to them for their help. Thank you to all whose names are not mentioned who

supported me in any way during this study.

Above all, I give glory to God for his blessings, protection and love.

“If you can’t fly then run, if you can’t run then

walk, if you can’t walk then crawl, but whatever

you do you have to keep moving forward.”

Dr. Martin Luther King Jr.

Table of Contents

I

Table of Contents

Table of Contents ....................................................................................................... I

List of Figures ......................................................................................................... VI

List of Tables .......................................................................................................... XII

List of Acronyms ................................................................................................... XIII

List of Symbols ...................................................................................................... XVI

Chapter 1 ................................................................................................................... 1

1 Introduction ........................................................................................................ 1

1.1 Background .................................................................................................. 1

1.2 Regional geological and tectonic setting ..................................................... 5

1.3 Landslide types and their failure mechanisms ........................................ 12

1.3.1 Slides .................................................................................................... 13

1.3.2 Falls ..................................................................................................... 13

1.3.3 Topples ................................................................................................. 14

1.3.4 Lateral spreads .................................................................................... 14

1.3.5 Flows .................................................................................................... 15

1.4 Problem statement ..................................................................................... 17

1.5 Research objectives .................................................................................... 19

1.5.1 Main objective ...................................................................................... 19

1.5.2 Specific objectives ................................................................................ 20

Table of Contents

II

1.6 Summary of methodology .......................................................................... 20

1.7 Structure of the Thesis .............................................................................. 23

Chapter 2 ................................................................................................................. 25

2 Predisposing and triggering factors of large-scale landslides in Debre Sina

area, central Ethiopian highlands .......................................................................... 25

Abstract ................................................................................................................ 25

2.1 Introduction ................................................................................................ 26

2.2 The study area ........................................................................................... 27

2.3 Materials and methods .............................................................................. 29

2.4 Results ........................................................................................................ 31

2.4.1 Geology and geomorphology of the study area ................................... 31

2.4.2 Description, typology, and distribution of landslides ........................ 38

2.5 Discussion ................................................................................................... 45

2.5.1 Lithology and structure ....................................................................... 45

2.5.2 Elevation, slope angle, and aspect ...................................................... 49

2.5.3 Rainfall ................................................................................................ 50

2.5.4 Earthquakes ........................................................................................ 51

2.6 Conclusions ................................................................................................ 54

Chapter 3 ................................................................................................................. 56

3 Tectonic conditioning revealed by seismic refraction facilitates deep-seated

landslides in the western escarpment of the Main Ethiopian Rift ....................... 56

Abstract ................................................................................................................ 56

Table of Contents

III

3.1 Introduction ................................................................................................ 58

3.2 The study area ........................................................................................... 60

3.3 Materials and methods .............................................................................. 61

3.3.1 Instrumentation and field procedures ................................................ 62

3.3.2 Data acquisition, processing, and presentation ................................. 64

3.4 Results ........................................................................................................ 68

3.4.1 Geology and geomorphology of the study area ................................... 68

3.4.2 Profile one and two (L1–L2) ................................................................ 72

3.4.3 Profile 3 (L3) ........................................................................................ 76

3.4.4 Kinematic analysis of slope failure ..................................................... 77

3.5 Discussion ................................................................................................... 82

3.6 Conclusions ................................................................................................ 89

Chapter 4 ................................................................................................................. 92

4 The effect of hydrogeological and hydrochemical dynamics on landslide

triggering in the central highlands of Ethiopia ..................................................... 92

Abstract ................................................................................................................ 92

4.1 Introduction ................................................................................................ 94

4.2 The study area ........................................................................................... 96

4.2.1 Geological setting ................................................................................ 99

4.3 Materials and methods ............................................................................ 102

4.3.1 Water sampling and analytical methods .......................................... 102

Table of Contents

IV

4.3.2 Geophysical survey ............................................................................ 104

4.4 Results and discussion ............................................................................. 107

4.4.1 Aquifer system and groundwater flow ............................................. 107

4.4.2 Hydrogeochemical facies ................................................................... 110

4.4.3 Mechanisms controlling water chemistry ........................................ 113

4.4.4 The implications of groundwater dynamics with landslides ........... 114

4.4.5 Evidence from isotopic signatures .................................................... 118

4.4.6 Vertical Electrical Sounding ............................................................. 122

4.5 Conclusions .............................................................................................. 128

Chapter 5 ............................................................................................................... 130

5 Slope stability analysis of deep-seated landslides using Limit Equilibrium and

Finite Element methods under static and seismic load in Debre Sina area,

Ethiopia……………………………………………………………………………………130

Abstract .............................................................................................................. 130

5.1 Introduction .............................................................................................. 131

5.2 Geology of the area ................................................................................... 134

5.3 Methods and materials ............................................................................ 137

5.3.1 Model generation ............................................................................... 138

5.3.2 Limit equilibrium analysis ................................................................ 139

5.3.3 Finite element analysis ..................................................................... 141

5.4 Results and discussion ............................................................................. 145

5.4.1 Limit equilibrium analysis ................................................................ 145

Table of Contents

V

5.4.2 Finite element analysis ..................................................................... 148

5.5 Conclusions .............................................................................................. 155

Chapter 6 ............................................................................................................... 158

6 Summary and future research perspectives .................................................. 158

6.1 Summary .................................................................................................. 158

6.2 Future research perspectives .................................................................. 163

Declaration of authorship ..................................................................................... 165

References ............................................................................................................. 167

Appendix ................................................................................................................ 189

List of Figures

VI

List of Figures

Figure 1.1: (a) Global landslide susceptibility map computed using slope, geology,

fault zones, road networks, and forest loss (Stanley and Kirschbaum, 2017); (b)

Global Landslide Catalog (2007–2016) showing the distribution of landslide

fatalities (Kirschbaum et al., 2015b). ....................................................................... 2

Figure 1.2: Landslide distribution of Ethiopia (modified from Woldearegay, 2013).

................................................................................................................................... 4

Figure 1.3: Stratigraphy of the Afar region (modified from Varnet, 1978 and

Beyene and Abdelsalam, 2005). ................................................................................ 7

Figure 1.4: Schematic geological cross sections across the western and southern

Afar margins (modified from Corti et al., 2015). The location of the cross-section

line shown in Fig. 1.5. ............................................................................................... 8

Figure 1.5: Digital elevation map of the Afar Region showing the main structural

divisions (modified from http:www.see.leeds.ac.uk/afar). ..................................... 10

Figure 1.6: 3D view of the Afar depression and the west and east flanking plateaus

(source: http://en.wikipedia.org/wiki/Image:AfarDrape.jpg). ................................ 11

Figure 1.7: Landslide classification based on the type of movement and material (Varnes,

1978 and Cruden and Varnes, 1996). ............................................................................. 16

Figure 2.1: Location map of the study area…………………………………………….28

Figure 2.2: Geological map of the study area. ....................................................... 33

Figure 2.3: Elevation map of the study area. ......................................................... 35

Figure 2.4: Slope angle map of the study area. ..................................................... 36

Figure 2.5: Slope aspect map of the study area. .................................................... 37

Figure 2.6: Landslide inventory and morphostructural map of the study area. .. 39

List of Figures

VII

Figure 2.7: (a) Translational slide occurred in 2005, (b) roto-translational rock slide

and rockfalls, (c) Yizaba Wein and Shotel Amba convex-concave landslides and (d)

rotational slide and earth flow dipping downslope towards Dem Aytemashy river.

................................................................................................................................. 40

Figure 2.8: (a) Debris flow demolished agricultural land in Nib Amba, (b) rock slide

in Nib Amba, (c) deep-seated rotational slides form a pond at the lower part of the

slide zone and (d) earth flow demolished farmland. .............................................. 42

Figure 2.9: (a, b) Earth slides around Nech Amba area occurred on May 6, 2016,

(c) pre-existing landslide scars and active landslides in the gorge of the Majete

river and (d) a quasi-rotational slide widening retrogressively. ........................... 43

Figure 2.10: (a) Photographs showing rock slides around Armaniya along the

asphalt roadside, (b) earth slides, (c) tension cracks in a black cotton soil at Shola

Meda, (d) asphalt road collapsed along Debre Sina and Armaniya. ..................... 45

Figure 2.11: Stereographic projection of joints/fractures orientation data: (a) rose

diagram showing strike direction, (b) rose diagram showing dip direction, (c) plots

of poles and (d) pole density contour diagram. ...................................................... 46

Figure 2.12: Lineament map of the study area. .................................................... 48

Figure 2.13: Rainfall data from the Debre Sina station from 1974 to 2016 compared

with landslide events. ............................................................................................. 51

Figure 2.14: Recorded earthquakes in the East African region from 1842 to 2011

(source: EAGLE data). ............................................................................................ 53


Figure 3.2: The layout of the seismic refraction survey. ....................................... 64

Figure 3.3: Time-distance curve along profile one and two (L1–L2) that black line

is the fit. ................................................................................................................... 67

List of Figures

VIII

Figure 3.4: Time distance curve along profile three (L3) that black line is the fit.

................................................................................................................................. 67

Figure 3.5: Geological and geomorphological map of the study area. ................... 70

Figure 3.6: Panoramic view of the main Yizaba Wein landslide and surroundings

from the east located in north direction of Debre Sina area. ................................ 71

Figure 3.7: Typical landslides in the study area: (a) translational slides in

porphyritic-agglomeratic basalt and seismic refraction line L1–L2 location, (b)

rotational slides on colluvial deposit and volcanic ash/tuff, (c) rock slides in

porphyritic basalt and seismic refraction line L3 location, (d) earth slide in clay soil

and colluvial deposit. .............................................................................................. 74

Figure 3.8: Geological cross section along selected line A–B. ............................... 74

Figure 3.9: Seismic refraction tomography 2D P-wave velocity cross-section along

profile L1–L2. .......................................................................................................... 75


profile L3. ................................................................................................................ 77

Figure 3.11: The main Yizaba Wein landslide located in north direction of Debre

Sina town: fault 1–green (NNE–SSW), fault 2–blue (NNW–SSE) and fault 3–

purple (WSW–ENE). ............................................................................................... 78

Figure 3.12: Stereonet of planar sliding kinematic analysis. ............................... 80

Figure 3.13: Stereonet of wedge sliding kinematic analysis. ................................ 81

Figure 3.14: Rose diagram showing strike direction. ............................................ 82

Figure 3.15: Delta-t-V inversion along profile one and two (L1–L2). ................... 84

Figure 3.16: Delta-t-V inversion along profile three (L3). ..................................... 86


List of Figures

IX

Figure 4.2: (a) Panoramic view of the main Yizaba and Shotel Amba landslides

from east, with examples of characteristic geodynamic features within the main

landslide body and its surroundings: (b) rotational slide, (c) rock slide, (d)

debris/earth slide, (e) debris flow, (f) earth flow, (g) translational slide occurred in

2005 and (g) large-scale sliding. ............................................................................. 99

Figure 4.3: Geological map of the study area (modified from Mebrahtu et al.,

2020a). ................................................................................................................... 101

Figure 4.4: Groundwater level map and groundwater flow directions based on

spring and river positions. .................................................................................... 108

Figure 4.5: Piper diagram showing compositions of different water types in the

study area. ............................................................................................................. 111

Figure 4.6: A schematic cross section (W–E), showing the hydrogeological

conceptual model of the Debre Sina landslide. The location of the cross section and

its view direction is shown in Fig. 4.3. ................................................................. 112

Figure 4.7: Gibbs diagrams for (a) cations and (b) anions indicating rock-water

interaction as the major process regulating the chemistry of the groundwater in

the study area. ....................................................................................................... 113

Figure 4.8: Categorization of the water samples resulting from a preliminary

hierarchy cluster analysis (HCA) based on major ions chemistry using the complete

linkage rule and Euclidean distances. ................................................................. 115

Figure 4.9: Pictures of typical landslide localities in the Debre Sina area: (a)

emerging springs in ignimbrite-volcanic ash/tuff, (b) spring water at the contact of

the top layer (colluvium) and underlying altered tuff, (c) seepage spring at the

highly fractured ignimbrite, (d) spring water outflows from the bottom of the

landslide and (e) ponded spring water at the toe of the landslide. ..................... 117

Figure 4.10: (a) Cross plot of δ18O versus δ2H of the water samples with the Addis

Ababa LMWL and the GMWL, (b) isotopic altitude effect of precipitation of the

List of Figures

X

study area and (c) cross plot of 18O versus electrical conductivity (EC) of the study

area. ....................................................................................................................... 120

Figure 4.11: Mean monthly rainfall of the area for the last 43 years (1974 to 2016)

and mean monthly rainfall for the years 2005, 2006, 2007, 2014 and 2016 for the

Debre Sina area. .................................................................................................... 121

Figure 4.12: (a) Geoelectrical section and (b) apparent pseudo-depth section along

profile line–1. ......................................................................................................... 124


profile line–2. ......................................................................................................... 125


profile line–3. ......................................................................................................... 127


2020a)………………………………………………………………………………………135

Figure 5.2: Panoramic view of the main Yizaba and Shotel Amba landslides from

east with examples of characteristic geodynamic features within the main

landslide body and its surroundings (modified from Mebrahtu et al., 2021): (a)

rotational slide, (b) rock slide, (c) debris/earth slide, (d) debris flow, (e) earth flow,

and (f) a quasi-rotational slide widening retrogressively with ponded spring water

at the toe of the Wanza Beret landslide. .............................................................. 136

Figure 5.3: Specimens prepared and tested under uniaxial, triaxial, and tensile

loading. .................................................................................................................. 138

Figure 5.4: (a) Slope cross-section and (b) discretized RS2 model of slope section

along the Shotel Amba section. ............................................................................ 144


along the Yizaba section. ...................................................................................... 144


along the Nib Amba section. ................................................................................. 145

List of Figures

XI


along the Wanza Beret section. ............................................................................ 145

Figure 5.8: 2D cross-section result from slide along the Shotel Amba section... 147

Figure 5.9: 2D cross-section result from slide along the Yizaba section. ........... 147

Figure 5.10: 2D cross-section result from slide along the Wanza Beret section. 148

Figure 5.11: (a) Finite element analysis for shear strain and (b) finite element

analysis for total displacement with maximum total displacement of 225 m along

the Shotel Amba section. ...................................................................................... 149


analysis for total displacement with maximum total displacement of 36.1 m along

the Yizaba section. ................................................................................................ 152



the Nib Amba section. ........................................................................................... 153



the Wanza Beret section. ...................................................................................... 154

List of Tables

XII

List of Tables

Table 1.1: Landslide classification based on types of material and mode of

movement (Cruden and Varnes, 1996). .................................................................. 12

Table 3.1: Parameters of seismic refraction used during fieldwork………………...64

Table 3.2: Kinematic analysis of planar and wedge failures. ............................... 81

Table 3.3: Classification of the various subsurface units on the basis of their

compressional velocities. ......................................................................................... 87

Table 4.1: Hydrochemical and isotope data of sampled groundwater and surface

water in the Debre Sina area. SP spring; R river; Ionic concentrations are

measured in mg/L………………………………………………………………………...105

Table 5.1: Material parameters for rock used in LE and FE models………………142

Table 5.2: Geomechanical parameters used for faults. ....................................... 143

Table 5.3: Calculated FS using LE and FE methods without horizontal seismic

coefficient (h= 0). ................................................................................................. 146

Table 5.4: Calculated FS using LE and FE with horizontal seismic coefficient (h=

0.2 and 0.3). ........................................................................................................... 151

List of Acronyms

XIII

List of Acronyms

AAS Atomic Absorption Spectrometry

AGC Automatic Gain Control

BED Boundary Element Method

BSM Bishop’s Simplified Method

BTS Brazilian Tensile Strength

CMER Central Main Ethiopian Rift

CMP Common Mid Point

DAAD German Academic Exchange Service

DFG German Research Foundation

EARS East African Rift System

EAGLE Ethiopia-Afar Geoscientific Lithospheric Experiment

EBC Ethiopian Broadcasting Corporation

EBCS Ethiopian Building Code Standard

EC Electrical Conductivity

EIGS Ethiopian Institute of Geological Surveys

EMA Ethiopian Mapping Agency

ETM+ Enhanced Thematic Mapper Plus

DEM Digital Elevation Model

FDM Finite Difference Method

FE Finite Element

FEM Finite Element Method

FS Factor of Safety

GIS Geographic Information System

GISP Greenland Ice Sheet Precipitation

GMWL Global Meteoric Water Line

GNIP Global Network of Isotopes in Precipitation

GSE Geological Survey of Ethiopia

GSI Geological Strength Index

HCA Hierarchical Cluster Analysis

IAEA International Atomic Energy Agency

List of Acronyms

XIV

IAEG International Association of Engineering Geology

ISRM International Society for Rock Mechanics

JSM Janbu’s Simplified Method

km kilometre

LE Limit Equilibrium

LEM Limit Equilibrium Methods

LMWL Local Meteoric Water Line

mg/L Milligram per Liter

m/s meter per second

ms milliseconds

mm millimeter

MER Main Ethiopian Rift

MoMP Ministry of Mines and Petroleum, Ethiopia

MoWE Ministry of Water and Energy, Ethiopia

MPM Morgenstern-Price Method

NMA National Meteorological Agency of Ethiopia

NMER Northern Main Ethiopian Rift

Pcc Piecewise cubic convolution

µS/cm Micro Siemens per centimeter

δ18O Isotope ratio of oxygen-18

δ2H Isotope ratio of hydrogen-2

‰ Per mil

PEM Point Estimate Method

PGA Peak Ground Acceleration

R River

RMS Root Mean Square

RS Remote Sensing

RUB Ruhr University Bochum

SID Station Identification

SLAP Standard Light Antarctic Precipitation

SM Spencer’s Method

SMER Southern Main Ethiopian Rift

List of Acronyms

XV

SMOW Standard Mean Ocean Water

SNNPR Southern Nations and Nationalities and People’s Region

SP Spring

SRF Strength Reduction Factor

SSR Shear Strength Reduction

SWL Static Water Level

TCS Triaxial Compressive Strength

TDS Total Dissolved Solids

TM Thematic Mapper

TS Total Station

UCS Uniaxial Compressive Strength

UTM Universal Transverse Mercator

VES Vertical Electrical Soundings

VSMOV Vienna Standard Mean Ocean Water

WET Wavepath Eikonal Traveltime

WFB Wonji Fault Belt

YTVL Yerer–Tullu Wellel Volcanotectonic Lineament

List of symbols

XVI

List of Symbols

Symbol description unit

E Young's modulus Pa

c cohesion Pa

c' effective cohesion Pa

Cr reduced cohesion Pa

Ei intact rock modulus Pa

Em rock mass modulus Pa

Fs Factor of safety -

h depth m

Kn joint normal stiffness Pa m-1

Ks joint shear stiffness Pa m-1

n number of velocity layers from intercept time -

Vn velocity of the nth layer m s-1

Vp P-wave velocity m s-1

vP ultrasonic pulse velocity m s-1

V1 first layer velocity m s-1

V1d inverse slope of the direct m s-1

V1r inverse slope of the reverse m s-1

V2 second layer velocity m s-1

V2d apparent up-dip velocity m s-1

V2r down-dip velocity m s-1

Tn nth intercept time s

ν Poisson’s ratio -

shear stress Pa

ρa apparent resistivity Ω-m

' effective normal stress on the surface of rupture Pa

h horizontal earthquake coefficient -

angle of internal friction °

' the effective angle of internal friction °

r reduced angle of internal friction °

Introduction

1

Chapter 1

1 Introduction

1.1 Background

Landslides are one of the most destructive natural hazards that pose a critical and

continuous threat to people around the world (Fig. 1.1). Landslides are caused by

different triggering factors such as heavy or prolonged precipitation, earthquakes,

rapid snow melting and a variety of anthropogenic activities. As Dai and Lee (2002)

report, natural hazards are believed to account for up to 4 % of the total annual

deaths worldwide, besides causing enormous economic losses and uprooting

habitation. Worldwide landslide activities are expected to continue in the 21st

century for the following reasons: (a) increased urbanization and development in

landslide-prone areas, (b) continued deforestation of landslide-prone areas, and (c)

increased precipitation caused by changing climatic conditions (Schuster, 1995).

As a result, the study of landslides has drawn global attention to increase

awareness about its socio-economic impacts and the pressure of increasing

population and urbanization on mountainous areas (Kanungo et al., 2006).

Landslides are also a problem in Ethiopia, and many parts of the country are prone

to natural and man-induced slope instability.

Ethiopia is located close to the active East African Rift System (EARS) which

results in numerous landslides in many parts of the country. Rainfall-induced

landslides are common problems in many areas of the hilly and mountainous

regions of the highlands of Ethiopia. Landslides are especially a common

phenomenon in the central highlands and rift escarpments of Ethiopia, which

brought a heavy impact on agricultural land, dwellers and infrastructure, and

often lead to the displacement and death of people. Landslides can be triggered by

both natural and man-induced changes in the environment. As Ayalew (1999)

mentioned, steep slopes, ''squeezed'' concave segments, highly jointed and

weathered rocks, and the action of both rivers and humans have created an

Introduction

2

environment that is conducive for triggering slope movements in many parts of

Ethiopia. People have moved into areas that are potentially threatened by slope

instability. Following this urbanization and expansion of construction into fragile

terrain without prior site investigation, these areas are being exposed to landslide

problems after rainfall during the rainy seasons (Temesgen et al., 2001; Abebe et

al., 2010; Woldearegay, 2013). There have been several landslide occurrences in

Ethiopia imposing considerable socio-economic problems.

Figure 1.1: (a) Global landslide susceptibility map computed using slope, geology,

fault zones, road networks, and forest loss (Stanley and Kirschbaum, 2017); (b)

Global Landslide Catalog (2007–2016) showing the distribution of landslide

fatalities (Kirschbaum et al., 2015b).

The highlands and mountainous areas of Ethiopia like the Debre Sina (Schneider

et al., 2008; Woldearegay, 2008), Kombolcha–Dessie road (EIGS, 1995; Ayenew

and Barbieri, 2005; Fubelli et al., 2008), Abay Gorge (EIGS, 1994; Ayalew and

Yamagishi, 2004), Jemma basin (Zvelebi et al., 2010), Goffa area (Asrat et al.,

Introduction

3

1996), Wollo area (Ayalew, 1999), Wondo Genet area (Temesgen et al., 2001),

Adishu area (Woldearegay et al., 2005) and many other parts of Ethiopia are

repeatedly facing problems associated with landslides (Fig. 1.2). The landslides in

these areas are affecting human lives, infrastructures, agricultural lands and the

natural environment.

In recent years, landslide incidences are increasing in the Ethiopian highlands due

to man-induced and natural causes, yet their causes are not well understood. For

instance, from 1993 to 1998, landslides or related ground movement problems

claimed about 300 lives, damaged over 100 km of asphalt road, demolished more

than 200 dwelling houses and also devastated more than 500 ha of agricultural

land in Ethiopia (Ayalew, 1999). Furthermore, 135 human lives were lost, about

3500 people were displaced and an estimated US$ 1.5 million worth of property

was damaged in the highlands of Ethiopia in the years 1998 to 2003 (Woldearegay,

2013). Even in the year 2019, a landslide that took place on 12 October 2019 after

10 hours of continuous rainfall in Konta special district of Southern Nations and

Nationalities and People’s Region (SNNPR) killed 23 people and demolished 5

dwelling houses according to a report by local media, Ethiopian Broadcasting

Corporation (EBC). Further, landslides in South Omo, SNNPR killed at least 8

people following heavy rains on April 18, 2020. There was also a landslide incident

in the South Omo zone (Ale Special Woreda) on April 30, 2020 following heavy rain;

12 people were killed, six houses completely demolished, and 2188 ha of farmland

were devastated. At least four people died while serval people were injured after

flash floods hit the city of Dire Dawa on April 24, 2020.

The margins of the western Afar depression are currently under threat with the

problem of landslides and mass movement. Out of many problematic regions, the

Debre Sina landslide is one of the largest deep-seated landslides in the country

(Figs. 1.1b and 1.2). The area is located in a tectonically active area with an

expansive character and has been affected by large-scale and deep-seated

landslides. Several landslides have occurred in the past and there are also

numerous evidences of active landslides in the study area. Particularly, as a result

of a single large-scale and deep-seated landslide that took place on 13 September

Introduction

4

2005 in the Debre Sina area, more than 3000 people have been displaced; 1250

dwelling houses and one elementary school demolished; four churches, four mills

and over 1500 ha of farmland were also severely disrupted (Woldearegay, 2008).

Figure 1.2: Landslide distribution of Ethiopia (modified from Woldearegay, 2013).

This landslide is probably one of the largest that occurred in recent times on the

East African continent. The landslide problem in the study area is very active and

still causing problems. There are many tension cracks which are developed in the

area, and these tension cracks are indications for probable some more slides to

occur in the near future. Therefore, in order to minimize such damages, a detailed

investigation of landslide-prone areas plays a crucial role. Adequate

characterization of landslides requires a deep understanding of causes and failure

mechanisms. This, in turn, requires a detailed study of the geological,

topographical and physical properties of rocks and soils that are found in unstable

slope profiles.

Regional geological and tectonic setting

5

1.2 Regional geological and tectonic setting

The geology of central Ethiopia is represented by three litho-stratigraphic units,

namely: (i) Precambrian crystalline rocks (ii) Mesozoic sedimentary rocks, and (iii)

Tertiary-Quaternary volcanic rocks with minor volcano-clastic sediments,

lacustrine sediments and superficial deposits (Fig. 1.3). Following the Late

Mesozoic-Early Tertiary transgression of the sea from the southeast, an epirogenic

uplift of Afro-Arabia occurred on an immense scale. The Ethiopian volcanic rocks

were divided into two main series: Trap Series or Plateau Series and Rift volcanic

(Mohr, 1971; Zanettin et al., 1974; Zanettin, 1993). The flood basalt successions in

the Ethiopian plateaus (northwest and southeast) formed during the period of 31–

28 Ma (Pik et al., 1998; Meshesha and Shinjo, 2007; Beccaluva et al., 2009).

In the central part of the northwestern plateau, the volcanic rocks are sub-divided

into Ashangi and Aiba basalts, Alaji Formation, and Tarmaber basalt (Zanettin et

al., 1974; Kazmin, 1979). The Ashangi basalt represents the earliest fissural flood

basalt volcanism consisting of predominantly mildly alkaline basalts with inter-

bedded pyroclastics, rare rhyolites and commonly injected by dolerite sills and

dykes (Zanettin and Justin-Visentin, 1974; Mengesha et al., 1996). The upper part

is more tuffaceous and contains interbedded lacustrine deposits with lignite seams.

The Aiba basalts consist of massive transitional flood-basalt flows, with

intercalated agglomerate beds. The Alaji Formation mainly consists of aphyric

flood basalts associated with rhyolite (ignimbrites) and subordinate trachytes

resting conformably on the Aiba basalts but in some places (e.g., Kesem and Muger

valleys as well as in most outcrops on the southeastern plateau) it directly lies on

top of the Mesozoic sediments (Mohr and Zanettin, 1988). This profuse volcanic

outpouring took place between 31 Ma and 29 Ma (Hofmann et al., 1997; Pik et al.,

1998; Ukstins et al., 2002; Coulié et al., 2003). This event was followed by shield-

volcano-building episodes from 23 Ma to 11 Ma (Kieffer et al., 2004). In the

southern part of the northwestern plateau, products of the shield-volcano were

followed by fissural eruption and grouped under Tarmaber-Megezez Formation.

They are made of lenticular, often zeolitized, alkali basalts with a large amount of

tuffs, scoriaceous lava flows, and peralkaline rhyolites with maximum thickness of


6

1,000 m close to the centers (Mohr and Zanettin, 1988). Along the rift escarpment

and plateau, voluminous basaltic rocks alternating with agglomerates and minor

silisic rocks form what is known to as the Trap Series (Kazmin, 1975). This

extensive basaltic rock (up to 1200 m thick) is believed to have erupted from

fissures during the middle Tertiary marking the proto-rift stage and initiation of

large-scale extensional movement affecting the horn of Africa. This stage is

terminated with gentle warping along the rift boundary accompanied with the

eruption of fissural silicic volcanics (rhyolites, ignimbrites and unwelded tuff)

(Zanettin et al., 1974).

The plateaus are covered by the oldest volcanic formation (collectively known as

the Trap Series) and consist of voluminous basaltic rocks alternating with

agglomerates and tuffs (Fig. 1.4). Dike swarms, acidic extrusions and typical red

paleosoils are also frequently found in association with the basalt. The Trap Series

is 200–1200 meters thick with the thickest section occurring in the proximity of

the rift escarpment (e.g., Mohr, 1967; Kazmin, 1975). This extensive formation

erupted from fissures during the early and middle Tertiary marking the initiation

of domal uplift (Proto-rifting stage) and large-scale extensional movements

affecting larger regions (Brotzu et al., 1986).

The stratigraphy of the volcanic rocks of the Main Ethiopian Rift (MER),

envisaging a lower basalt unit with trachybasalts and subordinate silicic flows (11

Ma to 8 Ma) followed by a widespread ignimbrite cover (Nazret Series) ranging in

age from 7 Ma to 2 Ma with an estimated thickness of 700 m (Corti, 2009). These

two units, common to the whole MER, are followed by Late Pliocene basalts with

pyroclastics fed by calderas which are limited to the northern and central sectors

(Fig. 1.4). The subsequent Quaternary volcanic unit, which outcrops throughout

the MER, is the Wonji Group associated with the oblique Wonji Fault Belt (Mohr,

1962). It includes basalt flows, scoria cones, and large silicic central volcanoes with

calderas experienced phreatomagmatic activity and historical flows. The rifting

stage is marked by shift in volcanism and tectonic activities from escarpment to

the axial zone (Kazmin and Berhe, 1978). This was followed by major faulting and

tilting of the escarpment and subsidence of the rift floor at around 4–5 Ma. This


7

was succeeded by the eruption of Bofa basalt (early rift floor basalt, Kazmin and

Berhe, 1978 and reference therein) which was interpreted as forming a

stratigraphic wedge between the Nazret Group and the Afar rift.

Figure 1.3: Stratigraphy of the Afar region (modified from Varnet, 1978 and

Beyene and Abdelsalam, 2005).

The MER is flanked by the Ethiopian Plateau in the west and the Somalian

Plateau to the southeast (Fig. 1.5). The MER is sub-divided into three main sectors

differing in trend, fault patterns and lithospheric characteristics (e.g., Mohr, 1983;

Hayward and Ebinger, 1996; Bonini et al., 2005) namely: (i) northern, (ii) central,

and (iii) southern sectors (Fig. 1.5). The EARS is one of the continental rifting,

which stretches for over 3000 km along its length. The MER is a segment of the

EARS. It extends for about 500 km in NE–SW direction within Ethiopia and

represents the link region between the Afar triple junction to the north and EARS


8

to the south (Fig. 1.5). The EARS marks the incipient plate boundary separating

Nubia and Somalia plates (Ebinger, 2005). The northern MER (NMER) extends

from the southern Afar depression in the northeast to the region of Lake Koka and

Gedemsa caldera (Fig. 1.5), showing a roughly NE–SW trend. The central MER

(CMER) extends southward in a rough N25°–30°E direction from the Lake Koka

up to Lake Awasa and the E–W Goba-Bonga tectonic lineament (Boccaletti et al.,

1998). The southern MER (SMER) extends southward of Lake Awasa up to the

overlapping region between the MER and the Kenya rift, characterized by a ~300

km-wide broadly rifted zone of basins and ranges (Moore and Davison, 1978;

Ebinger et al., 2000). The Debre Sina area lies along the rift escarpment of the

NMER (Fig. 1.5) straddling the Afar depression which have experienced rift

interplay between the Red Sea–Gulf of Aden and MER. The southern Afar rift,

where the study area situated, is a transition zone between the central Afar and

the MER. It is structurally characterized by north to northeast-trending dominant

structures in the west, and east–west trending in the east (Beyene and

Abdelsalam, 2005) and northwest-trending transfer fault zones which can be

traced to discontinuities in the western Ethiopian escarpment (Hayward and

Ebinger, 1996).

Figure 1.4: Schematic geological cross sections across the western and southern

Afar margins (modified from Corti et al., 2015). The location of the cross-section

line shown in Fig. 1.5.

The kinematically distinct Gulf of Aden normal faulting pattern (trending due to

east–southeast) found in the southern part (Tesfaye et al., 2003) and escarpment


9

with a length of about 250 km and an average crustal thickness of about 26 km. In

general, the three important structures namely: the NW–SE trending structures

(parallel to the general trend of the Red Sea); NE–SW trending structures (parallel

to the MER) and E–W trending (parallel to the Gulf of Aden) are joined in the

southern Afar rift (Fig. 1.6). The western bounding rift margins, where the study

area is located, is characterized by these three important regional structures

controlling the deep-seated landslides along the rift margins (e.g., the Debre Sina

landslide that occurred in 2005). The different MER sectors are characterized by

two distinct systems of normal faults that differ in terms of orientation, structural

characteristics (e.g., length, vertical throw), timing of activation and relation with

magmatism: (i) the border faults and (ii) a set of faults affecting the rift floor,

usually referred to as Wonji Fault Belt (e.g., Boccaletti et al., 1998; Mohr, 1962;

Gibson, 1969; Mohr and Wood, 1976). The border faults are normally long, widely

spaced, characterized by large vertical offset and variable orientation in the

different MER sectors. The Wonji Fault Belt (WFB) is a tectono-volcanic system

characterized by short, closely spaced, active faults that exhibit minor vertical

throw. The WFB is intimately associated with the intense Quaternary magmatism

of the rift floor.

These faults are well developed in the northern sector, where the WFB structures

form clearly defined right-stepping en-echelon segments obliquely cutting the rift

floor (Boccaletti et al., 1998). In addition to these, the Ethiopian rift shows an offset

around 8°30'N to 9°00'N latitudes marked by volcanoes and fracture systems

roughly trending east-west, termed as Yerer–Tullu Wellel Volcanotectonic

Lineament (YTVL) (Abebe et al., 1998; Mazzarini et al., 1999). The YTVL which

was traditionally considered distinct from the Cenozoic rift system and is

interpreted as an integral part of the MER evolution (Keranen and Klemperer,

2008).

Shift in tectonism from escarpment to the axial zone, enhanced by crustal thinning

and magma intrusion and subsequent strain softening produced major NNE

trending right lateral stepping en-echelon normal faults along discrete zones

known as the WFB. This event marks major ESE–WNW extension and associated

with extrusion of chains of basalt flows, scoria cones along the fault belt.


10

Figure 1.5: Digital elevation map of the Afar Region showing the main structural

divisions (modified from http:www.see.leeds.ac.uk/afar).

The study area is located in the southern Afar rift along the border zone of the

Ethiopian escarpment and the MER (Fig. 1.5). The western escarpment is

remarkably an elevated area with steep slope to the east marking the western

boundary of southern Afar depression. Internally it is generally rugged with

alternating hills and valleys along its strike (NNE–SSW). The EARS is 40–65 km

wide, generally N–S trending and extends for more than 3000 km from the Red


11

Sea region in the north to Tanzania further south (Baker et. al., 1972). Its

geodynamic evolution started in early Tertiary time and continued to the present

with episodic uplifting, volcanic activity and an associated fluvial-lacustrine

sedimentation along contemporary asymmetrical grabens formed in response to

enhanced instantaneous tectonic extension (e.g., Mohr, 1986; Chorowicz et. al.,

1987).

Figure 1.6: 3D view of the Afar depression and the west and east flanking plateaus

(source: http://en.wikipedia.org/wiki/Image:AfarDrape.jpg).

The Afar and MER segments are main sites where the lithosphere has

undergone the greatest amount of thinning (Fig. 1.6). These grabens traverse

the two broadly elongated western and eastern plateaus. Their development

has been attributed to the occurrence of a mantle hot spot beneath the uplifted

continental plateaus (e.g., Hofmann et. al., 1997; George et. al., 1998; George

and Rogers, 2002). The southern Afar region indicate the occurrence of

various kinds of volcanic rocks and volcanoclastic sediments related to two

major sequential stages of tectonic uplift and associated magmatism (e.g.,

Mohr, 1967; Merla et al., 1979; Brotzu et. al., 1986). In general, this extensive

tectonic activity is considered to be responsible for the development of the

http://en.wikipedia.org/wiki/Image:AfarDrape.jpg

Landslide types and their failure mechanisms

12

litho-structural relationships and present day morphological appearance of

the area. This has significant implications on the evolution of the deep-seated

landslides.

1.3 Landslide types and their failure mechanisms

Landslides are very diverse phenomena in shape and size, movement speed and

other characteristics. Many classifications have been proposed for landslides based

on the type of material, type of movement, causes, and many other factors. The

most widely used classification is the one developed by Varnes (1978), which takes

into account both the type of material and the type of movement in combination

for the classification of landslides into different types. This classification

distinguishes five types of mass movement (slides, falls, topples, spreads, and

flows) and combinations of these principal types along with different types of

material (bedrock, coarse soils, and predominant fine soils) (Fig. 1.7). The most

common classification for landslides is based on material properties and process

types (Table 1.1). Besides the main types of movement processes, there is one

complex class which contains movement processes with two or more different

processes acting together along with downslope movement of the landslide mass.

The most common types of landslides are described as follows and are illustrated

in Fig. 1.7.

Table 1.1: Landslide classification based on types of material and mode of

movement (Cruden and Varnes, 1996).

Process type Type of material

rock debris earth

Topple rock topple debris topple earth topple

Fall rock fall debris fall earth fall

Slide translational

rock slide debris slide earth slide rotational

Flow rock flow debris flow earth flow

Spread rock spread debris spread earth spread

Complex e.g., rock avalanche e.g., flow slide e.g., slump-earthflow


13

1.3.1 Slides

A slide is a downslope movement of soil or rock mass occurring predominantly on

surfaces of rupture or on relatively thin zones of intense shear strain (Cruden and

Varnes, 1996). The slide can be rock-slides or debris-slides when rocks or debris

slide down a pre-existing surface, such as a bedding plane, foliation surface, or a

joint surface (Fig. 1.7). Sliding mass may or may not experience considerable

deformation and could be rotational, translational or a combination of both, which

is called a compound slide (Bell, 1999). Rotational slide is a slide in which the

surface of rupture is curved concavely upward and the slide movement is roughly

rotational about an axis that is parallel to the ground surface and transverse across

the slide (Varnes, 1978). The head of the displaced material may move almost

vertically downward, and the upper surface of the displaced material may tilt

backwards toward the scarp (Highland and Bobrowsky, 2008). If the slide is

rotational and has several parallel curved planes of movement, it is called a slump.

Translational slide occurs when the mass displaces along a planar or undulating

surface of rupture, sliding out over the original ground surface (Cruden and

Varnes, 1996). The scale of rock slides could range from small-scale discontinuity

controlled plane or wedge failures to large-scale failures. A block slide is a

translational slide in which the moving mass consists of a single unit or a few

closely related units that move downslope as a relatively coherent mass (Fig. 1.7).

According to various authors (e.g., Terzaghi, 1950; Goodman and Kieffer, 2000),

the factors that governing large-scale slope stability are mainly: (a) stress

conditions, including the effects of water, (b) geological structures, particularly the

presence of large-scale features, (c) geometry of the slope, and (d) rock mass

strength. Failure modes in large-scale rock slope instabilities could be planar

shear, wedge failures or quasi-rotational shear failures.

1.3.2 Falls

Falls are abrupt downward movements of masses of geologic materials, such as

rocks and boulders, that become detached from steep slopes or cliffs (Fig. 1.7).

Separation occurs along discontinuities such as fractures, joints, and bedding


14

planes, and movement occurs by free fall, bouncing, and rolling. Falls are strongly

influenced by gravity, mechanical weathering, and the presence of interstitial

water. A fall starts with the detachment of soil or rock from a steep slope along a

surface on which little or no shear displacement takes place. The material then

descends mainly through the air by falling, bouncing, or rolling (Cruden and

Varnes, 1996). Fall movement is very quick, and typically involves slope angles

range from 45° to 90° and includes rock falls, debris falls, and earth falls (Fig. 1.7).

The falling material usually strikes the lower slope at angles less than the angle

of fall, causing bouncing. The falling mass may break on impact, may begin rolling

on steeper slopes, and may continue until the terrain flattens. The effects of

weathering, such as the freezing of water in joints (in cold countries), the pressure

of water in fissures, and root pressures may initiate failure in the weak rocks.

1.3.3 Topples

A topple is the forward rotation out of a slope of a mass of soil or rock around a

point or axis below the center of gravity of the displaced mass (Fig. 1.7). Toppling

is sometimes driven by gravity exerted by material upslope of the displaced

mass and sometimes by water or ice in cracks in the mass (Cruden and Varnes,

1996). Topples may lead to falls or slides of the displaced mass, depending on

the geometry of the moving mass, the geometry of the surface of separation,

and the orientation and extent of the kinematically active discontinuities.

Topples can consist of rock, debris (coarse material), or earth materials (fine-

grained material). Topples can be complex and composite. Topples range from

extremely slow to extremely rapid, sometimes accelerating throughout the

movement.

1.3.4 Lateral spreads

Lateral spreading is defined as an extension of a cohesive soil or rock mass

combined with a general subsidence of the fractured mass of cohesive material into

softer underlying material (Cruden and Varnes, 1996). The dominant mode of

movement is lateral accommodated by shear or tensile fractures (Varnes, 1978).

Lateral spreads involve the horizontal displacement of the surface and are


15

distinctive because they usually occur on very gentle slopes or flat terrain (Fig.

1.7). Loose cohesionless sediments commonly produce lateral spreads through

response to earthquake vibrations. The movement of lateral spreading is usually

complex, being predominantly translational, but also show rotational movement

and liquefaction, and consequent flow may also be involved (Varnes, 1978; Bell,

1999). The failure is caused by liquefaction, the process whereby saturated, loose,

cohesionless sediments (usually sands and silts) are transformed from a solid into

a liquid state. The lateral spread is controlled by different triggering mechanisms

such as (i) liquefaction of lower weak layer by earthquake shaking (ii) natural or

anthropogenic overloading of the ground above an unstable slope, (iii) saturation

of underlying weaker layer due to precipitation, snowmelt, and (or) groundwater

changes, (iv) liquefaction of underlying sensitive marine clay following an erosional

disturbance at base of a riverbank/slope, and (v) plastic deformation of unstable

material at depth (e.g., salt).

1.3.5 Flows

Flows are rapid movements of material as a viscous mass where inter-granular

movements predominate over shear surface movements and these can be debris

flows, mud flows or rock avalanches (Fig. 1.7), depending upon the nature of the

material involved in the movement (Varnes, 1978). They are distinguished from

slides by having higher water content and are thoroughly deformed internally

during movement (Hutchinson, 1995). The distribution of velocities in the

displacing mass resembles that in a viscous fluid. A flow is a spatially continuous

movement in which surfaces of shear are short-lived, closely spaced, and usually

not preserved. The two major types of flows are debris flows and earth flows. There

are five basic categories of flows (debris flow, debris avalanche, earth flow, mud

flow and creep) that differ from one another in fundamental ways (Highland and

Bobrowsky, 2008). The debris flows and earth flows are briefly outlined below.

A debris flow is a form of rapid mass movement in which a combination of loose

soil, rock, organic matter, air, and water mobilize as a slurry that flows downslope

(Fig. 1.7). Debris flows are commonly caused by intense surface-water flow, due to


16

heavy precipitation or rapid snowmelt, that erodes and mobilizes loose soil or rock

on steep slopes.

Figure 1.7: Landslide classification based on the type of movement and material (Varnes,

1978 and Cruden and Varnes, 1996).

Problem statement

17

Debris flows also commonly mobilize from other types of landslides that occur on

steep slopes, are nearly saturated, and consist of a large proportion of silt- and

sand-sized material (Highland and Bobrowsky, 2008).

Earth flows occur in moderate to steep slopes where the topsoil or overburden

seasonally becomes saturated by heavy rains (Fig. 1.7). The material slumps away

from the upper part of the slope leaving a scarp, and flows down to form a bulge at

the toe. The mass in an earth flow moves as a plastic or viscous flow with strong

internal deformation. Earth flow triggers include saturation of soil due to

prolonged or intense rainfall or snowmelt, sudden lowering of adjacent water

surfaces causing rapid drawdown of the groundwater table, stream erosion at the

bottom of a slope, excavation and construction activities, excessive loading on a

slope, earthquakes, or human-induced vibration. In the study area, the

characteristics of landslides are rotational slides, translational slides, rockfalls and

toppling, as well as debris and earth flows that occur as a result of heavy rainfall

and earthquakes.

1.4 Problem statement

Ethiopia is currently involved in massive infrastructural development (including

roads, and railways), urban development and extensive natural resources

management. However, during rainy seasons these infrastructure development

works face a huge risk of failure and damage from landslide and other slope failure.

Despite their huge economic, social and environmental significance, so far, mass

movements in Ethiopia have not been given due attention. Rainfall-triggered

landslides continue to cause loss of life and damages to infrastructure, agricultural

lands and the environment. Such hazards are expected to increase in Ethiopia as

more people move into the unstable terrains and as construction expands into

fragile terrains without proper prior site investigation. These problems can also be

further aggravated by climate change. It is, therefore, necessary to evaluate the

factors responsible for landslides in order to minimize the damage caused by

landslides.

Problem statement

18

Landslide-generated hazards in Ethiopia are becoming serious concerns to the

general public and to the planners and decision-makers at various levels of the

government. However, so far, little efforts have been made to reduce losses from

such hazards. The western Afar rift margins are densely populated and contain

several towns, infrastructures such as asphalt roads, large bridges, road tunnels,

and newly proposed railway routes. With the on-going infrastructural

development, urbanization, rural development, and with the present land

management system, it is foreseeable that the frequency and magnitude of

landslides and losses due to such hazards would continue to increase unless

appropriate actions are taken. In this whole socio-economic development,

landslides and related ground failures need to be given due attention in order to

reduce losses from such hazards and create safe geo-environment. Most of the

previous investigators recommended comprehensive studies on the geology,

geomorphology, structural settings, geotechnical characteristics, and hydrological

condition (surface water and groundwater) to clearly define the causes, failure

mechanisms, and mitigation options. It is important to know these factors and

assess their possible influence in inducing instability to the slopes. The

combination of these factors may possibly lead to landslides in a given area., so an

evaluation of these factors and their relation with the past landslides is necessary.

There is a strong need to evaluate the landslide condition in the study area in order

to characterize the landslides and the slopes that are prone to failure based on the

impacts of hydrogeological conditions, numerical modelling of slope stability and

depth to failure plane and possible slip surfaces. Among others, Woldearegay et al.

(2013) emphasized the significance of establishing landslide-groundwater and

landslide-rainfall relationships and numerical modellling for understanding the

initiations of failures of rock slides. Prediction of future landslide occurrence

requires an understanding of the conditions and processes controlling landslides.

Nevertheless, hydrogeological data and its implication on large-scale and deep-

seated landslides are extremely scarce in the study area. As it is mentioned above,

the problem is frequent during the rainy season; which indicates that the landslide

is mainly triggered by rainfall. Hence, lack of appropriate slope stability analysis

Research objectives

19

of rainfall-induced landslide types is believed to have played an adverse role in

aggravating the landslide problem in the study area.

This work aims at understanding the processes leading to the propagation of slope

failure, influencing factors, and failure mechanisms. Therefore, to identify the

most relevant influencing factors a multidisciplinary approach using detailed

geological and topographical, geological structures, groundwater condition

respective rock-water interactions, geophysical survey, and geotechnical

investigations were conducted. Accordingly, the findings of this research can help

to have a much better understanding of the overall landslide controlling

parameters and their mechanisms within the study area. This research work has

used a converging evidences approach from the results of geological and structural

settings, geo-morphometric analysis, hydrogeochemical and isotopic analysis,

seismic refraction survey and geotechnical investigations (kinematic analysis and

numerical modelling). The collected data is thoroughly interpreted using different

scientific techniques. In general, the results of this research work can play a vital

role in adding significant knowledge and disclose reliable insights into the

landslide monitoring and forecasting of the study area and surroundings. It can

also help to advance our understanding and propose possible mitigation measures

to minimize the effects of landslides on natural resources and natural

environment. Furthermore, this study could be helpful to answer similar landslide

and landslide-related hazard problems in rift margins and highland terrains

similar to Ethiopia.

1.5 Research objectives

1.5.1 Main objective

The main objective of this study is to understand the controlling parameters of

deep-seated landslide, the processes leading to the triggering of a landslide, and

the failure mechanisms in the Debre Sina area and its surroundings. This will be

valuable to the socio-economic planning and management of the environment

around the margins of the western Afar depression as well as other landslides

along the rift margins and associated highlands of Ethiopia.

Summary of methodology

20

1.5.2 Specific objectives

The specific objectives of this research are to (a) evaluate the major causes and

failure mechanisms of large-scale landslides, (b) understand the processes leading

to initiation and propagation of slope failure, (c) determine depth to bedrock and

failure plane and describe the existing situation of faults under the landslide slope,

(d) study the effects of groundwater and rainfall in causing slope failures (e)

evaluation of the rock-water interactions and (f) identify potential landslide areas

and perform a slope stability analysis through numerical modelling.

1.6 Summary of methodology

In order to achieve these objectives, a comprehensive detailed geological, morpho-

tectonic analysis, geomorphological, hydrogeological, hydrogeochemical, isotopic,

geophysical (seismic refraction and vertical electrical sounding) and geotechnical

(kinematic analysis and numerical modelling) investigations were conducted. In

addition to this, pertinent secondary data compiling and reviewing were carried

out.

Geological and geomorphological mapping and a number of field discontinuity

measurements (Appendix B) and observations were conducted throughout the

fieldwork. More than 40 years (1974–2016) of rainfall records of the meteorological

station at Debre Sina station were collected from the National Meteorological

Agency of Ethiopia (NMA). The geomorphic property of the area was generated

from a digital elevation model (DEM) using a Geographic Information System

(GIS) to outline the different geomorphic characteristics and landforms in the area.

The GIS and remote sensing (RS) analysis was also used to produce new geologic,

structural and geomorphologic maps of the area (Appendix A). Furthermore,

seismic refraction investigations were carried out to assess the mechanical and

geological conditions which determine the nature of sliding movements. The

seismic refraction survey was conducted along three nearly orthogonal survey lines

(in total having a horizontal length of 1 km) within the recently affected area.


21

Groundwater chemistry and stable isotopes analyses were used to characterize the

groundwater flow system and rock-water interactions. For this purpose, 65 water

samples were collected from the study area for hydrochemistry analysis and 39

water samples for stable isotopes analysis. Measurements of the field

hydrochemical parameters such as temperature, pH value and electrical

conductivity (EC) of sampled water were made in-situ, and each electrode was

calibrated. Titration analysis for the sensitive anions (HCO3) was also performed

in the field by employing a burette titration method. All the major ions except total

Fe (Fetot) were analyzed by using an ICS-1000 Ion Chromatography. The Fetot was

analyzed by using atomic absorption spectrometry (AAS). Vertical electrical

soundings (VES) were also carried out in order to trace the orientation and location

of the faults and geological contacts, which can have considerable effect on the

groundwater circulation, as well as to map the various aquifer systems.

Extensive laboratory tests were conducted, such as for determination of uniaxial

and triaxial strength, ultrasonic pulse velocity (vP), density, porosity and

permeability of the volcanic rocks to evaluate their geomechanical behaviour in

general and weathered materials in particular and their effect on instability

processes. Cylindrical core samples with a diameter of 30/40 mm and length 60/80

mm were drilled out of larger rock blocks. Three different densities measurements

for 21 rock samples: the bulk density of the dry sample material, the buoyant

density of water-saturated material (according to Archimedes' principle) and the

average grain density were also measured. In order to measure the degree of

saturation and water absorption, 21 cylindrical specimens with diameter 30/40 mm

and length 60/80 mm from different volcanic rocks were submerged in distilled

water under a constant air vacuum pressure. Finally, pore-volume, total pore area

and bulk density, as well as a value for effective porosity, were obtained. Total

porosity was obtained indirectly by pycnometer tests. The permeabilities of 19 rock

cores were also determined using a Darcy flow apparatus.

Uniaxial compressive strength (UCS) tests were used for the determination of

static Young's modulus, Poisson's ratio and the uniaxial compressive strength.

Young's modulus (E) and Poisson’s ratio (ν) derived from this test are the keys to


22

define a stress and strain relationship. 27 cylindrical samples were prepared from

aphanitic basalt, porphyritic basalt, rhyolite, trachyte, ignimbrite and

trachyignimbrite and welded tuff rocks (30/40 mm in diameter and 60/80 mm in

length) to perform a compression test. Young's modulus (E) was calculated by

fitting the linear portion of the axial stress-strain curve and Poisson’s ratio (ν) was

calculated by dividing the slope of radial strain curve by slope of axial strain curve.

Triaxial compressive strength (TCS) tests were used to measure the strength of

cylindrical rock specimens as a function of confining pressure. Multistage triaxial

tests and individual triaxial test were performed on samples where imminent

failure point or post-peak reduction in strength could be not easily recognized. 34

cylindrical samples (30/40 mm in diameter and 60/80 mm in length) were used to

perform individual and multi-stage triaxial tests. Each sample was tested under

dry state and it was subjected to different confining pressures. The strength of

rocks is represented by Mohr-Coulomb and the Hoek-Brown failure criterion.

Mohr-Coulomb criteria utilize the concept of cohesion (c) and friction angle () to

estimate the major principal stress at failure for a given minor principal stress.

The tensile strength of the rocks was also determined by Brazilian disc tests which

are indirect tension tests. The tests are indirect because no tensional external

loading is applied to the specimen. For this test, 27 circular disk samples with a

diameter of 30/40 mm and a thickness of 15/20 mm were used.

In this study, limit equilibrium (LE) analysis was carried out using SLIDE2

(Rocscience Inc. 2018) to compute the factor of safety. Among the available limit

equilibrium methods (LEM), Bishop simplified, Janbu simplified, Spencer, and

Morgenstern-Price were used for conducting the comparative study between limit

equilibrium and finite element methods. Finite element method (FEM) analysis

was performed together with the Point Estimate Method (PEM) using the RS2

software (Rocscience Inc. 2020). This software utilizes the shear strength reduction

(SSR) technique in computing the factor of safety as critical reduction factor. It

makes use of the SSR technique in which the shear strength parameters are

reduced in small increments until failure occurs in the slope. In these simulations,

elasto-plastic analysis is used to compute deformations and stresses. The Mohr-

Structure of the Thesis

23

Coulomb elasto-plastic material model was used for shear zones in order to allow

for plastic deformation and failure. The faults and interfaces also follow Mohr-

Coulomb failure criterion in order to evaluate the possibility of slipping failure

along the faults.

1.7 Structure of the Thesis

This thesis comprises six chapters and the chapters of the thesis are organized as

follows. The present thesis has a cumulative structure and consists of four studies

(chapter 2, 3, 4 and 5) as well as a summary and future research perspectives

(chapter 6) which focus on a systematic understanding of landslide influencing

factors and failure mechanisms in the margins of the western Afar depression in

Debre Sina, Ethiopia.

Chapter 1: describes the overall background of this research work including

motivation, regional geological and tectonic setting, landslide types and their

failure mechanisms, problem statement, research objectives and summary of

methodology. The introduction states the study aims and research problem

statements as well as the background to the study. The method describes how the

survey and case study was performed.

Chapter 2: addresses the typology, distribution of landslides, detailed evaluation

of representative landslides and understand the predisposing factors that control

the development of the landslide and failure mechanisms along the rift margins

and highland terrain linked to deep-seated potential landslides. It is mainly

evaluated based on the context of geology (lithology and structure), morphometric

analysis (elevation, slope angle and aspect), discontinuities analysis, rainfall, and

earthquakes. This chapter was published in the Bulletin of Engineering Geology

and the Environment Journal (Mebrahtu et al., 2020a).

Chapter 3: deals with the seismic refraction investigations along the deep-seated

landslide main scarp to identify the depth to failure plane, possible slip surfaces

and the mechanism of slope failures, the existing situation of faults and their

continuity under the landslide, and to determine the internal composition of the

sliding masses. This chapter also shows analysis of structurally controlled failures

Structure of the Thesis

24

using kinematic analysis possible failure mechanisms using Dips 7.0 program.

This chapter was published in the Geomorphology Journal (Mebrahtu et al.,

2020b).

Chapter 4: discusses the effects of the hydrogeological and hydrogeochemical

dynamics on landslide triggering by using converging evidences from geological,

geomorphological, geophysical, hydrogeological, hydrogeochemical and isotopic

investigations. It also shows the conceptual groundwater flow model of the Debre

Sina landslide area from the converging evidence of the different datasets. This

chapter was published in the Hydrogeology Journal (Mebrahtu et al., 2021).

Chapter 5: focus on the numerical modelling of slope stability using the limit

equilibrium methods and the finite element method. Four selected slope sections

were analysed for stability conditions of the slopes under static and pseudo-static

loading using the horizontal seismic coefficient to model their stability during a

seismic event. LEM and FEM using SSR for the analysis of slope stability problems

are presented in this chapter. Finally, the respective FS obtained from the LE and

FE analyses was compared. This chapter is submitted for publication.

Chapter 6: summarizes the main conclusions and recommendations on future

researches. Chapters 2, 3, and 4 have been published in international, peer-

reviewed scientific journals and chapter 5 contains a submitted manuscript.

Finally, the main findings on the failure mechanisms and slope stability

evaluations from the study area summarized in chapter 6, including

recommendations for further study.

Predisposing and triggering factors of large-scale landslides

25

Chapter 2

2 Predisposing and triggering factors of large-

scale landslides in Debre Sina area, central

Ethiopian highlands

This chapter is based on Tesfay Kiros Mebrahtu, Bedru Hussein, Andre Banning,

Stefan Wohnlich (2020a). Predisposing and triggering factors of large-scale

landslides in Debre Sina area, central Ethiopian highlands. Bull Eng Geol Environ.

80:1–19. DOI: 10.1007/s10064-020-01961-1.

Abstract

A large number of landslide events have repeatedly struck the border zone of the

northwestern plateaus of Ethiopia. Debre Sina area is one of the most tectonically

active areas located along the western margin of the Afar depression, which is

frequently affected by landslides. Despite that, urban and rural development is

currently active in almost the entire area. It is crucial, therefore, to understand

the main causes and failure mechanisms of landslides in the Debre Sina area and

its surroundings. The present study investigated landslides using field mapping of

geological and geomorphological features, remote sensing, geo-morphometric

analysis, structural analysis, rainfall data, landslide inventory, and earthquake

data. The results of the study indicate that large-scale and deep-seated landslide

problems appear to be caused by complex geological settings and rugged

topography. In particular, the location and morphology of the Yizaba Wein and

Shotel Amba landslides are strongly controlled by geological structures. Their

flanks are bounded by high angle faults, and their main basal failure surfaces have

developed within a W–E striking eastward-dipping normal fault zone. The complex

litho-structural and morphologic settings play a vital role in controlling the

geometry of the slip surfaces and the stability of the landslides.


26

2.1 Introduction

Landslides and related ground movements are among the common geo-

environmental hazards in many hilly and mountainous terrains of the world.

There have been several landslide events in Ethiopia that have resulted in

considerable socio-economic impact. The highlands of Ethiopia are highly

susceptible to slope instability due to heavy rainfall and land-use change, including

the effects of road construction (Woldearegay, 2013). Urban and rural development

is currently taking place in almost all areas. Landslides represent one of the main

constraints for the development of road infrastructures in many parts of Ethiopia.

As Woldearegay (2013) mentioned, several authors indicated that modification of

slope geometry through natural or man-made processes could influence the

stability of slopes (Varnes, 1978; Greenway, 1987; Bell, 1999). Nowadays, there is

a significant increase in landslides in Ethiopia as the road network has continued

to expand over recent decades. This problem is also significant in the Debre Sina

area and its surroundings. Active extensional tectonics, high heat-flow, and

intense volcanism associated with the East African Rift System are the main

factors for frequent hazardous geological phenomena in Ethiopia (e.g., Chorowitz,

2005; Abebe et al., 2007; Agostini et al., 2011; Kycl et al., 2017). According to

Ayalew (1999), major faults that run parallel to the Main Ethiopian Rift (MER)

have formed release surfaces for structurally controlled deep-seated landslides.

Most of the large-scale landslides in Ethiopia have occurred along the MER scarps

and also developed in plateau regions (e.g., Abramson et al., 1996; Ayalew, 1999;

Temesgen et al., 2001; Ayalew and Yamagishi, 2004; Ayenew and Barbier, 2005;

Nyssen et al., 2006; Fubelli et al., 2008, 2013; Moeyersons et al., 2008; Coltorti et

al., 2009; Van Den Eeckaut et al., 2009; Abebe et al., 2010; Zvelebil et al., 2010;

Vařilová et al., 2015). The margins of the western Afar depression are currently

under threat with the problem of landslides in many places. Particularly, the

Debre Sina area is known for its several landslides that have occurred in the past

(EIGS, 1979; Schneider et al., 2008; Woldearegay, 2008; Abay and Barbieri, 2012;

Alemayehu et al., 2012; Kropáček et al., 2015; Meten et al., 2015). There is also

evidence of active landslides in the area. As stated by different researchers


27

(Schneider et al., 2008; Abebe et al., 2010), the landslides in the rift margins of

Ethiopian highlands are associated with deep-seated and structurally controlled

deformations which require a detailed understanding of the geological and

structural settings in order to clearly define the failure mechanisms. Despite that,

less work has been conducted on landslides in the study area and its surroundings

as the area is hardly accessible.

A large-scale landslide incident occurred in the study area at Yizaba Wein and

Shotel Amba localities on September 13, 2005, and the slope instability problem

still remains very active. Adequate characterization of landslides requires a deep

understanding of the causes and failure mechanisms. This, in turn, requires a

detailed study of the geology, topography, and physical properties of rocks and soils

that occur in the areas forming unstable slope profiles. This study focuses on the

2005 large-scale landslide and recently reactivated landslides, which were induced

by a combination of specific local conditions and external factors. This work aims

at understanding the processes leading to the propagation of slope failure, the

influencing factors, and the failure mechanisms based on data from field

observations focusing on the geological and topographical conditions of the area,

morphometric analysis, tectonic activity, hydrometeorology, and detailed

evaluation of representative landslides with identical features and frequent

reactivation that are imposing potential risks on the local residents and nationally

important infrastructures passing through the area.

2.2 The study area

Debre Sina is situated in the border zone of the central-western highlands of

Ethiopia about 190 km north of Addis Ababa, where the MER widens into the Afar

depression (Fig. 2.1). The study area is bounded to the north and south by the UTM

1077165 m N and 1108635 m N; and to the west and east by the UTM 571065 m E

and 601125 m E. It covers an area of 946 km2, with an elevation ranging from 1154

to 3691 m above sea level (a.s.l). The area comprises an extremely steep

escarpment and a narrow strip of the plateau (Fig. 2.1). The escarpment comprises

steep mountain chains and rugged valleys which drain into the central-eastern and


28

western lowlands of Ethiopia. The study area is flanked by the lower and upper

normal faults and the Tarmaber–Mezezo mountain chain.

Figure 2.1: Location map of the study area.


29

The climate of the area is sub-humid to humid with an average annual

precipitation of about 1812 mm distributed with strong seasonality, having

maxima in summer and spring. Debre Sina is one of the wettest parts of the

country with a bimodal rainfall, with peaks in the months of June to August. The

average maximum temperature is 25 °C and, the average minimum is 10 °C.

The western margin of the study area was produced during the Tertiary-

Quaternary extensional phase, which was accompanied by down-warping of the

Afar depression and rift-ward tilting of faulted blocks (Zanettin and Justin-

Visentin, 1974; Almond, 1986; Mohr, 1986). The outcropping bedrock consists of a

sequence of Tertiary Trap Volcanic Series: Ashangi basalt formation and Tarmaber

basalt formation (Mohr, 1971; Kazmin, 1973; Zanettin et al., 1974; Kazmin, 1979).

The possible seismic source zones are very near to the study area (Ayele, 2009).

2.3 Materials and methods

This study was started by compiling and reviewing pertinent secondary data.

Topographic maps, regional geological maps, and satellite images were collected

from Ethiopian Mapping Agency, Geological Survey of Ethiopia, and US Geological

Survey, respectively. Detailed fieldwork investigations were carried out from April

to June 2016, October to November 2017, and June 2018. The fieldwork was

accomplished along selected traverse lines to cover the study area, focusing on the

topographical and geological settings. Following the traverse lines, many

observation points were selected in addition to the delineation and mapping of the

past landslides by Google Earth imagery. The landslide inventory was obtained

based on a 1:50,000 scale using aerial photographic interpretations for delineation

and identification of past landslides and geological and geomorphological field

mapping in the Debre Sina area and its surroundings. The distribution of

landslides was examined in the context of geology (lithology and structure),

topographic slopes, rainfall-landslides relationships, and earthquakes. There were

new landslide occurrences during the fieldwork. A number of surface discontinuity

measurements were conducted, and their kinematic relations with respect to the

general slope trend were determined using SSWIN 2.2 stereographic projection

facilities in a lower hemisphere equal area net. These have allowed visualization


30

of the main structural trends and interpretation of their effects on the general

conditions of stability of the lithological units.

Geological structures of the area were extracted from Landsat images (Thematic

Mapper (TM) and Enhanced Thematic Mapper Plus (ETM+) and fieldwork to

visualize their relationships with large-scale and deep-seated landslides. The

images have been interpreted using Erdas Imagine 2014 by applying different

enhancing and band composition techniques and digitized using ArcGIS 10.5.

Structural measurements of faults, fractures/joints, and lineaments were

conducted to reveal the structural predisposition of a slope to sliding. The strike

and dip of the fault planes and fractures/joints detected in each structural station

are reported in stereographic diagrams. Moreover, the geomorphic property of the

area was generated from DEM data using GIS utilities to outline the different

geomorphic characteristics and landforms in the area. Accordingly, the GIS and

remote sensing (RS) analysis were used to produce new geologic, structural, and

geomorphologic maps of the area. Processing of the DEM was carried out using a

high-resolution satellite image in a Geographic Information System (using

ArcView 3.3/ArcGIS utilities 10.5 and ILWIS 3.3) to extract elevation, slope

gradient, and slope aspect, which are key to identify significant features of

landslide events. These parameters are useful for the assessment of the

geomorphology of the area. Besides this, Erdas Imagine 2014, Global Mapper 18,

and CorelDRAW X7 were used for database creation, image processing and

interpretation, spatial data analysis, and high-quality output maps and

illustrations.

More than 40 years (1974–2016) of rainfall records from the Debre Sina station

were collected from the National Meteorological Agency of Ethiopia (NMA). These

data were analyzed to assess the distribution of precipitation variations and the

long-term development of rainfall in the study area. Furthermore, daily data from

the last 43 years were used to calculate the total annual precipitation in the years

and subsequently compared with the landslide occurrences/ reactivations.


31

2.4 Results

2.4.1 Geology and geomorphology of the study area

The outcropping lithology in the study area is represented by five Tertiary volcanic

units associated with volcanic ash and two Quaternary superficial deposits

(colluvial and alluvial deposits) (Fig. 2.2). Aphanitic basalt-porphyritic-

agglomerate units crop out in the gully areas and series of cliffs and benches in

deeply dissected valleys (Fig. 2.2). The basalt is grayish black to black or reddish-

brown colored and fine to medium-grained. It has a vitreous appearance, vesicular

to amygdaloidal with clearly developed columnar joints and intense fracturing, and

exhibits spheroidal weathering. The porphyritic variety is black to grayish-black,

medium to coarse-grained, massive, with plagioclase and olivine phenocrysts. The

agglomerate contains angular to sub-rounded glassy basaltic rock fragments as

groundmass ranging from few millimeters to few centimeters. There is a paleosoil

ranging in thickness from 10 to 20 cm between the porphyritic basalt and aphanitic

vesicular basalt. Springs and seepage emerge at the contact between the basalt

and the overlying colluvial sediments.

The ignimbrite-tuff-volcanic ash unit is widespread in the lower parts of the

escarpment. The association mainly consists of pumiceous lapilli tuff and volcanic

ash with subordinate ignimbrite, trachyte, and rhyolite. The ignimbrite beds form

gentle to steep slopes, elongated ridges, and isolated valleys (Fig. 2.2). The

ignimbrite is fine to medium-grained, highly consolidated, laminated, and

moderately to highly fractured and weathered. The tuff is medium-grained,

massive, weakly to moderately welded, often weathered, and friable, and exhibits

horizontal layering. The volcanic ash is fine-grained, locally exhibiting sub-

horizontal stratifications. The ignimbrite-tuff-volcanic ash beds form small cliffs,

are highly altered and intensely weathered, and are vertically jointed and highly

shattered by faulting. There are also dikes about 60 cm to 1.50 m thick which are

fine-grained and glassy in texture with N35°E, N60°E, and N20°W strike directions.

The porphyritic basalt-scoriaceous agglomerate unit consists of dominantly

porphyritic basalt and scoriaceous agglomerate with subordinate aphanitic basalt


32

and vesicular basalt. The rock exhibits notable textural and compositional

variations vertically: aphanitic basalt, porphyritic basalt, agglomeratic basalt, and

amygdaloidal basalt. The porphyritic basalt is medium to coarse-grained, with

plagioclase and olivine phenocrysts, massive, blocky in appearance, and exposed

in gentle slopes to steep cliffs. The vesicular basalt is a highly porous and friable

rock type, which is dominant in the western, northern, and southeastern parts of

the area. There is a paleosoil developed as a thin layer between the porphyritic

basalt and vesicular basalt. The porphyritic-agglomerate basalt shows a high rate

of spheroidal weathering and breaks easily to very small size material. The

weathering and fracturing increase towards the major joints and layering. It is

mostly found in the upper part, overlying the volcanic ash unit and underlying the

ignimbrite/rhyolite unit. The porphyritic and scoriaceous agglomerate basalt crops

out around Armanyia, and they are observed to be susceptible to local weathering

and erosion. The units found around Yizaba Wein and Shotel Amba localities are

dissected by faults oriented NNW–SSE, NNE–SSW, and E–W (Figs. 2.2 and 2.12).

The Tarmaber basalt, which is the most abundant in the study area, is mainly

exposed in the western part of the area in the high-rising Tarmaber–Mezezo

mountain chains. This rock formation forms vertical cliffs and ridges trending in

the N–S direction and some E–W offsets (Fig. 2.2). It is highly weathered, jointed,

and fractured and shows well-developed columnar joints. There are three sets of

joints trending in the N–S, E–W and NE–SW directions, with vertical to sub-

vertical dip angles. The upper ignimbrite unit crops out in the western part of the

area overlying the Tarmaber basalt. The upper ignimbrite is fine-grained, light

gray to light yellowish color, massive to bedded, highly weathered, and crossed by

sub-vertical to vertical fractures.

The colluvial deposits are common along the upper pediments adjacent to cliffs,

presumably transported downslope by the action of gravity and slope wash (Fig.

2.2). These deposits consist of unsorted to poorly sorted loose soil sediments and

rock fragments, with large blocks of basalt toppled from upslope cliff faces. Some

seepages or springs that drain from the highlands disappear in this thick colluvium

material and re-emerge at the lower slope breaks or stream banks. Most of the


33

seepages or springs also emerge along the contacts of the colluvium with the

ignimbrite-volcanic ash unit.

Figure 2.2: Geological map of the study area.


34

The alluvial deposits are restricted along the major riverbeds, riverbanks, and

their tributaries representing riverbeds and terraces (Fig. 2.2). The sediments

found at the riverbed and riverbanks are dominated by coarser materials (such as

sands, gravels, and boulders). This indicates that the streams are highly erosive

and transport debris from the landslide zones. The deeply dissected gullies are

filled by sediments up to several tens of meters thick, which consist mainly of

colluvial and alluvial deposits.

The geomorphology of the study area is very complex and strongly influenced by

tectonics, rock weathering, and erosion. Due to the Quaternary tectonic uplift

(Fubelli and Dramis, 2015) and the related deepening of the rift, the hydrographic

network increased the relief energy, giving rise to steep slopes and deeply cut

valleys in the bedrock and thus predisposing the slopes to the development of mass

movements. Some morphological aspects of the area have been visualized through

the morphometric analysis of a 30-m resolution DEM using utilities of ArcGIS. The

study area elevation ranges from 1153 m a.s.l. in the northeast and southeast

lowlands to more than 3690 m a.s.l. in the western catchment boundary at the

margin of the western rift escarpment (Fig. 2.3). The mean elevation is 2209 m

a.s.l. with a standard deviation of 620 m. There is a high topographic relief of 2.54

km that is marked by deeply dissected gullies and channels, rugged topography,

and high-altitude continuous ridge chains with steep escarpments. The land

surface on the central part is characterized by rugged morphology and in the west

is bordered by a high cliff. A gently to steeply sloping ground follows below the

main steep escarpment and extends downslope towards the Dem Aytemashy river

valley. The gullies located in the area form triangular-faceted landforms due to

weathering and erosion processes.

These peculiar characteristics of the study area presumably denote complex

landscape evolution in the geologic past. At very high elevation, there are

mountain summits that usually consist of slightly to moderately weathered rocks

whose shear strength is very high. The Debre Sina area is basically a warped

plateau with slope gradient in the range of 0° to 63° with a mean value of 11.23°

and standard deviation of 6.7° (Fig. 2.4). The slope range between 0° and 10°


35

accounts for 26.54% of the area which represents the lowlands, flatlands, linear

valleys, and plains in between relatively elevated hills. This constitutes mainly the

lowlands in the northeast, west, and southwest; lower valleys and benches of major

streams; and wide valley plains that stretch to the northeast and southeast.

Figure 2.3: Elevation map of the study area.


36

The slope class 10°–20° accounts for 49.58% of the area and marks small hills and

gentle slopes at the lower parts of escarpment. The slope class ranging between 20°

and 30° covers 20.47% of the area and denotes small scarps in deeply incised stream

gorges. The slope class 30°–40° accounts for 2.43% of the area and belongs to

prominent erosion and tectonic scarps, as well as elevated hills and ridges above

gently sloping grounds.

Figure 2.4: Slope angle map of the study area.


37

The slope class 40°–50° accounts for 0.64% of the area and marks the upper slopes,

local ridges, prominently mesas, and hills. The highest slope gradient above 50°,

which covers only 0.34% of the study area, is a characteristic of high altitude

continuous ridge chains with cliffs along the elevated mountains and volcanic

centers and to some extent the deeply dissected channels.

Figure 2.5: Slope aspect map of the study area.


38

Slope aspect defines direct contact with sunlight and winds, affecting indirectly

other factors that contribute to landslides, such as precipitation, soil moisture,

vegetation covers and soil thickness (Clerici et al., 2006). The mean slope aspect is

165° with a standard deviation of 105°. The computed slope aspect map was

classified into nine directional quadrants (Fig. 2.5) as flat (−1), north (0–35° to 324–

360°), northeast (35–71°), east (71–107°), southeast (107–143°), south (143–216°),

southwest (216–252°), west (252–288°), and northwest (288–324°). The east and

southeast facing slope aspects are very widely distributed with frequency values of

22.5% and 19.2%, respectively. The slopes facing north, northeast, and south have

the next higher frequency values, whereas slopes facing northwest, west, and

southwest have low-frequency values.

2.4.2 Description, typology, and distribution of landslides

The distribution of landslides in the study area was found to be notably

concentrated in the south and north vicinities of Debre Sina town (Fig. 2.6). Before

the occurrence of the more recent landslides, there were observations of symptoms

(such as tensile fissures in the soil) in the years 1977 and 1978. This was followed

by a major development of landslide in late 1979 (EIGS, 1979).

Many landslides have remained active for a long period, and a considerable

number of relict landslides were reactivated in 2004, 2005, 2006, 2007, 2014, and

2016 during intensive rainfall and tectonic activity (Woldearegay, 2008; Abay and

Barbieri, 2012; Kropáček et al., 2015). The oral information of the local people

indicates that the landslides in the study area have been active for at least the last

15 years. Following the Varnes (1978) classification, the most common types of

landslides in the study area include (a) rotational slides, (b) translational slides,

(c) rockfalls and toppling, and (d) debris and earth flow. The most spectacular ones

were observed at Yizaba Wein, Shotel Amba, Nib Amba, Nech Amba, Wanza Beret,

and Shola Meda areas (Fig. 2.6). Their main characteristics and the affected areas

are briefly outlined below.


39

Figure 2.6: Landslide inventory and morphostructural map of the study area.

2.4.2.1 Yizaba Wein and Shotel Amba landslides

The largest landslide in the Debre Sina area, ca. 40 km2 according to Woldearegay

(2008), affected the localities of Yizaba Wein and Shotel Amba in September 2005.

Here, the slope shows clear morphological evidence of a deep-seated landslide (e.g.,

breaks on the slope, depressions, terraced surfaces, and convex-concave forms)


40

(Fig. 2.7b, c). The landslide, which involves a large portion of the slope from the

top to the valley bottom, consists of two distinct and compound blocks: (i) an older

block and (ii) a more active one that underlies the former one (Fig. 2.7a). The active

block can be classified as a roto-translational rock slide (Borgatti et al., 2006) and

is characterized by a semi-circular head scarp (crown at 2457 m a.s.l.). Its mid to

lower part is concealed by a secondary active debris slide and debris/earth flow

(Fig. 2.7c). During the fieldwork, relatively fresh cracks 20–50 cm wide were also

identified in the lower part of the slope close to the debris flow (Fig. 2.7c).

Figure 2.7: (a) Translational slide occurred in 2005, (b) roto-translational rock slide

and rockfalls, (c) Yizaba Wein and Shotel Amba convex-concave landslides and (d)

rotational slide and earth flow dipping downslope towards Dem Aytemashy river.


41

The Yizaba Wein and Shotel Amba landslide is truncated by the most recent and

active body that involves a significant part of the slope and is characterized by

fresh morphological features (Fig. 2.7a, b). Furthermore, the rock mass moved

along the frontal part of the landslide causes the detachment of rock fragments

and blocks (rockfall, rock slide, and block toppling), which bounce towards the river

(Fig. 2.7b). Additionally, the constant deepening of the streams causes local

undercutting of the cliff/sliding block and triggers new shallow landslides from the

left bank of the watercourse, with a general retrogressive trend (Fig. 2.7d).

The landslide was initiated in the heavily fractured porphyritic basalt and highly

shattered ignimbrite and volcanic ash units. The bedrock is mostly covered by

colluvium deriving from the uppermost steep slope and cliff edges. The landslides

were mostly associated with colluvial materials, including boulders and a higher

proportion of granular soil. The hydrogeological conditions of the terrains are

favorable for the development of seepage within the pyroclastic sediments (tuff and

pumice horizons) and unconsolidated deposits during periods of rainfalls.

2.4.2.2 Nib Amba landslide

The Nib Amba area is widely affected by active and dormant landslides of different

types and sizes including translational and rotational slides, debris/earth slides

and flows, and rockfalls. Superimposed landslide bodies confirm that the spatial

distribution of the recent landslides is frequently influenced by the presence of

older landslides. The 2005–2007 events were characterized by multiple

retrogressive translational slides in the upper part and multiple advancing

rotational slides in the lower part, especially along the stream banks (Fig. 2.8b, c).

In May 2016, rainstorms triggered an earth flow with a volume of several cubic

kilometers and other landslides such as debris flow, shallow slides, and rockfalls

which caused damage to dwellings, agricultural lands, and the natural

environment (Fig. 2.8a, d).

Most failures have developed due to the soft interlayered pyroclastic sediments

(tuff and pumice) overcharged by the abundant rainfall. In 2005–2007, reactivated

landslides destroyed farmland and dwellings in the gorge of the Koda Menkeriya


42

river (Fig. 2.8b). Outcropping lithology and bedrock structure together with the

rugged topography were responsible for slope instability in this locality. The

landslides were mostly associated with soft interlayered pyroclastic sediments (tuff

and pumice) and black cotton soil. A quasi-rotational slide created a stagnant pond

at the frontal part of a slope movement (Fig. 2.8c).

Figure 2.8: (a) Debris flow demolished agricultural land in Nib Amba, (b) rock slide

in Nib Amba, (c) deep-seated rotational slides form a pond at the lower part of the

slide zone and (d) earth flow demolished farmland.

2.4.2.3 Nech Amba landslide

Shallow to deep landslides in the Nech Amba area involve weaker material (Fig.

2.6), mostly composed of transported loose deposits with a higher proportion of clay

underlain by pyroclastic sediment and agglomerate basalt (Fig. 2.9a, b). In some

areas, the failures are concentrated along streams and riverbanks within deep


43

gullies. Debris flows, rock slides and retrogressive rotational slides are common in

the area. These landslides are likely influenced by the presence of weathered rocks

and susceptible volcanic ash. The landslide was reactivated after heavy rainfall on

May 6, 2016, and a new episode of high-potential landslides occurred along the

crowns and river banks. Farmland and dwellings were devastated by reactivated

landslides in 2016. Moreover, there are irregular tension cracks in buildings and

tilted houses found near the river bank.

Figure 2.9: (a, b) Earth slides around Nech Amba area occurred on May 6, 2016,

(c) pre-existing landslide scars and active landslides in the gorge of the Majete

river and (d) a quasi-rotational slide widening retrogressively.

2.4.2.4 Wanza Beret landslide

The most remarkable and probably the second largest slope failure, a huge earth

slide, was recorded in Wanza Beret, highly rugged, deeply incised by various

gullies and rivers, and subjected to severe erosion. The main landslide body is

currently about 200 m wide and 70 m long (Fig. 2.9d). Various other types of


44

landslides (quasi-rotational slides, debris slides, earth slides, debris flows, and

earth flows), were identified in this area. A quasi-rotational slide is one which the

displacement of the compound block would have opened up a depression that was

rapidly infilled with debris which continues for some time after (Palmer et al.,

2007). This occurs in limited cases where the thickness of unconsolidated deposits

is thick enough to generate deep failure surfaces (Woldearegay et al., 2006). Slope

failures in highly weathered basalt, pyroclastic sediments, and unconsolidated

material are induced by the undercutting of the Majete river and its tributaries

(Fig. 2.9c). The thickness of unconsolidated deposits has a significant role to

generate deep sliding surfaces. Farmland and dwellings were destroyed by the

reactivated dormant landslides in 2005–2007 (Fig. 2.9c). These landslides were

mostly associated with colluvial inclusions of boulders and a higher proportion of

granular soil. A quasi-rotational slide has created a stagnant pond at the lower

part of the slope (Fig. 2.9d).

2.4.2.5 Shola Meda landslide

The Shola Meda landslide is considered the evolution of ground cracking and a

slow creeping zone (Fig. 2.10c). There are scars occurring intermittently with

smaller magnitude. In Shola Meda, an asphalt road crossing Tarmaber–Debre

Sina–Armaniya–Shewa Robit cracked at several places and has also been damaged

(Fig. 2.10d). Furthermore, a gravel road under construction heading from Debre

Sina to Shotel Amba was sliding at several places on May 6, 2016, following the

heavy rainfall (Fig. 2.10a, b). A considerable number of relict landslides in the

Debre Sina area such as Sina, Yizaba Wein, and Shotel Amba localities were

reactivated during the construction of a gravel road along the Debre Sina–Shotel

Amba line. In addition to the active tectonics and seismic forces, the slope

modification during the construction of the gravel road caused a man-induced slope

failure. There are several closely spaced, shallow landslides and ground cracks in

black cotton soil, a dark gray in color, fine to very fine-grained clay, slightly moist

to moist and rough in texture. During the fieldwork, the slope showed widespread

instability conditions with ground tension cracks, rock slides, and earth slides (Fig.

2.10a–c).


45

Figure 2.10: (a) Photographs showing rock slides around Armaniya along the

asphalt roadside, (b) earth slides, (c) tension cracks in a black cotton soil at Shola

Meda, (d) asphalt road collapsed along Debre Sina and Armaniya.

2.5 Discussion

The failure mechanisms of the large-scale landslides of the Debre Sina area were

evaluated based on the context of geology (lithology and structure), kinematic

analysis of discontinuities, rainfall, and earthquakes.

2.5.1 Lithology and structure

Several landslide occurrences are observed in the ignimbrite-tuff-volcanic ash and

colluvial deposit litho-stratigraphic units. The pyroclastics contain lapilli tuff, tuff

breccia, and tuffaceous rocks, which are susceptible to slaking, giving rise to

landslides. The highly weathered basaltic rocks and unconsolidated materials in

steep slope areas are mostly associated with the high landslide and rockfall

susceptibility zone. The porphyritic basalt-scoriaceous agglomerate cliffs that rise

above the ignimbrite-tuff-volcanic ash and colluvial deposit are particularly

subject to considerable rockfalls and topplings as well as rock slides (Fig. 2.7b).


46

The area covered with the Tarmaber basalt is least affected by landslide apart from

some rockfalls at the base of the cliffs. Field observations confirmed that many

types of landslides are densely distributed in the colluvial deposits, ignimbrite-

tuff-volcanic ash, porphyritic basalt, and scoriaceous agglomerate units. This can

be attributed to the high degree of weathering and fracturing, which in turn

reduces the strength of rocks. The intense fracturing and presence of faults favor

an easy movement along existing fault planes during saturation of the rocks or

soils and during seismic events or a combination of both conditions.

Figure 2.11: Stereographic projection of joints/fractures orientation data: (a) rose

diagram showing strike direction, (b) rose diagram showing dip direction, (c) plots

of poles and (d) pole density contour diagram.


47

The geological structures are coinciding with the head scarp of the old and new

failure planes, and several slide incidences also are observed in the ignimbrite-tuff-

volcanic ash and porphyritic basalt-scoriaceous agglomerate units. The structural

setting of the study area is associated with the extensional fault system of the rift

margin that bound the northwestern Ethiopian plateau to the west. The major and

longest normal faults are mainly composed of many overstepping small fault

segments propagating laterally, increasing in size and amount of displacement.

The main geological structures identified are faults, lineaments, and

fractures/joints (Fig. 2.12). A plot of contoured pole concentrations and a rose

diagram of 360 discontinuity measurements from the landslide areas are shown in

Fig. 2.11d.

There are four dominant sets of discontinuities (N–S, E–W, NE–SW, and NW–SE)

(Fig. 2.11a), which are mainly vertical to sub-vertical (tectonic joints). The E–W

and N–S striking lineaments record the highest frequency and cut across by the

NW–SE trending lineaments (Fig. 2.12). The N–S joint set dominantly dips

towards E, the E–W set dips either N or S, the NE–SW set dominantly dips towards

SE, and the NW–SE set dominantly dips towards NE. The dip angle ranges from

10° to 90° (Fig. 2.11b). The N–S trending normal faults are arranged in a stepwise

system towards east, which controls the morphology of the slope of the rift margin

escarpment. The major normal faults dip approximately 85°, 80° and 60° towards

the east; they cross the Yizaba Wein and Shotel Amba large-scale landslides. The

slope instability in the area is related to these faults, joints, and fracture zones.

Plots of poles cluster around the periphery of the stereographic net and exhibit

three maxima (Fig. 2.11c, d) that may suggest the occurrence of four trends of

faulting: N–S, E–W, NE–SW, and NW–SE. Plotting around the periphery of the

stereographic net indicates steep dips (Fig. 2.11d). Furthermore, the contoured

diagram exhibits three maxima that resolve into three great circle girdles. This

possibly signifies vertical to sub-vertical fractures at the intersection of joints

resulting in steep hydraulic gradient of groundwater. This situation might lead to

a rock topplings from the surrounding steep cliffs down to the river valley.


48

Figure 2.12: Lineament map of the study area.

In general, a frequency plot of all joint strike data shows widely ranging trends

that indicate fracturing of the lithologies in every direction (Fig. 2.11a). The


49

landslides displacement is orthogonal to the NNE–SSW, and N–S striking normal

fault systems that are affected by NW and NE striking trans-tensional components

(Kiros et al., 2018). The interaction of these fault systems produced a complex

displacement across and along the escarpment, manifesting oblique continental

rifting. The rock slopes stability is greatly affected by the discontinuities and their

interrelationship with the slope (Hoek and Bray, 1981). The discontinuities in the

area are generally open, smoothly undulating, lowly to highly persistent, and

intersecting each other. This is a crucial factor in slope stability, acting as either

conduit for groundwater flow or as aquitards. The NNE to NE fracture systems are

very well marked by topography breaks in the area. Landslides are strongly

oriented NNE–SSW and N–S, thus corresponding with the most consistent

lineament set.

2.5.2 Elevation, slope angle, and aspect

The distribution of landslides in relation to elevation is frequent in the middle

elevation (1800–2500 m). At intermediate elevations, slopes tend to be covered by

ignimbrite-tuff-volcanic ash, porphyritic basalt, and colluvium, which are more

prone to landsliding. The landslide risk is little to average at very low elevations

because the terrain itself is flat, although it is covered by layers of colluvial-alluvial

soils, but higher density springs and intense undercutting of the bottom of the

slope at the intermediate elevations initiate slope failure. The highest density of

landslides falls in the elevation class 1800–2500 m a.s.l., followed by elevation class

2500–3000 m a.s.l. But, the elevation class 1153–1500 m a.s.l. is characterized by

fewer landslide events (Figs. 2.3 and 2.14). The high density of landslides is mainly

related to the presence of highly fractured porphyritic basalt and highly shattered

ignimbrite as well as volcanic ashes which are susceptible to slaking. Furthermore,

the springs that emerge in the elevated parts of the area also suggest that the

hydrostatic pressure of groundwater can be a triggering factor to the landslides.

Most of the landslides occur in the slope gradients that range between 10° and 40°,

and the highest frequency of landslides is evident in the areas with slope gradient

ranging between 30° and 40° (Fig. 2.4). The density of landslides in the study area

with respect to aspect reaches the maximum values on the east-facing slopes,


50

followed by those facing southeast (Fig. 2.5). The easterly and southeasterly facing

slopes receive a high amount of sunlight and rainfall. This favors landsliding due

to fault orientation dipping towards the east, increased rate of saturation, and

weathering, particularly in loose pyroclastic sediments and colluvial deposits.

2.5.3 Rainfall

To determine the effect of rainfall on landslide occurrence in the area, the monthly

rainfall of the Debre Sina meteorological station was analyzed. The dates of

previously occurred landslides in the area were compiled from previous studies

(EIGS, 1979; Gebreselassie, 2007; Woldearegay, 2008; Abay and Barbieri, 2012),

and interviews with local residents during fieldwork. The hydrological year 1997

was exceptional in terms of the amount of rain (3593 mm). The most evident

reactivations of older landslides are the movements that occurred in spring and

summer 2005–2016 following long-lasting and above-average precipitation (Fig.

2.13). Localized landslide occurrences are common in every rainfall period in the

rift margin, including the Debre Sina area, especially along stream banks and road

cuts. The heavy rainfalls of May 6 and July 27, 2016 have triggered several shallow

to medium depth landslides in different parts of the area, mainly referable to

earth/debris flows, rockfalls, and earth/debris slides. Also, there was a flash flood

in Shewa Robit village in 2016, where the river burst its banks and caused

damages to built-up structures along the stream. Most of the landslide incidences

in the area occurred in the peak wet season (Fig. 2.13). All landslides occurred

when the annual rainfall was greater than the long-term average rainfall except

for the hydrological year 2016. This indicates that precipitation is one of the

potential triggering factors for the slope failure in the Debre Sina area. Besides,

the concave shape of the terrain is enhancing the convergence of groundwater flow

into the landslide area.

The landslides of May 6 and July 27, 2016 occurred after 24 h of continuous rainfall

of 54 mm and 60 mm, respectively. This shows that rainfall intensity alone could

be an issue for shallow landslides, but not for very large deep-seated failures. The

relationship between rainfall and landslide events indicates that deep-seated

failures prevail after a longer period of intensive rainfall. A single precipitation


51

event is unlikely to trigger a deep-seated slope failure of large extent. Most of the

landslides shown in Fig. 2.13 do not have known month or date of occurrences

except the landslides in 2016. The inlet graph of the mean monthly precipitation

in Fig. 2.13 shows periods of intensive rainfall in July, August and November

(1997). The mean annual precipitation for the whole period is 1812 mm and

marked by a horizontal black broken line, as shown in Fig. 2.13.

Figure 2.13: Rainfall data from the Debre Sina station from 1974 to 2016 compared

with landslide events.

2.5.4 Earthquakes

The central highlands of Ethiopia are in close proximity to the most seismically

active regions of the country, such as the Afar Triangle and the MER, where well-

documented damaging earthquakes are common (Samuel et al., 2012). The

northwestern plateau and southeastern plateau are split by the active East African

Rift Valley, which has a history of generating large earthquakes (Zygmunt et al.,

2014). More than 90% of the seismic and volcanic activities are connected with the

rifts, but the seismic hazards to life and property and the greatest damaging effects


52

are found on the plateaus where the majority of the population resides. The Debre

Sina area is known for its seismic activities recorded in chronicles as well as in

measured records. The earthquake events that occurred between April 1841 and

December 1842 (Gouin, 1979) along the plateau’s escarpment triggered landslides

and rockfalls that destroyed the town of Ankober, which is located nearby the study

area (Fig. 2.14). The causes of such slope failures were aggravated by heavy rains

that had saturated the thin layer of clayish soil, precariously exposed along very

steep slopes (Gouin, 1979). During February 1974, tremors were reported in the

Debre Birhan – Debre Sina region; some of these tremors were also felt in Addis

Ababa (Alemayehu et al., 2012). Such tremors were caused by a sequence of

earthquakes with magnitudes less than 4.5 originating from the region to the north

and northwest of Debre Sina. This portrays that the effects of continuous shaking,

even if light, on slopes of marginal stability, are cumulative.

According to Ayele et al. (2009), an earthquake of magnitude 5.0 MI ruptured in

an area 25 km northwest of the Ankober town at 17:56:07 GMT on 19 September

2009. The tremor was recorded in Addis Ababa, about 160 km from the epicenter.

Furthermore, the area has been reported to have experienced a 6.0 magnitude

earthquake in 1983. This is the maximum recorded from the area where the

tectonic stress is mostly released by smaller magnitude shocks (Ayele, 2009).

During the fieldwork, interviewed local people and district administrators

regarding the reactivated landslide in 2016 said that there were earthquakes

shaking on 29 October 2015, 24 January 2016 and 1 May 2016, in the Debre Sina

area and surroundings. This is in good agreement with suggestions by Kropáček

et al. (2015) that the sliding events are driven by a combination of geologic and

tectonic predispositions together with external factors such as long-term water

saturation and seismic events. Figure 2.14 shows that a concentration of epicenters

follows the MER structures. However, the epicenter distribution is more

concentrated along the northeastern and western escarpment margin. Faults of

the rift margin suffer from occasional earthquake tremors leading to the activation

of unstable ground. Furthermore, there are several earthquake tremors recorded

in the period 2005, 2006, 2016, and 2017 in the region. Out of these, the seismic

tremor known to have been felt in the area close to the time of the deep-seated


53

landslide that took place on 13 September 2005 in the Debre Sina area

(Woldearegay, 2008) is that of the major seismo-tectonic event in north-central

Afar, in September 2005 (Yirgu et al., 2006). Although the earthquake around the

time of the main deep-seated landslide occurred afterward, foreshocks of lower

magnitude could have brought about a certain degree of instability in the already

susceptible terrain.

Figure 2.14: Recorded earthquakes in the East African region from 1842 to 2011

(source: EAGLE data).

As reported by Ayele (2009) and supported by subsequent researchers such as

Yirgu et al. (2006), and Abay and Barbieri (2012), the tectonic activity and

earthquakes might also be related to the activation of landslides in the form of


54

deep-seated gravitational creep. Particularly in the Debre Sina area, the role of

tectonic activity 2016 inducing landslides has been significant. The historical

earthquakes distribution can indicate that the seismic hazard in the central-

western highlands of Ethiopia is even higher. As there are intensive urban and

infrastructural developments taking place along hills and rugged mountains in the

study area and surroundings, very serious attention is required to consider the

seismic hazard in the area during planning, design, and construction phases.

2.6 Conclusions

The study area is located along the rift margin escarpment, a zone of notably high

potential for landslides. It is flanked by the lower and upper normal faults and the

Tarmaber–Mezezo mountain chain. This article presents six typical examples of

recent slope movements in the area. The studied landslide areas have been

evaluated in terms of their historical development, current status (based on the

field survey results), and geological and topographical conditions. Deeper

landslides are found in Yizaba Wein, Shotel Amba, Nib Amba, Nech Amba, and

Wanza Beret areas, whereas shallower depth landslides are found in the Shola

Meda area.

The area is covered by volcanic rocks and superficial deposits ranging in age from

Tertiary to the recent. The landsliding events were driven by a combination of

geologic and tectonic predispositions, together with external factors such as long-

term water saturation and/or seismic events. The rocks exhibit a variety of planar

discontinuities that originated during the volcanic formation and subsequent

tectonic disturbances. The presence of highly fractured porphyritic-agglomeratic

basalt, highly shattered ignimbrite and volcanic ash, which are prone to water

absorption and susceptible to slaking, was identified as one of the reasons for a

high concentration of landslides and main triggering factors of reactivation in the

observed cases. Therefore, it is evident that the inherent variation in the physical

property of the lithologic sequence and their structures influence the slope

stability.


55

Overall assessment of the morphometric analysis revealed that the slopes ranging

from 10° to 40°, with an elevation of 1800–2500 m and aspect to east and southeast,

are highly prone to sliding. The study area is densely traversed by faults and

lineaments with a variable pattern that denotes the formation of the variable

hydraulic system affecting mechanisms of surface and groundwater paths and thus

degrades rock mass strength but also increases the weight of the slope mass, i.e.,

increasing pull of gravity. The landslide displacement is orthogonal to NNE–SSW,

and N–S striking normal fault systems affected by NW and NE striking trans-

tensional components.

In general, the deep-seated translational and rotational landslides of the area are

controlled by different predisposing factors such as (i) geological-structural setting,

(ii) complex morphology of the slope, (iii) presence of closely spaced normal fault

segments with steep slope angles, and (iv) deepening of the Dem Aytemashy,

Majete and Shenkorge streams (active erosion and gullying). This study shows the

importance of recognizing both the predisposing factors and failure mechanisms

along rift margins and highland terrains linked to deep-seated potential landslides

and their disastrous consequences.

Acknowledgments

We express our deep gratitude to the anonymous esteemed reviewers and the

editors of the Bulletine of Engineering Geology and the Environment Journal for

their constructive comments and suggestions.

Funding information

The first author would like to thank the German Academic Exchange Service

(DAAD) for the scholarship grant to pursue the PhD study. This work was

supported by the Ruhr University Research School PLUS, funded by Germany's

Excellence Initiative (DFG GSC 98/3).

Tectonic conditioning revealed by seismic refraction

56

Chapter 3

3 Tectonic conditioning revealed by seismic

refraction facilitates deep-seated landslides in

the western escarpment of the Main Ethiopian

Rift

This chapter is based on Tesfay Kiros Mebrahtu, Michael Alber, Stefan Wohnlich

(2020b). Tectonic conditioning revealed by seismic refraction facilitates deep-

seated landslides in the western escarpment of the Main Ethiopian Rift.

Geomorphology 370, 107382. DOI: 10.1016/j.geomorph.2020.107382.

Abstract

Landslide is a geo-hazard phenomenon that has been taking lives and causing

severe property damages all over the world mostly in mountainous areas. The

Main Ethiopian Rift has a unique tectonic setting with complex geological and

geomorphological features, coupled with continuously deteriorating environmental

conditions, which made its escarpments vulnerable for landslides. The study area

is located near the Debre Sina town, within the Yizaba Wein locality, which has

been severely affected by frequent landslide problems. This work was carried out

using a multidisciplinary approach based on geological, geomorphological,

kinematic analysis and geophysical survey. Seismic refraction investigations were

carried out along the Yizaba Wein landslide main scarp to determine the depth to

the bedrock and to the failure plane, to assess the stability of the slope, to locate

possible structural features and to identify the extent of recent landslide activity,

and to study the subsurface situation. The seismic measurements were made along

three nearly orthogonal survey lines in the recently affected area. A high-

resolution 2D P-wave survey was conducted using a 24-channel seismic unit. The

seismic refraction results revealed four layers of geomaterials with distinct

physical characteristics that contained a subsurface landslide anomaly within the

layers. The layers were interpreted according to the major lithological units, from

https://doi.org/10.1016/j.geomorph.2020.107382


57

top to bottom: (i) clay, loosely cemented colluvial sediments and highly weathered

agglomeratic basalt; (ii) highly to moderately fractured porphyritic basalt,

ignimbrite-volcanic ash and rhyolite/trachyte; (iii) moderately to slightly fractured

ignimbrite, rhyolite/trachyte and basalt and (iv) very strong, massive, fresh

rock/bedrock. Faults and weak zones have also been identified in the bedrock based

on the abundance of fractures and subsidiary faults resulting from damage of rocks

and change of lithology due to variable fault rock formation strongly influencing

the wavefield distribution which usually causes a local decrease of the velocity

value. The main findings show that the landslide in the Yizaba Wein locality was

caused by its complex geological-structural setting and downslope movement of the

underlying pyroclastic sediment facilitated by heavy rainfall. Considering the

similar geological and tectonic settings, similar mechanisms can be assumed for

other landslides along the rift margins and associated highlands of Ethiopia.


58

3.1 Introduction

Mass movements are among the common geohazards that cause major economic,

social and environmental problems in the world. A gravity driven down-slope

movement of soil or rock masses without requiring a fluid as a primary

transporting agent, is generally known as a landslide (Brunsden, 1999). The

unique tectonic setting and accompanying active rifting processes have exposed

Ethiopia for seismic and volcanic risks. Besides, its complex geological and

geomorphological features, coupled with the continuously deteriorating

environmental conditions, have made most parts of the country vulnerable for

landslides. Previous and recent tectonic activity are possibly directly or indirectly

influencing the occurrence of landslides by various factors. Landslides are one of

the most destructive natural hazards in the highlands and the rift margins of

Ethiopia (Ayalew, 1999; Temesgen et al., 2001; Ayenew and Barbieri, 2005).

Fatalities, destroyed infrastructure and agricultural lands as well as distortion of

the natural environment have been the consequences of several landslides (Ayalew

and Yamagishi, 2004; Woldearegay et al., 2005). Landslides are abundant and

frequent in central highlands and rift escarpments of Ethiopia, but their failure

mechanisms are not well understood yet. A single major deep-seated landslide that

took place on 13 September 2005 in the Debre Sina area demolished 1250 dwelling

houses and more than 3000 people were displaced. Moreover, one elementary

school, four churches, and four mills were destroyed (Woldearegay, 2008). The

damage caused by the landslide in the Debre Sina area and its surroundings is

immense in terms of human life, engineering structures, agricultural land, natural

environment, etc.

The magnitude of landslide occurrences and its resulting damage to infrastructure,

property, agricultural land and environment have been increasing from time to

time and there is still an active landslide hazard especially along the Yizaba Wein

locality. To protect human live and to reduce the financial loss due to landslide

incidences by appropriate mitigation measures, the driving factors of the

landslides need to be understood. A clear understanding of the geological setting,

as well as proper evaluation of the geophysical and geohydrological characteristics


59

of soil and rock masses, is the key for a slope stability assessment. Geophysical

techniques can be extremely helpful investigating landslides and have been

successfully applied to detect failure surfaces and the associated hydrogeological

regimes (Jongmans and Garambois, 2007). In landslide studies, geophysical

methods are used to determine the approximate thickness of the landslide debris

using different techniques, such as electrical resistivity tomography and seismic

refraction that are based on the determination of electrical resistivity/conductivity

and velocity of elastic waves that provide insight about characteristics of the

underlying earth materials (Pellegrini and Surian, 1996). As stated by Prokeˇsov´a

et al. (2014), geophysical investigation methods can help to locate historic and

potential landslide areas, but also in monitoring landslides during slip. Lateral

extend as well as depth to the failure surface can be determined through

geophysical investigations (Cummings and Clark, 1998; Bruno and Marillier,

2000; Mauritsch et al., 2000). Further parameters, such as groundwater levels or

moisture content, which are directly or indirectly related to slope failure, can be

derived from geophysical measurements (McCann and Foster, 1990; Hack, 2000;

Jongmans and Garambois, 2007; Heinze et al., 2019). Combined with an in-depth

understanding of the failure mechanisms and triggering factors, a characterization

of the landslide can be achieved. Reliable information on the current state, extent,

structure, sliding plane location, groundwater table are required for hazard

assessment (Bell et al., 2006). Seismic refraction has been used to identify

lithological layers, delineating failure surfaces and determining physical

properties of the landslide material (e.g., Bogoslovsky and Ogilvy, 1977; Bichler et

al., 2004; Glade et al., 2005; Otto and Sass, 2006; Göktürkler et al., 2008). Seismic

refraction is widely used in investigating the mechanical and geological conditions

which determine the nature of sliding movements.

Combined engineering geological and geophysical investigations for landslides

were conducted using seismic refraction, vertical electrical sounding (VES),

electrical resistivity profiling and magnetics in some parts of Ethiopia (Shimeles

and Getnet, 2016), but not in study area of this work. The study area was visited

by some researchers but there were no particular works on the deep-seated

landslide using geophysical methods. To deal with this problem seismic refraction


60

method was applied along with geological and geomorphological mapping and

assessment of landslide prone areas in the study area. The main goal of this work

is to identify geological and geomorphological features of the deep-seated landslide

of the Yizaba Wein locality in order to understand the predisposing factors that

control the development of the landslide, to examine the internal structure of

landslides and to assess the stability of the slope. We analysed a landslide that

occurred at the Yizaba Wein locality in 2005 to identify depth to failure plane,

possible slip surface and the mechanism of slope failures, existing situation of

faults and their continuity under the landslide, and to determine the internal

composition of the sliding masses for the characterization of similar landslides

using seismic refraction. While seismicity was recorded and reported around the

occurrence of the landslide in 2005, the focus point of this work is on the pre-

existing lithology and its relationship to the tectonic setup of the region rather

than on a possible seismic trigger.

3.2 The study area

Ethiopia is composed of four major physiographic regions, with the Main Ethiopian

Rift (MER) separating the northwestern and the southeastern plateau. In the

northeast, normal faults sloping towards the northwestern and noutheastern

plateau bound the Afar depression (Fig. 3.1). The study area is located in the

border zone of the Ethiopian highlands and the MER about 190 km to the north of

Addis Ababa (Fig. 3.1). The area is characterized by high elevation and topographic

variations, with elevation ranging from 1130 to 3696 m above sea level (a.s.l). The

area is transected by deeply dissected valleys and channels, canyons, rugged relief,

N–S trending outstanding ridge chains with steep cliff escarpments. The study

area contains several stream networks that originate in the highlands and proceed

further outwards through deep gorges and rift valleys. The average annual

temperature and rainfall is 15 °C and 1812 mm, respectively. The drainage pattern

is well defined with parallel to sub-parallel, dendritic, rectangular and trellis

patterns developed along faults and master joints in hard rocks. The most

spectacular landslide phenomena were observed at the Yizaba Wein locality, which

is to the north of Debre Sina town (Fig. 3.1).


61



Fieldwork investigation was carried out from January to February 2018 along

selected traverse lines to delineate and map the landslide features, geological, and


62

topographical settings of the area. Moreover, the delineation and identification of

past landslides was defined by Google Earth imagery. Lineaments were extracted

from Landsat Enhanced Thematic Mapper Plus (ETM+) images and fieldwork.

Structural measurements of faults and joints were carried out to reveal the

structural predisposition of slope to deep sliding. In particular, stereographic

projection kinematics analysis method was performed in the main landslide area

to analyze the potential failure modes and the corresponding probabilities. The

strike and dip of the fault planes were reported in stereographic diagrams. A

geological cross-section was constructed along the seismic refraction profile in

order to link surface and sub-surface geological data.

Seismic methods are the most commonly conducted geophysical surveys for

engineering investigations. The seismic refraction uses seismic energy, which

returns to the surface after travelling through the ground along refracted ray paths

(Kearey et al., 2002). The first arrival of seismic energy at a detector offset from a

seismic source always represents either a direct ray or a refracted ray. Due to this

fact, refraction surveys can be carried out with the focus only on the first arrival (or

onset) of seismic energy, and time-distance plots of these first arrivals are interpreted

to derive information on the depth of refractive interfaces. The seismic refraction

survey was conducted on a hummocky, sloping surface of a landslide along three

nearly orthogonal survey lines within the recently affected area. For this study

three profiles were carried out with a horizontal length of 1 km. The profiles were

orientated perpendicular to the deep-seated landslide main scarp. The first and

second profile (hereafter L1–L2) extend in the NW to SE direction having a total

length of 600 m, while the third profile (hereafter L3) stretches from SW to NE with

a horizontal length of 400 m (Figs. 3.9 and 3.10). The topographic surveying work

using a total station unit (TS) was intended to determine the direction and layout of

seismic lines, marking and labeling of seismic shot points and geophone locations

along the corresponding lines.

3.3.1 Instrumentation and field procedures

The seismic refraction study was carried out by using SmartSeis unit of Geometrics

seismograph of the United States of America (USA), which is a 24-channel


63

seismograph, microprocessor-based with the latest hardware and software and

interfaced with a plotter. This unit includes a state-of-the-art internal computer for

pre- and post-recording, filtering, and partial processing facilities. The signals

were generated using a 7 kg sledgehammer and a metal plate as a seismic source.

PE-3 geophones with a natural-frequency of 10 Hz were used to receive signals.

The seismic profiles were later processed using Rayfract TM v 3.34 (April 2016)

tomography software (http://rayfract.com), together with a classic refraction

processing software package. The survey was undertaken using in-line profiling

and reversed recording method. A spread length of 115 m with 5 m geophone

spacing was used. For a continuous subsurface mapping, successive spreads with

2 geophone overlap were used. In order to have a regression of arrival times from

the various subsurface formations, on each spread, a total of five shots were

employed where the topography allowed to do so. A total of seven shots per spread

were used with a shot point spacing of 25 m (Fig. 3.2). The shots are two offset

shots (direct and reverse), direct, central and reverse shots. The energy source used

was a 7 kg sledgehammer. Several hammer impacts were stacked for each shot

point. Mostly, the signal is stacked in the seismograph 5 to 15 times to increase

amplitude of the refracted signal and to cancel random noise, as a result first

breaks are enhanced and also signal to noise ratio is improved. The survey lines

were laid down and pegged at 5 m intervals with coordinates and elevation

recorded on each pegging point. Besides, in order to avoid noise created by different

unnecessary sources, the geophones were buried into a 20 cm deep hole. The

refraction field parameters and geometry of seismic shooting (field observations) are

summarized in Fig. 3.2 and Table 3.1.

Survey results are mainly affected by data quality, data processing, and

interpretation procedures. Data quality in the present survey area could be

influenced by ambient noise, geologic noise, surface irregularities (topographic

problems, and outcropping rock which presents problems in establishing contact

between geophones and the ground surface). The depth of investigation from the

present seismic refraction survey was attained with a maximum of 75 m.


64

Figure 3.2: The layout of the seismic refraction survey.

Table 3.1: Parameters of seismic refraction used during fieldwork.

Description Type/magnitude

Energy source Sledgehammer

Geophone distance 5 m

Length of spread 115 m

Shot position –22.5, 2.5, 27.5, 52.5, 77.5, 102.5, 127.5…

Recording Time 128 ms – 512 ms

Sampling 0.0625–0.125

Filtering Open

Amplification Auto trace (mostly)

3.3.2 Data acquisition, processing, and presentation

The field data was processed in a conventional manner, by manual work in which the

P-wave velocities determined from the first arrivals were used for the

interpretations. The first arrivals in the seismic signals are picked to determine the

transit time from shot position to the geophone. In the inversion process, a tomogram

of the P-wave velocity distribution in the subsurface is calculated based on the

determined transit times (Kearey et al., 2002). The frequency filtering options of the

instrument were utilized to eliminate unwanted signals or noise. The seismograms

were produced using the internal computer of SmartSeis exploration seismograph

with the following standard processing parameters:

Display mode: Average limited

Time scale: Normal

Print time scale: Compressed


65

Automatic Gain Control (AGC): Fixed Gain

Low cut filter: Open or 10 Hz

High cut filter: Open or 180 Hz

3.3.2.1 Conventional approach

From the first arrivals, time-distance (T–D) curves were plotted and verified for

reciprocity, parallelism, number of possible layers which could be mapped, and

presence of overlapping events for the same refractor over each spread.

Subsequently, velocity analysis was performed taking into account the survey layout

geometry, number of shots available, and number of overlapping events for each

spread. For the first and second layers the velocities have been determined using the

inverse slope of the lines of best fit of the inner time-distance plots. The noise has

been removed using the inherent filtering facility of the TERRALOC internal

computing system. The first arrival times were picked up using auto pick option of

the system for all records and few adjustments were made manually where it seemed

necessary. Various corrections were made on the recorded data (elevation, shot point,

and phase) prior to analysis, to remove influences caused by surface irregularities.

In some cases, due to the limited energy by the sledgehammer, the signal to noise

ratio was found to be small; in such cases, secondary arrivals were used to correct

the first arrivals. After plotting these corrected times against source-geophone

distance (time-distance curve) the reciprocity and parallelism between the different

refractors have been evaluated and the necessary processing was made (Figs. 3.3

and 3.4). An initial velocity-depth model was generated by sorting the travel times

and assuming several horizontal layers with constant internal velocity gradients

(Gebrande and Miller, 1985; Rohdewald, 2011).

From the travel time curves, the number of layers can be estimated as well as first

assumptions of their apparent velocities can be made. The inverse slope of the first

line segment gives the velocity of the first layer while the apparent velocity of the

second layer is determined from the inverse slope of the second line segment. The

first layer velocity (V1) was determined by averaging the inverse slope of the direct

(V1d) and reverse (V1r) recordings since those are direct waves, not refracted ones

(Eq. (3.1)).


66

V1 =(V1d+V1r)

2 (3.1)

The second layer velocity (V2) was estimated by taking the harmonious mean of

the apparent up-dip (V2d) and down-dip (V2r) velocities (Eq. (3.2)), as there has been

no sufficient overlap of travel times (Kearey et al., 2002).

V2 =2V2dV2r

(V2d+V2r) (3.2)

Depth equation at impact points for an arbitrary n number of velocity layers from

intercept time is given as follows (Eq. (3.3)) (Sjögren, 1980).

ℎ(𝑛 − 1) = 1 2⁄𝑇𝑛𝑉𝑛𝑉(𝑛−1)

√𝑉2𝑛−𝑉2(𝑛−1)−

𝑉𝑛𝑉(𝑛−1)

𝑉2𝑛−𝑉2(𝑛−1)∑ ℎ𝑣√(

1

𝑉2𝑣−

1

𝑉2𝑛)𝑣=𝑛−2

𝑣=1 (3.3)

where, Vn= velocity of the nth layer, Tn= nth intercept time and h= depth.

Depth analysis and velocity distribution in the bottom refractor are carried out by

using the T–0 method (Gurvich, 1972), and the mean-minus T and ABC-curve

methods (Sjögren, 1980; Robert, 1995). The advantage of Mean-Minus-T method

over the others is that it verifies the lateral velocity variation as accurate as

possible (Sjögren, 1984). It also unmasks the irregularity of the refractor

topography in depth determination. Velocities estimated by the above methods

were compared and calculated. This, in turn, has allowed to improve accuracies in

the determination of refractor velocities and therefore layering thicknesses.

3.3.2.2 Improved/latest approach

After the necessary partial data processing within the seismic unit, manual data

processing which includes correction of arrival times and analysis of preliminary

seismic models, a software-assisted data processing has also been carried out. This

was done using the Rayfract 32 software package. This package provides an image

of the subsurface based on seismic first break energy propagation modeling. In this

method, a starting model is calculated based on the travel times. Using the 2D

Wavepath Eikonal Traveltime (WET) inversion, the P-wave velocity distribution

was calculated based on the derived travel times (Schuster and Quintus-Bosz,

1993; Lecomte et al., 2000). In the course of the present work, the partially


67

processed, corrected and checked data have been used to produce WET

tomographic sections through smooth inversion and refined Delta-t-V inversion

(Figs. 3.15 and 3.16) methods which were in turn later refined by the WET

tomographic inversion method (Figs. 3.9 and 3.10). The refraction seismic time-

distance curves first break and their corresponding fits difference between

measured (instantaneous) and inverted (CMP layer) are shown in (Figs. 3.3 and

3.4). The P-wave velocities remaining root-mean-square (RMS) error between

modelled and measured data is less than 2%.

Figure 3.3: Time-distance curve along profile one and two (L1–L2) that black line

is the fit.

Figure 3.4: Time distance curve along profile three (L3) that black line is the fit.


68

3.4 Results

3.4.1 Geology and geomorphology of the study area

The geology of the area is mainly composed of (i) aphanitic basalt-porphyritic-

agglomeratic, (ii) ignimbrite-tuff-volcanic ash, (iii) porphyritic basalt-scoriaceous

agglomerate, (iv) Tarmaber basalt, (v) colluvial deposit and (vi) alluvial deposit

(Fig. 3.5). The aphanitic basalt-porphyritic-agglomerate units crop out in the gully

areas and series of steep cliffs and benches in deeply dissected valleys (Fig. 3.5).

The ignimbrite-tuff-volcanic ash unit crop out in lower part of the slope, underlain

by aphanitic basalt-porphyritic-agglomerate and overlain by porphyritic basalt-

scoriaceous. The ignimbrite-tuff-volcanic ash beds form small cliffs, which are

highly altered and intensely weathered, vertically jointed and highly shattered by

faulting (Mebrahtu et al., 2020a). The porphyritic basalt-scoriaceous agglomerate

unit consists of dominantly porphyritic basalt and scoriaceous agglomerate with

subordinate aphanitic basalt and vesicular basalt. These units show abundant

spheroidal weathering and breaks easily to very small size material. The Tarmaber

basalt unit is mainly exposed in the western part of the study area in the high

rising mountain chains. This rock formation forms vertical cliffs and ridges

trending in N–S direction.

The colluvial deposits cover the slopes with lower inclination, presumably

transported down-slope by gravity. They contain rock fragments and soil derived

from fragmented and weathered bedrock. In particular, the outcropping rocks in

the Yizaba area are deeply weathered and covered by colluvial materials, locally

including pyroclastic materials. The alluvial deposits are restricted along the

major riverbeds, riverbanks and their tributaries representing riverbeds and

terraces (Mebrahtu et al., 2020a). They are derived from the weathering,

transportation, and reworking of different rocks from the steep cliffs and

escarpment. The study area has been affected by N–S, E–W, NNE–SSW, NE–SW,

NW–SE, NNW–SSE, and WSW–ENE major trends of faulting (Fig. 3.5). These

faults experienced weathering down to greater depths causing thick soil layers and


69

show a high groundwater potential, which might trigger landslides through high

pore pressures and reduced friction.

The geomorphological survey was carried out to identify and map the main surface

features associated with gravitational processes of the topographical surface. The

geomorphological survey shows that various landforms have been identified in the

area, including structural, fluvial and slope forms. The morphological setting of

the area is very complex and strongly influenced by rock weathering, erosion, and

tectonics. The geomorphology of the study area is closely related to the

development of the MER system and recent river erosion. Deep fluvial dissection

has imprinted the geomorphological landscapes of the area and its surroundings.

The deep-seated landslides also modify the slope morphology of the study area. The

Yizaba–Shotel Amba slope shows clear morphological evidence of a deep-seated

landslide (e.g., breaks on the slope, depressions, terraced surfaces, and complex

convex–concave forms) (Fig. 3.6).

The slope morphology is also clearly controlled by the superimposition of mass

movements on a large-scale. Tension cracks of different size and opening are

recognized in the main body of the Yizaba landslide (Fig. 3.5). The major ground

deformation observed in the area during the fieldwork is represented by tensile

fissures, forming back scars and back tilting blocks (Fig. 3.6). Based on the

landslide classification following Cruden and Varnes (1996), the shape of the

cracks indicates translational and rotational character of the movement in the

upper part of the affected slope. Shear cracks developed systematically in sliding

colluvial deposits and pyroclastic sediments (Fig. 3.5). The Yizaba Wein landslide

can be divided into two main parts, which are indicated by the old and new scarps.

The upper part of the slope was formerly a head scarp, and the lower portion a

convex-concave slope with a gentler top surface. The new scarp runs parallel to the

topographic break of a pre-existing landslide (Figs. 3.5 and 3.6). Landslides and

deepening stream channels play an important role in the landscape development

of the study area. Below the new scarp, the bedrock is exposed and consists of

highly weathered porphyritic basalt and tuff-volcanic ash. The frontal part of the

moving mass is widely affected by multiple, retrogressive and successive debris


70

slides and earth flows induced by the lateral erosion of the stream, particularly

active in the rainy season. The instability in the Yizaba Wein locality is in the

recurrent stage with instabilities occurring on areas showing to have undergone

previous slides.

Figure 3.5: Geological and geomorphological map of the study area.


71

The upper portion of the scarp slopes varies approximately between 30° and 60° to

the east and is clear of landslide debris. The lower portion of the scarp is mantled

by a colluvial wedge.

Figure 3.6: Panoramic view of the main Yizaba Wein landslide and surroundings

from the east located in north direction of Debre Sina area.

The morphology of the area is also conditioned by numerous mass movements,

including rotational slides, translational slides, rockfalls and toppling, rock slides,

debris slides, and earth flows (Figs. 3.5 and 3.6). The spatial distribution of the

landslide types essentially depends on the type of materials involved (lithology and

structure of the bedrock, nature of overburden materials) and the slope angle. The

instability of the entire Yizaba slope system was influenced by the pressure of

active landslides from the Yizaba slope. The area is typically characterized by an

irregular hummocky topography, with deeply incised valleys and sharp steep

slopes, gentle slopes, cliffs, and escarpments crossed by the Dem Aytemashy river

and its tributaries. They are commonly deeply weathered and covered by colluvial

materials. These materials are shifted downslope by debris flow and earth flow

movements and tend to fill up small valleys and inter-hill depression (Fig. 3.5).

Moreover, large-scale slides affect the colluvial materials and pyroclastic


72

sediments, covering the hummocky topography. The colluvial deposit is diffusely

affected by shallow-deep multiple retrogressive and successive rotational slides.

The colluvial deposits are considered likely to be derived from ancient landslides,

probably associated with movement along layers of weathered tuff/volcanic ash

within the volcanic sequence.

The NE–SW trending Yizaba–Shotel Amba fault escarpment (Figs. 3.5 and 3.6) is

affected by a huge retrogressive rotational and translational rock slides and

toppling. Recent rock slides and toppling deposits derived from the overlying

porphyritic basalt cliffs occupy the margins of the valley and there are debris slides

that extend down to the river on the western side (Fig. 3.6). The intervening cliffs

are often formed in open-jointed rocks, posing slope stability problems. The gullies

that drain these valleys convey large volumes of sediment during the wet season,

mostly in the form of debris slides and flows. Springs are mainly observed in the

main scarp of the landslide and below the scarps (Fig. 3.5). The Dem Aytemashy

river, which plays an important role for the Yizaba area deep-seated landslide

development by slope undercutting, is structurally controlled by a NE–SW fault

segment that has guided the river incision and produced gully with toe erosion and

slope undercutting (Fig. 3.5). They often correspond to sets of fractures existing as

result of tectonic deformations. During the rainy season, runoff from the western

mountains forms a number of gullies. In addition to this, valley widening occurred

as a result of river erosion and landslides. Extreme undercutting of the base of the

sliding slope by the rivers contributed to the destabilization of the slope and thus

facilitated the landslide in the study area.

3.4.2 Profile one and two (L1–L2)

The profile L1 was surveyed above the main scarp (head scarp zone) towards NW

direction, whereas the profile L2 is below the main scarp (main slide block zone)

towards SE direction (Fig. 3.7a). The ground surface along the profile L1–L2 is

relatively flat in the southeastern part of the line and gently slopes in the

northwestern direction. The inverted P-wave velocity tomograms cover a broad

range of velocities ranging from 400 m/s to 5400 m/s (Figs. 3.9 and 3.15). The smallest

velocities (Vp 500 m/s) are found in the lowermost flat slope. In the P-wave


73

tomogram (Figs. 3.8 and 3.9) four different lithologies can be distinct based on

lithological outcrops in the area. The lower P-wave range ( 1000 m/s) in profile

L1–L2, reflects the response from unconsolidated sediments exposed mostly on the

lower relatively flat terrain, gently sloping elevated parts and at the top part

towards NW. Line L2 is roughly 200 m away from the main scarp covering the

damaged zones of the main body towards the SE. The top material is similar to the

sliding mass of the steeper slope, so that the lithological layers identified in profile

L1 are continued, though vertically displaced by a potential fault. Moreover, the P-

wave range of the lower layer (1000–1500 m/s) indicates a mixture of the upper

material with agglomeratic basalt. The top unit disappears between stations 400–

435 and around station 560 (Fig. 3.9), which is moved towards the lowermost gentle

slope. The unit has a thickness in the range of 7–15 m, with further increasing

thickness towards the SE of station 500. This unit is underlain by relatively

cemented highly weathered/decomposed materials and agglomerates represented

by the P-wave velocity range (1000–1500 m/s). The above two sequences of

sediments comprising the top velocity layer, with P-wave velocities below 1500 m/s,

have thickness up to 30 m.

Generally, this layer thickness ranges between 10 and 30 m excluding the above

two zones where it vanishes. Its maximum thickness is associated with the

downthrown part of the line between stations 300 and 320. The intermediate unit

with the P-wave range (1500–2000 m/s) is mapped along the entire length of line

L1. This unit is exposed to the surface around station 400, where it attains a

maximum thickness of about 30 m coinciding with a fault (Figs. 3.7a–b and 3.9).

This unit appears relatively compact with the P-wave range (2000–2500 m/s)

towards its bottom. These units, having the same range of velocities, appear to

correspond to different lithologic units as observed at the outcrops. This could be

an indication of water saturation in the soil due to seepage and drainage through

fissures and fractures. Consequently, groundwater in the elevated parts of the area

is expected to be manifested in the form of seepages (springs), which could

contribute pressure in the direction of flow which is towards the lower elevated

area, as indicated by the arrows in Fig. 3.9.


74

Figure 3.7: Typical landslides in the study area: (a) translational slides in

porphyritic-agglomeratic basalt and seismic refraction line L1–L2 location, (b)

rotational slides on colluvial deposit and volcanic ash/tuff, (c) rock slides in

porphyritic basalt and seismic refraction line L3 location, (d) earth slide in clay soil

and colluvial deposit.

Figure 3.8: Geological cross section along selected line A–B.


75

In the lower elevated sides, to the southeast of station 350, the groundwater level

could be between 2250 and 2270 m a.s.l. The intermediate unit with P-wave

velocities between 1500 and 2400 m/s may be partially saturated. The underlying

relatively higher velocities with 2500 m/s are interpreted as comprising the upper

part of the bedrock (Fig. 3.9).


profile L1–L2.

The intermediate units with a P-wave range between 2500 and 3500 m/s may be

related with moderately to slightly fractured volcanic rocks. The thickness of this

unit is in the range of 7–20 m. The P-wave velocities of this unit increase depth-

wise suggesting an increase in the degree of compaction and quality of the bedrock.

Accordingly, the P-wave velocities above 3500 m/s reflect very fine, stronger and

very sound rocks, and probably aphanitic basalt. The depth to the sound bedrock

is found within 45–75 m. The maximum depth is associated with the fault zone

around station 500 m. Moreover, geological structures are inferred around stations


76

330, 400 and 520 associated with potential future landslides and presently instable

areas (Fig. 3.9). The P-wave velocity interpretation of the three profiles is based on

lithological outcrops, surface observations and geological setup.

3.4.3 Profile 3 (L3)

The inverted P-wave velocity ranges from 400 m/s to 4400 m/s (Figs. 3.10 and 3.16).

The smallest velocities in the P-wave tomogram (Vp 500 m/s) are found in the

lowermost gentle slope. Again, the tomogram allows the identification of the same

four lithological units. Thus, the uppermost units are represented by P-wave

velocities below 1000 m/s attributed to clayey soil, highly weathered/decomposed

materials and colluvium with thicknesses generally increasing towards northeast

from about 3 m to a maximum of 20 m around station 380. Towards the NE, the P-

wave velocity is significantly reduced, coinciding with a drop in ground elevation.

The bedrock is mostly covered by colluvium deriving from the uppermost gentle

slope. At the bedrock, the P-wave velocity is strongly discontinuous, increasing

from Vp 500 m/s to Vp 1000 m/s. The underlying unit with a P-wave velocity

between 1000 and 1500 m/s has a varying thicknesses between 5 m and 25 m.

Maximum variations are noted between stations 200 and 250 where the unit

narrowly extends depth-wise towards the lower elevated side (Figs. 3.10 and 3.7c).

This unit is associated with highly fractured and weathered volcanic rocks and

colluvium.

The third lithological unit has a P-wave range between 1500 and 2500 m/s. It

reflects wider variations in the central part of the area probably due to

compositional variations. Its maximum thickness is above 45 m and situated

around station 240 where this unit extends depth-wise towards the lower elevated

zone associated with the inferred fault (Fig. 3.16). The upper part of the bedrock is

marked with P-wave velocities of 2500 m/s extending along the entire length of the

profile. The degree of compactness and quality of the rocks increases with depth as

reflected from the increasing P-wave velocities. However, a fault is found around

station 230 and the bedrock P-wave velocity is relatively lower probably attributed

to brecciating and fracturing. The depth to the relatively sound bedrock (Vp 3500

m/s) undulates between 50 m and 72 m. The maximum depth (72 m) is found


77

associated with the inferred fault in the central part of the line. In addition to the

major fault in the central area, a fracture zone and an inferred fault are also shown

around stations 125 and 370, respectively (Figs. 3.10 and 3.16).


profile L3.

3.4.4 Kinematic analysis of slope failure

Kinematic analysis was performed to determine the failure modes of the rock slope

by interpreting the stereonet. The back-analysis of the stability of the landslide

area was prior to the landslide event in 2005. In this study, Rocscience Dips Version

7.0, a graphical and statistical analysis of orientation data software, was utilized to

analyze and visualize the orientation data. In a kinematic analysis various modes of

potential rock slope failures such as planar sliding, wedge sliding and flexural

toppling that occur due to the presence of discontinuities in unfavourable orientation

can be studied (Hoek and Bray, 1981; Goodman, 1989). The stability of the rock

slopes is highly influenced by the orientation of discontinuities like joints, faults,


78

bedding planes, etc. The three faults located along the main scarp, as shown in

Figs. 3.5 and 3.11 are plotted on the stereonet according to the fault orientation

details as given in Table 3.2 and friction angle of 30°. First, the mode of failure of

each sliding has been defined by providing the kinematic properties, such as planar

sliding, wedge sliding and flexural toppling. There are two types of failure modes in

the study area: planar sliding and wedge sliding (Figs. 3.12 and 3.13). The analyses

for planar and wedge failures were done on each plot using a basic friction angle

of 30° (Barton, 1976) and mean dip direction of 85°. The friction angle is used as a

first approximation to the stability analysis of the slopes along faults in hard rock.

Figure 3.11: The main Yizaba Wein landslide located in north direction of Debre

Sina town: fault 1–green (NNE–SSW), fault 2–blue (NNW–SSE) and fault 3–

purple (WSW–ENE).

The variables that contribute to the probability of failure are slope orientations,

friction angle and lateral limits. The input data for the kinematic analysis of

planar and wedge failures are summarized in Table 3.2. Especially discontinuities


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have a strong impact on the strength of a material, as they reduce frictional and

cohesional strength and provide water pathways to deeper layers. As shown in Fig.

3.11, the slope is dominated by various parallel tensional fractures, which might

allow surface runoff to penetrate the subsurface. The interaction of these fault

systems produced complex displacement across and along the study area. The

density of the fractures in tilted basalt and ignimbrite-volcanic ash materials are

weak zones facilitating water infiltration into the volcanic ash levels. The green,

blue and purple broken lines as shown in Fig. 3.11 represent fault 1, fault 2 and

fault 3, respectively. According to the interpretation of the topographic profiles, the

maximum estimated thickness of the active main landslide body is approximately

150–200 m, whereas the maximum thickness of the old slope failure is up to 300–

400 m (Kropáček et al., 2015).

The stereonet presented in Fig. 3.12 is about planar sliding kinematic analysis

failure mode in pole vector mode. From Fig. 3.12, the region highlighted in red is

the critical zone for planar sliding where it is inside the daylight envelope and

outside the plane friction cone. Any pole failure within the daylight envelope will

not slide if frictionally unstable. On the other hand, any pole falling outside of the

represent planes with a dip steeper than the friction cone represents planes will

slide if kinematically possible. All poles that are plotted in the region in red are

representing a risk of planar sliding. The kinematic analysis result of planar

failure of the landslide event in 2005 with an average slope angle of 75° is as shown

in Fig. 3.12. For fault 3 which is circled in red, 1 out of 3 poles is in the critical

region. Therefore, the risk of the occurrence of planar sliding is about 33.33%. This

indicates that sliding along any single fault plane is likely to occur in the geological

plane having a dip direction of 90°.

Fig. 3.13 depicts the stereographic projection of the wedge sliding kinematic

analysis. In wedge sliding, a wedge is formed through intersection of two planes

and slides along those intersections. Wedge failure occurs when the line of

intersection dips steeper than the friction angle. The kinematic is mainly

controlled by orientation and geometry of the slope plane and the intersections as

well as the friction along the plane surface. The slope plane defines the day-


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lighting condition (Fig. 3.13). Any intersection point which plots outside the pit

slope great circle satisfies the daylighting condition. The plane friction cone is the

angle measured from the equator of the stereonet. The critical zone for wedge

sliding is the red area in Fig. 3.13. If the intersection point falls in the red area in

Fig. 3.13, the intersection of two joint planes could cause a wedge sliding. On the

other hand, the secondary critical zone is highlighted in yellow in Fig. 3.13.

Figure 3.12: Stereonet of planar sliding kinematic analysis.

The kinematic analysis result of wedge failure of the landslide event in 2005 with

an average slope angle of 75° is as shown in Fig. 3.13. The risk of the occurrence of

wedge sliding is about 66.67%. The wedge formed by the intersection of fault 2 and

fault 3 is more critical than fault 1 and fault 3 as it lies within the primary critical

zone for wedge sliding. In both cases the wedge sliding takes place predominantly

along the plane formed by fault 3 as the dip vector of this plane is larger than the

friction angle and at the same time, the differences between dip direction of the

slope and fault plane 3 being less than the lateral limit of 30°.


81

Figure 3.13: Stereonet of wedge sliding kinematic analysis.

Table 3.2: Kinematic analysis of planar and wedge failures.

SID Dip direction Dip Trend Plung

Fault 1 105 85 285 5

Fault 2 65 80 240 10

Fault 3 85 60 270 30

Mean 85 75 265 15

Variance 400 175 525 175

According to Kropáček et al. (2015), the Yizaba Wein landslide of 2005 had an

estimated depth of 150–200 m and a volume of 1.7 km3, affecting an area of 6.5

km2. The rose diagram of the lineaments orientation data indicates four trends;

N–S, E–W, NE–SW and NW–SE (Fig. 3.14). Among these the N–S and E–W trends

are widespread in the area with the highest frequency. The discontinuities in the

area are generally open, smoothly undulating, lowly to highly persistent, and

intersecting each other (Mebrahtu et al., 2020a). The average spacing of the

discontinuity sets ranges from 0.23 to 0.54 m.


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Figure 3.14: Rose diagram showing strike direction.

3.5 Discussion

Four soil and rock layers incorporating subsurface landslides were identified by

the seismic refraction method. The seismic refraction study revealed zones of

overburden material comprising, from top to bottom: clay, loosely cemented

colluvial sediments and highly weathered material (Vp 1000 m/s) of 7–15 m

thickness and highly weathered agglomeratic basalt (1000–1500 m/s) with

thicknesses up to 30 m. These units are highly susceptible to sliding when it gets

moist, because the volcanic ashes are prone to slaking and acts as lubricant

material (Fig. 3.9). The intermediate units comprise weathered and fractured

volcanic rocks (1500–2000 m/s) with a maximum thickness of about 30 m and

moderately fractured porphyritic basalt, ignimbrite-volcanic ash and

rhyolite/trachyte, relatively compact (2000–2500 m/s). The depth to the upper part

of the bedrock marked by the P-wave velocity of 2500 m/s is in the range of 40–50

m. The third unit probably comprises moderately to slightly fractured (2500–3500

m/s) ignimbrite, rhyolite, trachyte and/or basalt. The thickness of this unit is

between 7 m and 20 m in profile L1–L2 and 5–35 m in profile L3. The quality of the

bedrock increases with depth from slightly weathered/fractured rocks to fresh rocks


83

probably comprising aphanitic basalt (Vp 3500 m/s). Depth to very fresh sound

bedrock ranges between 45 m and 75 m (Figs. 3.15 and 3.16).

The 2D P-wave tomography section showed in Figs. 3.15 and 3.16 allows to draw

assumptions about the layers stratification. Each layer shows minor distortions in

its morphology but in general the lithology along profile L3 oriented in NE–SW is

similar to the ones seen in L1 and L2, oriented NW–SE. This agrees well with the

assumed continuity of the major lithological units. The result obtained from the

2D profile model is in good agreement with the results obtained from the geological

mapping. Further, the measured P-wave velocities of the different geological units

coincide with previous obtained wave velocities in the literature (Ellis and Singer,

2007; Schon, 2011). The morphology in combination with the tectonic assemblage

and the intense weathering processes strongly favors the mass movement. The

seismic refraction line shown in Fig. 3.15 emphasized a scarp in the bedrock which

is located along the prolongation of a NNE–SSW trending fault. The scarps and

linear depressions developed as a result of deep-seated gravitational slope

deformation that preceded the failure. The landslides occurred in the lower parts,

gravitationally deformed slopes leaving unstable slopes above. Field observations

confirmed that many types of landslides are densely distributed in the porphyritic

basalt, ignimbrite, volcanic ash, and colluvial deposit.

It is known, that seismic refraction is very suitable to reveal the sliding surface

due to the usually significant velocity contrast between sliding mass and bedrock

(Heincke et al., 2006; Donohue et al., 2012; Yamakawa et al., 2012). From a geo-

hazard point of view, the subsurface for profile L1–L2 is divided into three parts:

Zone 01 (area to the SE of station 300), relatively stable bedrock, with minor

undulation and no significant geological structure; zone 02 (area between stations

300 and 420), gently sloping ground surface with recently moved mass, and not yet

stable. The direction and slip surface of the landslide that has taken place in 2005

is situated above the position of the interpreted major fault (Fig. 3.9, arrow A). The

inferred fault around station 330 may also mark an old slide boundary. Zone 03

(the area between stations 420–520): sloppy terrain, embodied between two faults,

bedrock with 3500 m/s not seen or deep, favourable morphology, and unstable zone.


84

The velocity model essentially shows low-velocity slide material overlying a

concave, higher velocity layer (Fig. 3.15). The high-velocity material at the base of

the profile is interpreted to represent bedrock and dips southeastward towards the

river. Fig. 3.9 arrow B shows possible future slide slip direction which may coincide

with the inferred fault around station 520. The slip surface generally coincides

with the 2000 m/s isoline (marked by yellow broken line as shown in Figs. 3.15 and

3.16) due to the presence of highly fractured and saturated nature of the

underlying rocks.

Figure 3.15: Delta-t-V inversion along profile one and two (L1–L2).

The seismic refraction tomography shows distinct refractors which are interpreted

as the surface of the bedrock. Most likely, this surface corresponds to the main slip

surface of the landslide. The porphyritic-agglomeratic basalt, ignimbrite-volcanic

ash and colluvial deposit dipping downslope favor the development of rather deep-

seated rotational slides over discrete strata and have been significantly reactivated

in the last years. In the new scarp and its upper part, the input of debris from the

upper head scarp zone is a relevant process. This occurs by the progression of the

marginal portion of the sliding colluvial sediments and volcanic ash slab (Fig.


85

3.15). The upper layers of unconsolidated deposits and porphyritic-agglomeratic

basalt rocks experiences significant water transit towards the deeper layer of

ignimbrite-volcanic ash. However, the pyroclastic sediments are impeding the

vertical percolation of rainwater due to their low permeability and hence force the

rainwater to flow laterally. The volcanic ash beds also act as excellent slip surface,

especially when they are weathered. The increasing pore-water pressure and

seepage force in the colluvial deposit may favor slope instabilities. This shows that

the sliding surface is controlled by the ignimbrite-volcanic ash as the ignimbrite-

volcanic ash causes full saturation due to tail water corresponding to increased

pore pressures due to its low hydraulic conductivity. In general, based on the

seismic refraction survey in the Yizaba Wein locality, the study site can be

characterized as a deep-seated landslide including bedrock and surficial deposits.

Fig. 3.15 shows that a very clear linear trend between stations 200 and 250 is

marked by the intermediate and lower P-wave velocities (1000–2000 m/s)

associated with the inferred fault. The uneven surface of the bedrock (3300–3500

m/s) in Figs. 3.15 and 3.16 could be interpreted as a continuation of fracturing and

displacement. The zone to the northeast of this profile is vulnerable to sliding,

which is favoured by the possible groundwater pressure that flows from elevated

areas and circulates through the structure as well as morphology. Furthermore,

the subsurface condition for the profile L3; zone to the NE of station 200 is

vulnerable to sliding. Sliding is anticipated to occur along the slip surface and

direction indicated by the arrow around station 200 (Fig. 3.16). Based on the

tomogram and field observations, the active slide mass is interpreted to be

superimposed on a larger slide, with fracturing and displacement extending into

bedrock. Besides, the subsurface unit's classification was made based on the P-

wave velocities (Table 3.3). This classification is based only on the actual seismic

results, surface observations and local geological setup of the study area. The study

area experiences high tectonic activity with intense fracturing due to its location

at the western margin of the MER. The result obtained from the 2D profile models

agrees well with the geological mapping and field observations. As observed in the

curves of Figs. 3.15 and 3.16, strong lateral variations are present along the 2D


86

seismic refraction profiles. These variations are most likely caused by tectonic

fractures affecting subsurface units before the movement.

Figure 3.16: Delta-t-V inversion along profile three (L3).

The presented seismic sections revealed that geological structures are inferred in

profile L1–L2 around stations 330, 400 and 520 associated with old, present and

future slides. Moreover, two geological structures are identified in profile L3

around station 230 (a major fault) and station 370. In addition to this, a fracture

zone is also indicated around station 125 in profile L3. It is known that faults have

a passive role in the development of large gravitational deformations (Di Luzio et

al., 2004; Bois et al., 2008; Agliardi et al., 2009). Figs. 3.11 and 3.12 depict that the

deep-seated landslides displacement which occurred in 2005 is mainly controlled

by NNE–SSW, NNW–SSE, and WSW–ENE trending lineaments which relate to

accommodative fault zones.

After heavy rainfall, high pore-water pressure can build up in the faults. The pore-

water pressure is the main triggering factor for rock slides, rockfalls, and toppling

in the study area. This condition reveals the predisposing role of the structural

setting of the entire slope on the landslide's kinematics. This is in good agreement


87

with Woldearegay (2008) showing that the landslide is located in a tectonically

active and extensive area. The geological structures are coinciding with the head

scarp of the old and new failure planes, and several slide incidences also are

occurred (Fig. 3.11). Surface deformation enhance and accelerate surface processes

(Densmore, 1997; Strecker and Marrett, 1999; Ambrosi and Crosta, 2006; Bucci et

al., 2013; Scheingross et al., 2013). Landslides play a prominent role in the present-

day geomorphological evolution of the Yizaba Wein area and its surroundings. The

landscape development of the study area was dominated by mainly fluvial and

mass movement processes.

Table 3.3: Classification of the various subsurface units on the basis of their

compressional velocities.

Main

Layers

Layering

sequence

P-wave

velocity

(m/s)

Included lithologic units Nature

Layer 1

Uppermost < 1000

Clay, loosely cemented colluvial

sediments and highly

weathered/decomposed materials Soft formation

(easily sliding)

Lower part 1000–1500

Highly weathered/ decomposed

materials, loosely cemented colluvium

and agglomeratic basalt

Layer 2 Intermediate

(upper part) 1500–2500

Highly to moderately fractured

porphyritic basalt, ignimbrite and

rhyolite/trachyte, agglomerates,

cemented tuff

Slightly hard to

moderately hard

(sliding)

Layer 3 Upper part 2500–3500

Moderately to slightly fractured

ignimbrite, rhyolite, trachyte and/ or

basalt

Hard formation

(mostly stable)

Layer 4 Bottom > 3500 Very strong, massive, fresh basalt Very hard

formation (stable)

Highly permeable fissures and cracks can alter the pore-water pressure of clays

soil quickly triggering slope failure (Iverson et al., 1997; Van Asch et al., 1999). In

this regard, the upper part of the slope is dominated by soil and rock types of high

permeability. Rainwater is infiltrating through these slope masses and then

migrating laterally downward into the middle and lower sections of the slope. The

springs observed at the base of the slope provide the pore pressure regime, which

determines critical conditions for the study area slope stability. The seismic

refraction results convey that groundwater is expected within the highly

weathered and fractured units associated with geological structures represented

by a P-wave velocity range (1500–2500 m/s). The aforementioned seismic refraction


88

agrees well with the typical field observations in the study area. There is an

emergence of springs and seepages at the contact of between porphyritic-

agglomeratic basalt and ignimbrite-volcanic ash. This shows that the basalt and

ignimbrite rocks are acting as a pathway for groundwater, while the volcanic ash-

tuff is acting as a barrier for the flow of water within the unstable slopes. Thus,

the volcanic ash which is prone to water uptake can cause high pore pressures

facilitating landslides. During the field investigation landslides manifestation

triggered by rainfall were identified in the area. As reported by Abay and Barbieri

(2012) and Alemayehu et al. (2012), heavy rainfall was the ultimate mobilization

of the landslide.

A small graben feature occurs on the crest of the main block which shows evidence

of seasonal pond development (Fig. 3.11). It appears that the majority of present-

day ground movements relate to the reactivation of existing landslide deposits and

the gradual regression of head scarps associated with extremes of rainfall and river

erosion. Besides heavy rainfall, tectonic and seismic dynamics cause slope failure,

too. Landslides concentrate along tectonically active mountains zones (e.g., Dramis

and Sorriso-Valvo, 1994; Strecker and Marrett, 1999; Alexander and Formichi,

2006) as the fault scarps represent steep slopes with tectonically disturbed

lithology. The central highlands of Ethiopia are close to the Afar Triangle and the

MER where well-documented devastating earthquakes are common (Samuel et al.,

2012). As reported by Ayele et al. (2009) and others, such as Yirgu et al. (2006),

and Abay and Barbieri (2012), there might be a coupling between tectonic activity

and creeping deep-seated landslides. This is in good agreement with suggestions

by Abebe et al. (2007). Although the earthquake in September 2005 around the

time of the main landslide occurred afterwards, foreshocks of lower magnitude

could have brought about a certain degree of instability in the already susceptible

terrain.

High magnitude earthquakes trigger landslides and induce significant

morphological changes to a large area, particularly when the epicenter is located

in a mountainous terrain (Jibson et al., 2006; Wang et al., 2009; Yin et al., 2009;

Gorum et al., 2011; Wartman et al., 2013). Particularly in the Debre Sina area, the


89

tectonics has a great impact on slope failure, as the area is in close proximity to

the most seismically active regions in the country. In the town of Ankober, which

is close to the study area, an earthquake of M6.0 in 1983 as well as M5.0 in 2009

were recorded (Ayele et al., 2009). There are also several smaller events with

magnitudes between 4 and 5. From this perspective, the Yizaba case study shows

that tectonic activity plays a well-defined role in promoting landslides with

seismicity as a possible predisposing factor and by determining the lines of

weakness along which the landslides may have developed. While not focus point of

this study, there is evidence of seismic activity as a triggering factor in the study

area as indicated by interviews from local inhabitants and large rock mass slides,

which are known to be associated with seismic activity (Gouin, 1979; Hearn, 2018).

However, for the study area landslide, the complex geological-structural setting,

slope gradient, heavy rainfall and the subsequent water pressure lead to an

increase in shear stress triggering slope failure. For the future, comprehensive

data generation using borehole drilling, periodic monitoring of the rate of

movement and additional geophysical data collection using different techniques

has to be conducted to have a wider and more complete view of the area from a

mechanical and geological point of view.

3.6 Conclusions

This study shows that the seismic refraction is an important technique to assess

the stability of the slope and to understand the predisposing factors that control

the development of the landslide. The tomogram of P-wave velocity reveals the

lithological layers of the landslide. The depth of investigations from the present

seismic refraction survey was attained with a maximum of 75 m. The seismic

refraction data shows that the currently active landslide is superimposed by a

larger slide including parts of the bedrock. The geomorphological analysis showed

that the complex landslide sloped surfaces in the detachment zone are associated

with a hummocky and step-like morphology as a result of successive or

retrogressive sliding. Landslides and deepening stream channels play an

important role in the landscape development of the study area. It has been shown

that the 2D P-wave velocity tomographic section reveals the slip surfaces,


90

geological structures contributing to the old and present slides, failure

mechanisms and influencing factors. In addition, failure-prone areas are indicated.

The kinematic assessment shows that the rock slope has a higher probability of

failure in the wedge sliding failure mode (66.67%) compared to planar sliding

(33.33%). The results of the kinematic analysis manifested a planar failure along

fault 3 and wedge failures due to intersection of faults 1, 3 and due to intersection

of faults 2, 3. The stability of the slope is largely controlled by the geology, structure

and slope gradient, particularly the ignimbrite-volcanic ash formation, which is

highly susceptible to sliding. The seismic sections across the main landslide scarp

show highly permeable unconsolidated deposits and highly fractured and

weathered porphyritic-agglomeratic basalt overlying ignimbrite-volcanic ash with

a shallow water table. Such type of geological setup facilitates significant seepage

forces and increased pore pressure within the unconsolidated deposits and the

volcanic ash during the rainy season. The vertical and sub-vertical faults and

fractures provide a favourable regime for surface and groundwater flow and also

facilitate water seepage. The intense fracturing can create weak zones that

accelerate the infiltration of water which can be responsible for the build-up of high

hydrostatic pressure resulting to lowering of normal stress in the rock mass giving

rise to landslides. Therefore, it is evident that the complex geological-structural

setting, slope gradient and hydrogeological conditions of the terrain contribute

significantly to the Yizaba Wein locality landslide failure.

The high-resolution 2D P-wave tomographic survey has provided useful

information related to the mechanical and geological conditions of the subsurface

that could be among the landslide causing conditions in the investigated area. These

results are another great example for the suitability of refraction seismic for

landslide investigations. The multidisciplinary approach using geomorphological,

geological and geophysical data proved sufficient to elucidate the connection

between tectonic and landslides. The study area is in close proximity to one of the

most seismically active regions in the world. Therefore, we recommend evaluating

seismic activity and its effects on deep-seated landslides to obtain a complete view

of the area from a tectonic perspective. The applied technique allows


91

understanding both the deep-seated landslides and the present-day

geomorphological characteristics of the area. The integration of 2D seismic

refraction into geomorphological studies enables the interpolation of a geological

model of the landslide. The application of geophysics, especially refraction seismic

in landslide studies and geomorphology can help to answer unresolved questions

in geomorphological research, such as sediment thickness and subsurface

conditions. Therefore, the applied methodology can be recommended as a tool for

assessing slope stability conditions as well as for planning possible mitigation

measures.

Acknowledgments




Excellence Initiative (DFG GSC 98/3). Many thanks also go to Dr. Tesfaye

Asresahagne (Geomatrix Plc) for providing field logistics. We express our deep

gratitude to the anonymous esteemed reviewers and the editors of the

Geomorphology Journal for their constructive comments and suggestions.

The effect of hydrogeological and hydrochemical dynamics on landslide

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Chapter 4

4 The effect of hydrogeological and

hydrochemical dynamics on landslide

triggering in the central highlands of Ethiopia

This chapter is based on Tesfay Kiros Mebrahtu, Andre Banning, Ermias Hagos,

Stefan Wohnlich (2021). The effect of hydrogeological and hydrochemical

dynamics on landslide triggering in the central highlands of Ethiopia. Hydrogeol J

29, 1239–1260. DOI: 10.1007/s10040-020-02288-7.

Abstract

The volcanic terrain at the western margin of the Main Ethiopian Rift in the Debre

Sina area is known for its slope stability problems. This report describes research

on the effects of the hydrogeological and hydrochemical dynamics on landslide

triggering by using converging evidence from geological, geomorphological,

geophysical, hydrogeochemical and isotopic investigations. The chemical

characterization indicates that shallow to intermediate aquifers cause

groundwater flow into the landslide mass, influencing long-term groundwater-

level fluctuations underneath the landslide and, as a consequence, its stability.

The low content of total dissolved solids and the bicarbonate types (Ca–Mg–HCO3

and Ca–HCO3) of the groundwater, and the dominantly depleted isotopic

signature, indicate a fast groundwater flow regime that receives a high amount of

precipitation. The main causes of the landslide are the steep slope topography and

the pressure formed during precipitation, which leads to an increased weight of

the loose and weathered materials. The geophysical data indicate that the area is

covered by unconsolidated sediments and highly decomposed and weak volcanic

rocks, which are susceptible to sliding when they get moist. The heterogeneity of

the geological materials and the presence of impermeable layers embodied within

the highly permeable volcanic rocks can result in the build-up of hydrostatic

pressure at their interface, which can trigger landslides. Intense fracturing in the


93

tilted basalt and ignimbrite beds can also accelerate infiltration of water, resulting

to the build-up of high hydrostatic pressure causing low effective normal stress in

the rock mass, giving rise to landslides.


94

4.1 Introduction

Rainfall-triggered landslides occur frequently in the central highlands of Ethiopia.

These highlands are highly populated regions in which more than 60% of the

country’s population is settled. The mean annual rainfall in these regions exceeds

1200 mm and accounts for some 70% of the total precipitation the country receives

each year (Ayalew, 1999). The highlands are highly rugged. The topographical

variation, land-use, geology, and the surface water and groundwater flow systems

here are strongly characteristic for this type of region, as are rainfall-triggered

landslides (Woldearegay, 2013). Most of the landslides in these highland regions,

including the largest ones, are triggered by heavy precipitation occurring at the

end of the rainy periods in July and August (Ayalew, 1999). In addition, fast-

moving slope failures, such as rock slides, occur due to seismic triggering by

earthquakes from the Afar depression (Fig. 4.1), the Main Ethiopian Rift (MER)

and their escarpments (Abebe et al., 2010). There were also reactivated landslides

during the fieldwork in this research conducted from April to June 2016 and

October to November 2017, following heavy rainfall and earthquake incidents.

Rainfall-induced mass movement hazards in the Debre Sina area are closely linked

with hydro-meteorological hazards such as slope instability and erosion. This is

shown by the strong association of landslides with streams or river incision and

with gully erosion. During the past few years, natural disasters within the central

highlands of Ethiopian have increased in both frequency and intensity, and have

had severe social impacts. According to previous studies (e.g., Ayalew, 1999;

Temesgen et al., 2001; Ayalew and Yamagishi, 2002; Ayenew and Barbieri, 2005;

Woldearegay et al., 2005), the landslides have severely affected human lives,

infrastructures, agricultural lands, and the natural environment in various parts

of the highlands and rift margins of Ethiopia. Besides rising public awareness of

the landslide hazard, there has been little to no changes in land-use or other

protective measures in this region.

Hydrogeological data, and hence hydrogeological assessment, on the large-scale

and deep-seated landslides are extremely scarce in the study area and other parts

of the central highlands of Ethiopia. The infiltration of precipitation into the


95

subsurface is one of the main factors that initiates and controls the mobilization of

rock (Iverson, 2000). Groundwater-level rise is often the critical factor for slope

failure because it induces high pore-water pressures which can reduce the

frictional strength of slopes. As stated by different researchers (e.g., Bogaard et al.,

2000; Tullen et al., 2002; Malet, 2003; Lindenmaier et al., 2005), large landslides

usually imply a complex hydrogeology with various flow paths. Defining the origin,

age and ongoing processes of groundwater flow while it passes through a landslide

area, can contribute to understanding the hazard (de Montety et al., 2007).

Hydrogeochemistry can also be an important contribution to characterize

landslides, as hydrogeochemistry and isotope ratios can be characteristic (Epstein

and Mayeda, 1953; Tóth, 1999; Guglielmi et al., 2000; Wang et al., 2001; Guglielmi

et al., 2002; de Montety et al., 2007; Cervi et al., 2012). Groundwater that flows

within volcanic rocks alters the chemical composition of the groundwater itself,

with influences from the precipitation, mineralogy of the watershed aquifers,

climate, topography, and anthropogenic and volcanic activities (Edmunds et al.,

1992). The interaction of groundwater with these factors leads to the formation of

different hydrochemical facies which can be correlated with location, geology,

climatic conditions and topography (Clark and Fritz, 1997). A number of research

projects in different countries have shown that the patterns of hydrogeochemistry

and δ18O and δ2H isotopic compositions in the water can provide a useful tool for

landslide investigations (Di Maio et al., 2004, 2014; Gaucher et al., 2006; Picarelli

et al., 2006; de Montety et al., 2007; Calmels et al., 2011; Cervi et al., 2012; Vallet

et al., 2015), while there have been no similar studies undertaken in Ethiopia and

particularly in the Debre Sina area, where the study area resides.

Landslide incidences in the central highlands of Ethiopia and the western margin

of the Ethiopian Rift escarpment are increasing at an alarming rate. Therefore,

investigation of hydrogeological dynamics is vital for understanding the influence

of water on the potential to move mass. Nevertheless, hydrogeological data

associated with large-scale and deep-seated landslides are extremely scarce in the

study area. Considering the scale of the landslide problems and the socio-economic

development in the area, there is an urgent need to understand the hydrological

processes, to evaluate the soil/rock-water interactions and to determine the nature


96

of sliding movements. These are essential for appropriate hazard maps and

realistic predictions, as well as for developing systems for early warning of

landslide hazards in the margins of the western Afar depression. The study area

is hardly accessible and the active movement of landslides may quickly destroy

instrumentation. Therefore, to identify the most relevant influencing factors, a

comprehensive study of the geology, groundwater flow conditions, respective rock-

water interactions, and geophysical investigations was conducted. Thus, the

objective of this study is to implement a converging evidence approach towards

understanding the main landslide triggering factors and the land-mass failure

mechanisms through (i) detailed study of the geological and structural settings of

the study area by using remote sensing data, field geological mapping and

geophysical techniques (ii) conceptualization of the hydrological processes and

rock-water interactions by using hydrogeochemical and isotope approaches and

(iii) integration of the geological, hydrogeological, hydrogeochemical and isotope

data. The data presented and the issues raised in this paper will also be useful for

similar studies across the East African Rift and in similar tectonic settings.

4.2 The study area

The study area is located in the central-western highlands of Ethiopia, forming

spectacular escarpments along the margins of the southwestern Afar depression,

which is tectonically active (Fig. 4.1). It is geographically bounded by UTM

1077165 m N and 1108635 m N, and the UTM 571065 m E and 601125 m E. The

steep escarpment and the narrow strip of the plateau over-look the Afar

depression. The steep mountain chains and rugged valleys that drain into the

central-eastern and western lowlands of Ethiopia characterize the wider area. The

elevation ranges from 1,130 m above sea level (a.s.l.) in the southeast and

northeast parts, to 3,696 m near Tarmaber on the plateau (Fig. 4.1). The high-

elevation ridge chain occupies the western part of the study area and represents

the area's highest peaks. It includes highly elevated and N–S trending outstanding

ridge chains with steep cliff escarpments.


97


The Ethiopian highland is characterized by variable climatic conditions. The

climate of the study area is significantly colder and wetter than the rest of Ethiopia


98

due to the high elevation and high gradient. The area has sub-humid to humid

climate and bi-modal type of rainfall (rainy months: March to Mid-May and Mid-

June to September). The annual rainfall distribution is characterized by

pronounced seasonality, with the heaviest rains occurring in July and August. The

area has an annual average precipitation of about 1812 mm which is estimated

based on 43 years of complete precipitation records (Mebrahtu et al., 2020a). In

general, the western highlands bounding the rift valley receives high rainfall,

above 1200 mm/year, whilst the rift floor gets little seasonal rain, often less than

600 mm/year. The air temperature has a maximum value of 25 oC and a minimum

value of 10 oC. The mean annual temperature is 15 oC. The lower and middle parts

of the area are densely populated and massively cultivated. People in this region

are still actively involved in agriculture. The area is characterized by deeply

dissected valleys and channels, rugged relief, mesas, plains, high-elevation

continuous ridge chains with steep cliff escarpments, highly variable topographic

features and complex geology, which reflect the past geological and erosional

processes (Mebrahtu et al., 2020b). This implies that the area is subjected to

dynamic geomorphic processes of erosion, transportation, and material deposition.

Stream networks originate in the highlands, and then they proceed further

outwards through deep gorges towards the rift valley. The drainage pattern is well

defined with parallel to sub-parallel dendritic patterns developed along faults and

master joints in the hard rocks.

Intensive rainfall induces fast-moving slope failures, which have affected the

Debre Sina area several times in recent years. The increasing impact of

anthropogenic activities (land-use changes, especially deforestation and intensive

agriculture, quarrying, road construction, urbanization, etc.) has also contributed

to slope instability and landslide hazards over the last two decades (e.g., Ayalew,

2000; Nyssen et al., 2003; Zvelebil et al., 2010). Settlements at the foot of steep

slopes and close to the streams that carry flood flows and debris from adjacent

mountains are especially in danger. As mentioned by Woldearegay (2008), the

localities Yizaba Wein and Shotel Amba areas were strongly affected by a single

major deep-seated landslide that took place on 13 September 2005 and the slope

instability problem still remains very active (Fig. 4.2). The Debre Sina landslide


99

that reactivated during summer 2005 had existed for the previous 15 years. Since

the original landslide activity started, the slide has continued to move at a

relatively high rate. The most common types of landslides in this area are

rotational slides, translational slides, rockfalls and toppling, rock slides, debris

slides, and debris and earth flows (Fig. 4.2).

Figure 4.2: (a) Panoramic view of the main Yizaba and Shotel Amba landslides

from east, with examples of characteristic geodynamic features within the main

landslide body and its surroundings: (b) rotational slide, (c) rock slide, (d)

debris/earth slide, (e) debris flow, (f) earth flow, (g) translational slide occurred in

2005 and (g) large-scale sliding.

4.2.1 Geological setting

The study area is marked by its complex lithological and tectonic settings. The

Cenozoic era is characterized by extensive faulting accompanied by widespread

volcanic activity and uplift. The area is represented by two major litho-

stratigraphic formations, which are the Tertiary volcanic rocks associated with

volcanic ash and the Quaternary superficial deposits. The major rock and soil types


100

in the area include: aphanitic basalt, porphyritic basalt and agglomerate basalt

(aphanitic basalt–porphyritic–agglomerate); ignimbrite–tuff–volcanic ash;

intercalated porphyritic basalt and scoriaceous agglomerate (porphyritic basalt–

scoriaceous agglomerate); Tarmaber basalt; upper ignimbrite; and unconsolidated

deposits (colluvial and alluvial deposits; Fig. 4.3). The volcanic rocks have

experienced intense weathering, which resulted in the occurrence of deep

weathering profiles and weathered landforms. The aphanitic basalt–porphyritic–

agglomerate units crop out in the gully areas and series of cliffs and benches in

deeply dissected valleys. The unit exhibits notable textural and compositional

variations vertically, which are constituted by the aphanitic basalt, porphyritic

basalt, and scoriaceous agglomerate basalt.

The ignimbrite–tuff–volcanic ash unit mainly consists of pumiceous lapilli tuff and

volcanic ash with subordinate ignimbrite, trachyte, and rhyolite. The ignimbrite–

tuff–volcanic ash beds form small cliffs that are highly altered and intensely

weathered, and are vertically jointed and highly shattered by faulting (Mebrahtu

et al., 2020a). The porphyritic basalt–scoriaceous agglomerate unit consists of

dominantly porphyritic basalt and scoriaceous agglomerate with subordinate

aphanitic basalt and vesicular basalt. The porphyritic–scoriaceous agglomerate

basalt shows a high rate of spheroidal weathering and breaks easily to very small

sized material, and the weathering and fracturing prevails more in the major joints

and layering. This unit is highly weathered and fractured, favoring the circulation

and storage of subsurface water. The Tarmaber basalt unit is mainly exposed in

the western part of the study area in the high-rising mountain chains (Fig. 4.3).

This rock formation forms vertical cliffs and ridges trending in the N–S direction

as well as some E–W offsets and shows well-developed columnar joints (Mebrahtu

et al., 2020a). The upper ignimbrite unit is exposed in the western part of the study

area overlying the Tarmaber basalt. This unit is fine-grained, highly weathered,

and crossed by sub-vertical to vertical fractures (Mebrahtu et al., 2020a).

The slopes with lower inclination are covered by Quaternary sediments. The

colluvial deposits are associated with rock pediments originating mainly from the

basalt, presumably transported downslope by the action of gravity and slope wash


101

(Mebrahtu et al., 2020a). These colluvial deposits mainly contain rock fragments

and soil derived from fragmented and weathered bedrock.


2020a).


102

The alluvial sediments are deposited along the major riverbeds and convey large

volumes of sediment during the wet season, mostly in the form of debris slides and

flows. They are derived from the weathering, transportation, and reworking of

different rocks from the steep cliffs and the escarpment. The area is traversed by

four major trends of faults (N–S, E–W, NE–SW, and NW–SE) (Fig. 4.3) and they

can be assumed to be a major conduit for groundwater flow. However, this behavior

can be sometimes lost or sealed by precipitation of secondary material or clay

(Guglielmi et al., 2000). The landslides coincide with the tectonically active

geological structures.


4.3.1 Water sampling and analytical methods

A multi-techniques investigation strategy combining hydrogeochemical, isotopic

and geophysical methods, was followed in this study. Groundwater chemistry and

stable isotopes analyses were used to characterize the groundwater flow system

and rock-water interactions. The water samples were collected in two field

campaigns (April–June 2016 and October–November 2017) from 65 sites. The

samples were taken directly from two different sources, which include cold springs

and rivers. Spring water samples were collected directly at their discharge points

under natural pressure by using a plastic syringe. The water sampling point

locations were systematically selected in order to be representative of: different

rock formations in the stratigraphic column; recharge and discharge areas; and

landslide areas and surroundings. The locations of the sample sites are shown in

Fig. 4.3, and the hydrochemical results, including isotope data, are presented in

Table 4.1. All the water samples were filtered through a 0.45 µm membrane on site

and filled 50-ml polyethylene bottles. The samples taken for major cations analysis

were acidified to pH 2 with HNO3 (nitric acid). All the major ions except HCO3–

were analysed using ion chromatography (Dionex 1000 Ion Chromatography

System) in the laboratory of Applied Geology at the Ruhr University of Bochum

(RUB), Germany. HCO3– was analysed in the field by employing a burette titration

method by using HCl, and the total Fe (Fetot) was determined by atomic absorption


103

spectrometry 240F (AAS) in the RUB. Measurements of electrical conductivity

(EC), pH and temperature were made in-situ by using WTW Multi340i handheld

meters, and each electrode was calibrated. Conventional field hydrogeological

observations and a landslide inventory were done to support the results from

hydrochemical and isotope analyses.

The stable isotopes (18O and 2H) were measured in 39 water samples (33 cold

springs and 6 rivers). The sampling bottles were repeatedly rinsed with the water

to be sampled and then completely filled leaving no space for air. Water samples

for stable isotopes (18O and 2H) were collected in high-density polyethylene bottles

(50 ml) and analyzed following standard procedures at the laboratory of Isodetect


Earth Sciences at the Technical University of Darmstadt, Germany. The

determination of 18O and 2H in groundwater was carried out by using a laser

absorption device (PICARRO L2130-i δD/δ18O Ultra High-Precision analyzer). The

measured values of the samples (mean of 10 individual measurements) were

calibrated with international standards (Standard Light Antarctic Precipitation

(SLAP), Standard Mean Ocean Water (SMOW), Greenland Ice Sheet Precipitation

(GISP)) and any drift or memory effects were corrected. The general measurement

error is ± 0.1‰ or ± 0.5‰ (standard deviation) based on the Vienna Standard Mean

Ocean Water (VSMOV). The measurement inaccuracy (simple standard deviation)

of the analysis carried out reached a maximum of ± 0.09‰ for 18O and ± 0.3‰ for

2H.

Long-term isotopic data of rainfall (from 1961 to 2016) from the Addis Ababa (Fig.

4.1) Global Network of Isotopes in Precipitation (GNIP) station (190 km from the

study area) is taken from the International Atomic Energy Agency database

(IAEA, 2020). The resulting stable isotope data are interpreted by plotting them

with the Global Meteoric Water Line (GMWL) (Craig, 1961), and the Local

Meteoric Water Line (LMWL) of Addis Ababa (Kebede et al., 2008). In this study,

Statistica version 8.0 was used to conduct hierarchical cluster analysis (HCA). The

softwares ArcGIS 10.5 (Esri), Geochem (US Geological Survey), computer program

Diagrammes v 6.5 and CorelDRAW X7 were used for database creation, spatial


104

data analysis and to prepare high-quality maps and illustrations. The final results

and interpretations were then combined to formulate a conceptual hydrogeological

model of the Debre Sina landslide area.

4.3.2 Geophysical survey

4.3.2.1 Data acquisition and processing

Vertical electrical soundings (VES) were applied to approximately delineate

horizontally layered strata and to investigate the vertical layering. The electrical

resistivity data were collected using the ABEM Terrameter SAS 4000/ SAS 1000

with steel electrodes, cables on reels and other accessories. Four electrodes were

placed along a straight line on the earth surface. Current was injected into the

earth through two electrodes (A and B) and the resulting voltage differences were

measured at two potential electrodes (M and N). The Schlumberger array (A M N

B) was used, with the distance between current electrodes five times the one of the

voltage electrodes. The VES was carried out on the profile lines with AB/2 and

MN/2 spacing ranging from 1.5 m to 220 m and 0.5 m to 20 m, respectively. The

resistivity data were collected at eight points, two of which were in the Yizaba area,

while the remaining six were in the Armaniya area (Fig. 4.3). The apparent

resistivity (ρa) data were then plotted against the electrode spacing (AB/2) in order

to obtain a resistivity-depth model for iteration on the IPI2win software (IP2win,

2003). The iterations were completed once a RMS error 5% was obtained. The

final RMS errors in this study vary between 1.59% and 4.41%. Finally, the results

were interpreted both qualitatively and quantitatively. In the quantitative

interpretation, a pseudo-depth section and geo-electric section were created. The

measured potential difference demonstrates the effects of different geological

materials within the area. The raw resistivity data are plotted as a pseudo-depth

section using the IPI2win software, which demonstrates the vertical variation of

measured resistivity as a function of electrode spacing (AB/2) and guides the

construction of the geo-electrical section using surfer 17 software and MATLAB.

Based on the geo-electric sections, the subsurface structures are quantitatively

characterized.


105

Table 4.1: Hydrochemical and isotope data of sampled groundwater and surface water in the Debre Sina area. SP spring; R river; Ionic concentrations are measured in mg/L.

Sample X_coord Y_coord Elev EC pH T TDS Li+ Na+ K+ Mg++ Ca++ Fetot HCO3- F- Cl- NO3

- SO4-- δ2H δ18O d-excess

ID (m a.s.l.) (µS/cm) (°C) (mg/L) [‰] [‰] [‰]

SP01 582657 1091105 2850 131 7.5 18.0 181 0.1 5.5 1.1 5.4 13.5 0.1 145.0 0.1 1.0 7.9 1.3 -13.91 -5.08 26.80

SP02 582895 1089302 2826 101 6.9 20.4 158 0.1 4.2 0.9 3.1 12.5 0.1 120.0 0.1 1.3 12.3 3.3 -8.27 -4.48 27.60

SP03 583879 1089927 2616 115 7.4 17.1 131 0.5 7.8 3.0 3.0 12.0 0.2 90.0 0.2 1.6 8.8 3.9 -9.27 -4.36 25.60

SP04 583689 1090650 2626 118 7.9 16.9 136 0.1 5.3 1.6 3.7 13.6 0.1 95.0 0.1 2.7 10.8 3.4 -5.49 -3.94 26.00

SP05 583689 1090882 2585 113 7.6 16.9 168 0.1 5.0 1.5 3.7 13.4 0.1 130.0 0.1 1.9 9.7 3.0 -8.37 -4.15 24.80

SP06 585290 1088159 2425 158 6.3 18.7 156 0.5 28.0 4.0 2.0 12.0 2.3 95.0 0.2 2.0 5.3 7.3 -5.91 -4.02 26.20

SP07 586172 1088369 2316 259 6.6 20.9 234 0.1 9.1 3.4 9.1 30.1 0.1 150.0 0.1 5.0 19.4 7.5 -4.39 -3.72 25.40

SP08 586780 1088368 2306 220 6.4 21.4 215 0.5 11.0 6.2 6.2 26.0 0.1 135.0 0.3 3.7 21.0 5.0 -5.53 -3.82 25.00

SP09 586018 1091691 1975 307 8.1 24.2 399 0.1 15.8 6.9 9.8 34.4 0.1 310.0 0.3 3.4 12.0 6.6 -1.42 -2.91 21.80

SP10 586202 1092581 2035 315 7.9 22.8 442 0.1 48.0 4.4 2.7 20.7 0.1 345.0 0.6 3.9 10.1 7.2 -3.86 -3.50 24.10

SP11 585012 1092580 2129 153 8.0 20.0 220 0.5 8.3 3.0 5.0 19.0 0.1 160.0 0.4 2.6 17.0 4.1 -6.13 -4.02 26.10

SP12 585193 1094664 2200 153 6.6 26.3 141 0.5 7.4 5.0 4.0 18.0 0.2 85.0 0.2 3.0 12.4 5.4 -5.78 -3.91 25.50

SP13 586690 1094093 1887 215 6.5 23.1 195 0.5 14.0 4.0 5.9 25.0 0.1 120.0 0.6 4.8 13.5 7.4 -4.49 -3.06 20.00

SP14 588535 1093884 1988 376 7.5 22.7 329 0.1 18.8 2.2 10.9 47.0 0.1 220.0 0.4 11.2 13.9 4.8

SP15 592420 1097151 1531 544 7.2 25.3 423 0.5 26.0 4.0 14.0 64.4 0.1 260.0 0.3 12.6 28.0 13.1 2.39 -1.94 17.90

SP16 591183 1097502 1473 418 8.4 25.8 343 0.5 35.0 5.3 12.0 48.0 0.1 210.0 0.6 8.8 8.7 14.7 8.21 -0.87 15.10

SP17 591693 1099083 1549 542 8.3 28.4 427 0.1 52.5 1.0 18.0 40.0 0.1 235.0 1.0 17.0 41.0 22.0 7.64 -0.73 13.50

SP18 572226 1096951 2934 155 7.3 17.8 123 0.5 7.0 0.5 4.0 17.0 0.1 83.0 0.3 3 7.5 1.0

SP19 589347 1090612 2128 297 6.8 21.4 266 0.5 31.0 7.2 9.4 23.0 0.2 180.0 0.5 3.1 8.9 2.4 -3.74 -3.36 23.20

SP20 594605 1092884 2125 286 7.2 25.8 419 0.5 21.0 5.0 8.4 40.0 0.1 310.0 0.4 10.1 10.6 13.5

SP21 582566 1088267 2806 119 6.6 15.3 184 0.1 6.1 1.2 3.4 14.3 0.1 140.0 0.1 1.4 14.1 3.8

SP22 582341 1091954 2826 139 7.6 15.7 216 0.1 5.9 1.9 4.1 17.6 0.1 155.0 0.1 1.4 28.0 1.6

SP23 582218 1092826 2805 125 8.1 16.2 215 0.1 6.2 1.3 3.7 15.0 0.1 170.0 0.1 1.0 16.2 1.4 -9.96 -4.49 25.90

SP24 582326 1094739 2580 99 8.1 17.8 172 0.1 5.0 1.9 3.3 11.3 0.1 140.0 0.1 1.6 6.5 2.4

SP25 581491 1095576 2532 72 6.4 15.7 129 0.1 5.6 2.2 2.1 7.9 0.1 90.0 0.1 1.0 16.4 3.6

SP26 584064 1094702 2420 185 7.7 16.3 108 0.5 5.8 2.8 4.0 15.0 0.1 73.2 0.2 3.3 0.1 3.0

SP27 584335 1094375 2269 465 7.0 19.4 128 0.5 14 4.9 3 14 0.1 79 0.6 4.7 4.6 3.3 -3.9 -2.24 14.0

SP28 583844 1094191 2416 183 6.6 18.7 137 0.5 6.7 2.9 4 16 0.1 70 0.2 2.9 31 2.6

SP29 583980 1094020 2375 216 6.8 21.5 127 0.5 6.4 9.8 4 15 0.17 70 0.2 7.4 9 4.4 -5.9 -2.80 16.5

SP30 584250 1093933 2317 207 6.5 21.3 121 0.5 5.9 7.7 4.0 17.0 0.5 61.0 0.2 3.8 4.4 16.9 -7.05 -2.96 16.64

SP31 584847 1094360 2234 177 6.4 19.6 103 0.5 7.3 4.5 4.0 15.0 0.1 61.0 0.3 3.0 2.5 4.9 -5.92 -2.85 16.87

SP32 586002 1093953 2073 373 7.0 22.0 262 0.5 16.0 1.1 9.5 38.0 0.1 189.1 0.3 2.2 0.1 5.7


106

Sample X_coord Y_coord Elev EC pH T TDS Li+ Na+ K+ Mg++ Ca++ Fetot HCO3- F- Cl- NO3

- SO4-- δ2H δ18O d-excess

ID (m a.s.l.) (µS/cm) (°C) (mg/L) [‰] [‰] [‰]

SP33 583842 1092047 2562 199 7.7 17.1 97 0.5 5.8 1.6 4.0 18.0 0.1 61.0 0.1 1.1 3.6 1.2

SP34 587303 1090769 2167 205 8.0 16.8 109 0.5 6.0 2.4 5.0 17.0 0.1 73.2 0.2 1.5 0.1 3.3

SP35 586248 1091063 2143 307 9.8 21.3 201 0.5 11.0 3.1 8.1 30.0 0.1 140.3 0.3 1.9 0.7 4.0 -1.58 -2.18 15.84

SP36 587111 1084884 1870 284 8.4 21.4 177 0.5 11.0 2.9 7.1 28.0 0.7 103.7 0.1 1.9 2.6 18.9 -3.11 -2.53 17.14

SP37 582718 1091290 2864 177 7.6 14.8 99 0.5 5.0 1.4 4.0 16.0 0.1 67.1 0.1 1.4 2.4 1.4 -13.29 -4.10 19.53

SP38 583970 1090132 2621 193 7.9 16.1 127 0.5 9.5 11.3 3.0 12.0 0.1 73.2 0.1 13.9 0.1 3.5 -11.03 -3.80 19.39

SP39 583808 1089437 2642 222 7.9 17.5 141 0.5 11.0 2.0 5.1 19.0 0.1 97.6 0.1 2.3 2.2 1.2

SP40 581361 1085598 2715 176 6.9 13.8 118 0.5 4.0 5.6 5.0 15.0 0.1 73.2 0.1 10.6 1.5 2.9 -10.10 -3.79 20.22

SP41 581761 1086736 2821 155 6.9 13.4 97 0.5 4.0 0.6 3.0 13.0 0.1 67.1 0.1 2.7 3.9 1.8 -8.78 -3.54 19.54

SP42 582096 1087709 2845 234 6.6 14.8 159 0.5 5.9 3.9 6.7 23.0 0.1 103.7 0.1 10.6 1.7 3.3 -8.73 -3.58 19.88

SP43 585110 1087341 2424 294 6.9 22.4 191 0.5 8.6 2.0 8.4 30.0 0.1 115.9 0.2 9.1 12.3 4.2 -4.30 -2.53 15.95

SP44 580555 1088521 3173 109 6.9 18.0 96 0.5 4.0 0.6 3.0 13.0 0.1 67.1 0.1 2.7 3.9 1.8 -15.20 -4.20 18.40

SP45 582739 1091179 2960 173 7.4 17.0 115 0.5 5.2 1.0 3.8 15.0 0.1 85.4 0.1 0.9 2.7 0.9 -14.60 -4.10 18.20

SP46 598008 1105381 1273 535 7.3 24.0 522 0.5 47.0 3.9 14.8 53.1 0.8 329.4 0.1 22.6 8.0 42.7 -1.20 -1.60 11.60

SP47 581091 1093710 3113 94 6.6 16.3 63 0.5 3.0 0.5 2.0 10.0 0.1 16.0 0.1 7.0 20.0 4.0

SP48 579777 1099074 2900 87 7.1 15.8 70 0.5 5.0 0.9 2.0 9.0 0.1 45.0 0.1 1.0 4.4 3.0

SP49 579605 1100941 3005 83 6.9 18.0 61 0.5 2.0 2.0 1.0 8.0 0.1 35.0 0.1 3.0 4.4 6.0

SP50 579807 1099181 2918 214 7.5 15.6 178 0.5 11.0 5.0 4.0 24.0 0.1 128.0 0.4 3.0 1.3 2.0

SP51 582334 1091496 3021 141 7.4 17.2 113 0.5 5.0 0.6 4.0 17.0 0.1 79.0 0.2 1.0 5.8 1.0

SP52 576252 1093000 3003 161 7.3 18.0 127 0.5 8.0 3.0 4.0 18.0 0.1 85.0 0.4 4.0 3.1 2.0

SP53 582409 1091394 2979 172 8.0 17.0 138 0.5 6.0 2.0 5.0 19.0 0.1 98.0 0.3 4.0 0.4 4.0

SP54 581506 1089390 3147 116 7.9 15.0 92 0.5 5.0 0.9 3.0 13.0 0.1 61.0 0.1 1.0 5.8 2.0

SP55 582066 1091594 3058 160 8.0 18.0 127 0.5 7.0 2.0 4.0 19.0 0.1 89.0 0.3 3.0 2.2 1.0

SP56 582643 1089410 2873 174 7.4 21.0 135 0.5 6.0 3.0 4.0 19.0 0.1 89.0 0.2 5.0 5.3 4.0

R01 583771 1090042 2596 93 7.4 17.8 110 0.1 5.0 1.7 2.5 10.1 0.2 75.0 0.1 1.7 9.9 4.1 -4.21 -3.82 26.30

R02 583413 1091305 2557 122 7.8 15.8 199 0.1 6.4 1.7 4.0 15.0 0.1 160.0 0.1 2.4 8.0 1.3 -11.46 -4.72 26.30

R03 586865 1091394 2015 268 7.8 22.0 235 0.1 18.2 7.3 6.7 27.5 0.1 150.0 0.4 3.4 9.2 12.8 -1.69 -2.66 19.60

R04 586773 1092303 1825 234 8.0 20.7 206 0.5 22.0 4.0 5.7 24.0 0.2 130.0 0.4 2.4 8.1 8.8 -2.84 -3.36 24.10

R05 587526 1094130 1692 244 7.2 24.5 246 0.5 22.0 3.0 5.5 27.0 0.1 170.0 0.4 2.6 6.6 8.6

R06 583253 1092922 2507 166 8.1 15.0 86 0.5 5.7 2.1 4.0 13.0 0.1 54.9 0.1 1.2 2.9 1.6

R07 584282 1093360 2275 190 7.4 19.1 93 0.5 5.0 1.7 4.0 13.0 0.1 61.0 0.2 2.2 3.5 2.4 -5.34 -2.97 18.38

R08 585977 1092051 1939 241 8.4 22.0 214 0.5 19.0 3.4 5.2 23.0 0.1 146.4 0.3 5.6 2.9 8.2 -4.01 -2.50 15.97

R09 591792 1100433 1433 573 8.5 25.5 470 0.5 40.0 4.3 15.0 59.6 0.1 311.1 0.6 12.9 15.8 11.2

Slope stability analysis of deep-seated landslides

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4.4 Results and discussion

4.4.1 Aquifer system and groundwater flow

The groundwater flow in the study area is controlled by geological structures,

topography and rock type. The groundwater flow direction in the whole basin

coincides with the topography following the surface-water flow direction (Fig. 4.4)

because small intermittent and particularly perennial rivers form local drainage

basins and shallow aquifers. The flow is partly controlled by the structure and

partly by the geomorphology of the area. Local groundwater flow directions vary

from place to place according to the local topography (Fig. 4.4). The groundwater

divide between the Rift Valley and the Jemma basin do not generally conform to

the surface-water divide; the divide is slightly shifted to the west into the Jemma

basin particularly in the northwestern part of the study are.

The lithostratigraphic, geomorphologic, isotopic and hydrochemical evidence

indicates that two groundwater flow systems (shallow/local and intermediate-deep)

exist in the study area. The shallow groundwater flow is mainly localized to the

highland areas and adjacent escarpments and its water table is a subdued replica

of the surface topography (Figs. 4.5 and 4.6). It is generally characterized by lower

concentrations of dissolved ions, depletion in heavy isotopes and higher d-excess.

A significant part of the groundwater is discharged to rivers (in the form of base-

flow) and as contact springs within the highland plateau and its margins. The

intermediate-deep groundwater flow is strongly influenced by the

lithostratigraphy and the major faults in the area rather than the surface

geomorphology. At the top of the slope, groundwater directly flows through the

vertical to sub-vertical joints with different flow paths mainly guided by the highly

permeable gravitational features that correspond to interconnected tension cracks.

This intermediate-deep groundwater is recharged through the deep-seated

fractures adjacent to the major faults in the highland plateaus and discharges

mainly in the form of high-discharge springs and base-flow in the eastern lowland

sections of the Dem Aytemashy, Robi, Majete and Shenkorge rivers (Figs. 4.4 and

4.6). The shallow aquifer system is drained by the perennial springs located at the


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top of slopes, at the basal aquifer in the lower part of the slopes and at the landslide

toe.

Figure 4.4: Groundwater level map and groundwater flow directions based on

spring and river positions.


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The Fig. 4.4 depicts that there is a swamp area (discharge of groundwater) around

Argaga/Asfachew in the north-central part of the area, which corresponds well with

field observations in the study area. This marshy area is developed where the rocks

are impermeable and rock-water intact near the surface. Around Yizaba and

Majete areas, the groundwater contours are closed indicating flow from all

directions towards the center (Fig. 4.4). As the water table drops below the stream

level, water infiltrates from streams and rivers into the aquifer. Aquifers along the

rivers are recharged by the surface water of streams, and the flow of many streams

is controlled by geological structure. The area is characterized by large faults that

play important roles in the occurrence and movement of the groundwater. The

plateau volcanic rocks retain rainwater for a long time and create favorable

conditions for infiltration through a highly weathered, jointed and permeable

upper layer. The shallow groundwater is partly drained by rivers and the

remaining water recharges the underlying aquifers. The highly permeable fracture

network facilitates subsurface flow to the lowlands as the primary recharge source

of the deeper aquifers.

The highly fractured volcanic rock of the plateau, consisting of basalt, ignimbrite,

rhyolite and/or trachyte, is one of the major water-bearing formations in the area.

It also covers large gently-to-steeply undulating areas of the eastern part of the

area. In accordance with to the distribution of springs (Fig. 4.3), the inter-bedded

volcanic rocks of the ignimbrite–tuff–volcanic ash act as a semi-confined aquifer.

The vertical to sub-vertical joints and tensional features in the Tarmaber basalt

covering the plateau area create a favourable condition for rainwater percolation.

Most springs are located at topographic breaks, such as hillsides.

Generally, there is a clear zonation in the total ionic concentration of natural

waters following the direction of groundwater flow from the highlands to the lower

elevations. This zonation corresponds with the spatial variations of recharge and

discharge conditions and the geological setting. The slight increase in the total

ionic concentration towards the lowland implies that the residence time of the

groundwater and the magnitude of rock-water interaction are likely to increase in

the same direction. The faults in the area are not only weak zones, but also mostly


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characterized by deeper weathering and higher potential for concentrated

groundwater flow, which can act as a lubricant and produce water pressure,

causing landslides. The most favorable condition for landslides in the Debre Sina

area and its surroundings is considered to be the fractured state of the bedrocks,

especially near the tectonic lines. As a result, most mass movement occurs in the

NNE–SSW and N–S directions, which coincides with set of lineaments. The lapilli

tuff, tuff breccia and tuffaceous strata within the pyroclastic unit make the strata

susceptible to slaking. The slaking can also be one of the triggering factors of

landslides in the area. Triggering mechanisms can also be aggravated by the

development of pore-water pressure, seepage forces, seepage erosion and

mechanisms related to high plasticity. The hydrogeological conditions of the

terrains are generally favourable for the development of seepage forces within the

pyroclastic sediments (tuff and pumice horizons) and unconsolidated deposits

during periods of rainfall.

4.4.2 Hydrogeochemical facies

All the groundwaters and the surface waters in the area are fresh, characterized

by low total dissolved solids (TDS) ranging from 61 to 522 mg/L. The pH values

show that the groundwater is slightly acidic to alkaline (6.3 – 9.8) in springs and

rivers (Table 4.1). The chemical groundwater types of an area can be distinguished

and grouped by their position in a Piper diagram (Piper, 1944). Different

hydrochemical facies were identified in the study area on the basis of the Piper

diagram (Fig. 4.5). There are four major water types, identified as Ca–Mg–HCO3,

Ca–HCO3, Ca–Mg–Cl–SO4 and Na–HCO3, classified according to their dominant

chemical composition (Fig. 4.5). Groundwater and surface water from the higher

elevations typically have a Ca2+(Mg2+)–HCO3– hydrochemical facies, whereas

groundwater in the lower altitude displays a Na+–HCO3– type (SP46). There is a

general compositional change from a Ca2+(Mg2+)–HCO3– type water to a Na+–

HCO3– hydrochemical facies along the groundwater flow path from higher to lower

altitude (Fig. 4.6). This result is consistent with other hydrochemical studies

conducted along the central-western highlands and margins of the Afar depression

(Darling et al., 1996; Chernet et al., 2001; Ayenew, 2005).


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As topography controls the fluxes in the hydrological cycle, it also controls the

hydrochemical signature of the groundwater. The TDS of the groundwater

increases towards lower altitude as the hydrochemical facies changes along its flow

paths. The low TDS and bicarbonate groundwater type in the highland part

indicate the fast hydrogeological regime of the plateau receiving a relatively high

volume of precipitation. The TDS content increases along the flow direction as

water flows from the recharge to the discharge areas. In the study area, Ca–Mg–

HCO3 is the dominant water type in the basic volcanics and Na–HCO3 in the acidic

volcanic rocks. In general, the TDS increases from the infiltration area along the

watershed on the plateau to the drainage area formed by the valleys of the Robi

River, Shenkorge River, and their tributaries (Fig. 4.4).

Figure 4.5: Piper diagram showing compositions of different water types in the

study area.

From west to east the Na+ concentration increases due to cation exchange.

Similarly, there is a facies change from being slightly mineralized in the west, to a


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significantly mineralized water type in the east. High Na+ and K+ concentrations

in springs located in the lower part of the landslide show that the water was in

contact with acidic volcanic rocks at the head scarp. In addition to cation exchange,

weathering of silicate minerals controls the hydrogeochemical facies. The

hydrochemical data provided useful insight into the main hydrogeochemical

processes involved in the water mineralization. Water groups represented by Ca–

Mg–HCO3 are weakly mineralized waters within the basaltic and scoriaceous

aquifers. Water groups represented by Ca–Na–HCO3 and Ca–HCO3 are draining

the fractured rhyolites, ignimbrites, tuff, and trachytes, and, as could be expected,

have a more dilute chemistry. The Na–Ca–Mg–HCO3 and Ca–Na–Mg–HCO3 water

types are mixtures of the water types Ca–Na–HCO3 and Ca–HCO3. Ca–Mg–HCO3

and Ca–HCO3 groups represent shallow groundwater circulation and short

residence time that contain early stages of geochemical evolution (recent recharge)

or rapidly circulating groundwater that has not undergone significant rock-water

interactions (Edmunds and Smedley, 2000; Kebede et al., 2005; Kebede et al.,

2008).

Figure 4.6: A schematic cross section (W–E), showing the hydrogeological

conceptual model of the Debre Sina landslide. The location of the cross section and

its view direction is shown in Fig. 4.3.


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The location of the landslide occurs within a formation that is poorly welded and

composed of tuffaceous material that readily weathers to clay minerals and is

capped by pervious basalt and/ or ignimbrite and colluvial deposit (Fig. 4.6).

4.4.3 Mechanisms controlling water chemistry

Gibbs plots are employed to understand the processes affecting the geochemical

parameters of groundwater (Gibbs, 1970, 1971). In these diagrams, TDS is plotted

against the concentrations of Na+/ (Na+ + Ca2+) for cations, as well as TDS versus

Cl−/ (Cl− + HCO3−) concentrations for anions. From these diagrams the natural

mechanism controlling groundwater chemistry, including the rock-weathering

dominance, evaporation and precipitation dominance, can be derived. The Gibbs

plot of samples from the study area (Fig. 4.7) shows that all of the groundwater

samples fall into the rock-weathering dominance group. The results indicate that

the surface water had active interaction with groundwater, since the samples are

not located in the rainfall dominance cell. All data points in the domain of rock-

water interaction (Fig. 4.7a, b) indicate that chemical weathering controls water

chemistry.

Figure 4.7: Gibbs diagrams for (a) cations and (b) anions indicating rock-water

interaction as the major process regulating the chemistry of the groundwater in

the study area.

As stated above, the interaction between rocks and water results in leaching of ions

into the groundwater system, which influences the water chemistry. The chemistry

of the spring water is mainly controlled by the residence time and the intensity of


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recharge. The upslope springs show low-mineralized water types whereas the

springs at the toe of the landslide area show higher mineralized water.

4.4.4 The implications of groundwater dynamics with landslides

Hierarchal cluster analysis (HCA) is a typical multivariate statistical algorithm

that puts observed data into meaningful clusters in their hierarchal order (Davis,

2002). Three groundwater groups have been identified from the preliminary HCA

based on major-ion chemistry (Na+, K+, Mg2+, Ca2+, HCO3–, SO4

2–, F–, Cl–) of the

water samples collected in this study (Fig. 4.8). The groundwater samples from the

higher-altitude areas lie within the low-EC group I (Fig. 4.8). Most samples that

lie in this group are characterized by low concentrations of all major ions. These

samples are located in the highlands bounding the Rift Valley and are

characterized by low salinity, with TDS below 216 mg/L. They were collected

mainly from basaltic and scoriaceous aquifers, indicating fast groundwater flow.

The vertical/sub-vertical joints and tensional fractures create a favourable

condition for rainwater percolation. This area is generally acting as a recharge zone

for the surface water as well as subsurface water that flows to the down-slope

areas. This low EC (72–222 µS/cm) characteristic arises from the sample-point

location within the recharge area (low residence time) and the presence of aquifer

material with lower solubility. This indicates that groundwater in the highland

areas is getting recharge from rainwater.

The samples in group II have an EC range of 115–465 µS/cm and low

concentrations of all the major ions similar to group I, but they are found at middle

altitudes. These samples were collected close to the escarpments, and recharge

seems to have taken place by precipitation over the highlands and transported

through large faults. They were collected from highly fractured and shattered

ignimbrite, rhyolite, trachyte associated with basalt, indicating that the

groundwater movement is shallow to intermediate. However, they have relatively

higher concentrations of Na+, K+, Cl– and SO42– as compared to group I, which is

mainly related to the solution or interaction between water and secondary

minerals or clay that precipitate into faults. The local freshwaters in the middle

altitude are controlled by the normal faults ultimately derived from fast circulating


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recharge from the high rainfall of the plateau. Groundwater mainly emerges as

high-discharge cold springs on the slope and at the bottom hill of the escarpment

formed by steep faults. The groundwater samples in group II therefore correspond

to groundwater in the intermediate flow systems. In the middle part of the study

area, transitional types Ca–Na–HCO3 and Ca–HCO3 occur.

Figure 4.8: Categorization of the water samples resulting from a preliminary

hierarchy cluster analysis (HCA) based on major ions chemistry using the complete

linkage rule and Euclidean distances.

The water samples of group III were collected from the lower-elevation areas

(below 1500 m asl) in the eastern and northeastern parts of the study area, which

are covered with volcanic ash-dominated units and sporadic colluvial-alluvial

deposits. In these litho-units, the groundwater movement is slow, which together

with the presence of soluble minerals, enhances the effects of rock-water

interaction giving rise to relatively higher concentrations of Na+, K+, Cl– and SO42–

. The EC values of the groundwater samples within this group are between 215

and 573 µS/cm and increase towards the Shewa Robit valley. This indicates that


116

there is intermediate to deep groundwater circulation and relatively higher

residence time of the groundwater. The low hydraulic gradient of the groundwater

in the lowland plain (Fig. 4.4) also indicates slow groundwater velocity. This leads

to the longer residence time and enhancement of rock-water interaction. Sodium

bicarbonate-rich groundwaters as well as higher sulphate concentrations were

found in this discharge area. But there are also localized freshwaters at the lower

elevation, indicating that there is also fast circulating recharge from the high

rainfall of the plateau along regional faults.

The average isotopic composition of the water samples collected from the study

area is − 5.70‰ for δD and − 3.31‰ for δ18O, which is not very far from the long-

term weighted average isotope composition of the summer rainfall for Addis Ababa

IAEA station (Kebede et al., 2008). This suggests that the groundwaters in the

study area are mainly recharged from the summer rainfall on the highlands under

cold air conditions. Therefore, they are generally of meteoric origin and they are

not affected by some processes (like evaporation) during or before recharge. The

residence time is short, the soil/rock-water interaction is low, and the water is little

mineralized mainly in the highland and intermediate regions. Therefore, it is

possible to conclude that the main cause of the landslide is not the active soil/rock-

water interaction. It is rather because of the steep slope topography and the

pressure formed during precipitation, which leads to an increase in the weight of

the loose and weathered materials (increasing its shear stress). The material loses

its shear resistance which finally results in land mass failure or landslide. The

springs and water ponding in the study area are usually seen at the upper failure

section or main scarp of landslides, and their discharge comes from the overlying

unit with a high discharge (Fig. 4.9).

Intermittent springs emerge along the highly-conductive layers of porphyritic-

agglomeratic basalt, as well as where there is an intersection with the low-

conductivity pyroclastic layers (Groups I and II, Fig. 4.9). In areas where such

highly permeable zones/layers are covered by colluvium, the groundwater can build

up pore-water pressure from below and favour the triggering of shallow landslides.

Below the spring horizons, humid zones are also formed which could additionally

favour landslide triggering, especially during heavy or long-term rainfall. Such


117

ponded water (Fig. 4.9e) can infiltrate into the slope and increase pore-water

pressure which decreases the shear strength, thereby causing instability to the

slopes. In many of the landslide-affected sites, springs and seepage zones were

observed to emerge along more fractured zones of the rocks or along the coarser

soil horizons (Fig. 4.9a–d).

Figure 4.9: Pictures of typical landslide localities in the Debre Sina area: (a)

emerging springs in ignimbrite-volcanic ash/tuff, (b) spring water at the contact of

the top layer (colluvium) and underlying altered tuff, (c) seepage spring at the

highly fractured ignimbrite, (d) spring water outflows from the bottom of the

landslide and (e) ponded spring water at the toe of the landslide.


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4.4.5 Evidence from isotopic signatures

4.4.5.1 Groundwater recharge

Groundwater recharge depends on the intensity of rainfall, permeability of the

lithological units, and the topography that controls the groundwater infiltration

and surface runoff. Knowing the origin of groundwater can help to understand the

cause of the slope instabilities and to evaluate the influence of water on the moving

mass. Groundwater recharge in the study area is mainly from precipitation. The

water vapour from the moisture sources undergoes isotope fractionation before

becoming rainfall during transportation towards the continent (Dansgaard, 1964).

During this process, δD and δ18O values in rainwater can be correlated with the

relationship δD = 8δ18O + 10, given by the Global Meteoric Water Line (Craig,

1961). Furthermore, the isotopic composition of rainfall is dependent on a number

of factors such as altitude, latitude, season, temperature and rainfall amount

(Ayenew et al., 2008; Girmay et al., 2015).

The δ18O and δD values in the study area range from –0.73 to −5.08‰ and from

8.21 to −15.2‰, with average values of −3.31 and −5.70‰, respectively (Table 4.1).

The d-excess value in the studied water samples ranges from 11.6 to 27.6‰.

Significant evaporation from surface waters might have caused the higher d-excess

as the vapour re-condenses in the atmosphere (Clark and Fritz, 1997). The cross

plot of δ18O and δ2H values of the water samples collected in the study (Fig. 4.10a)

shows that local precipitation is the major source of recharge to the aquifers of the

area. Slight shifting of groundwater samples towards the left in ellipses A and B

(Fig. 4.10a) is mainly attributed to their location at a higher altitude, and hence,

the combined influence of the altitude and the difference in isotopic composition of

its local air mass from that of Addis Ababa (Girmay et al., 2015). As a result of the

cold and humid summer weather of the relatively elevated localities in the area, a

depleted and high d-excess air mass is expected below the cloud base (Girmay et

al., 2015). Addis Ababa is located close to the lakes region of the Ethiopian Rift

valley and, hence, the isotopic exchange of rain droplets with the relatively

enriched vapour from these continental water bodies can result in a relatively

enriched precipitation and groundwater recharge (Kebede et al., 2005). The


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samples in ellipse A are at a relatively higher altitude than those in ellipse B, as

demonstrated by more depleted samples in ellipse A.

The samples in A, B and C approximately correspond to the cluster groups I, II and

III. The distance effect can also have a slight impact, as the area is located far from

the moisture source of the summer precipitation in this region as compared to

Addis Ababa. Kebede et al. (2005) also indicated that the major source of recharge

to the Ethiopian groundwater is the summer rainfall, and the distance and altitude

effects are prominent factors in depleting the precipitation in the central highlands

of Ethiopia. The depleted signature of the samples SP07, SP09, SP10, SP13, R03,

and R07 are similar to the others in ellipse B, while they are from relatively lower

altitudes, which indicates that there is fast groundwater flow along regional open

fractures. The enrichment of some spring samples in polygon C (Fig. 4.10a)

signifies the influence of local recharge from nearby surface waters. In most cases,

the springs and surface water on the landslide areas are being supplied with

groundwater that is recharged from higher elevations above the landslide complex.

During the fieldwork, interviews with local residents in the recently affected area

indicate that numerous springs emerged at Yizaba, Shotel Amba, Nech Amba, Nib

Amba, and Wanza Beret localities (Fig. 4.3) following the landslide incidents. And

the springs in Yizaba locality are observed to change their flow directions from

time to time which can indicate that there is still active mass movement in the

area.

Many landslides have occurred within the formation that is poorly welded and

composed of tuffaceous and volcanic ash materials, which readily weather to clay

minerals and are capped by highly brecciated ignimbrite. Springs are common at

the interface between the fractured rock and its underlying weathered part or

volcanic ash or a paleosoil that occurs between various lava flows. The pyroclastic

unit contains lapilli tuff, tuff breccia and tuffaceous strata, which are susceptible

to slaking. Thus, the stable isotope results indicate that rainfall is one of the main

triggering factors of the slope instability in the area associated with degrading rock

mass strength and increase of the weight of the slope mass, i.e. increasing the pull

of gravity.


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Figure 4.10: (a) Cross plot of δ18O versus δ2H of the water samples with the Addis

Ababa LMWL and the GMWL, (b) isotopic altitude effect of precipitation of the

study area and (c) cross plot of 18O versus electrical conductivity (EC) of the study

area.

The possibility of identifying a relationship between rainfall and reactivations of

the landslide was investigated by using historic records of precipitation in the area.

The mean annual precipitation measured in Debre Sina station for the period 1974

– 2016 was 1812 mm, while the mean precipitation in the period from June to

September was 1037 mm. Figure 4.11 shows that the rainfall intensity in July and

August of the years 2005, 2006, 2007 and 2014 was considerably above average.

The most evident landslide reactivations were the movements that occurred in the


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summer season of the years 2005 – 2007, 2014 and 2016 following long-lasting and

above average precipitation (Fig. 4.11).

According to Woldearegay (2008) and Abay and Barbieri (2012), the ultimate

mobilization of the landslides in the area has occurred in the month of September

and some in October signifying the effect of the heavy rains on stability. Ayalew

(1999) also found significant landslides in the Ethiopian highlands occurring

between September and October. The water enters through the open tensional

cracks or pipes during periods of high precipitation, building up a rapid pore-water

pressure. This phenomenon, together with the effect of surface erosion around the

lower parts of the slopes and an increase in bulk density at the top, might cause

sudden and catastrophic failure.

Figure 4.11: Mean monthly rainfall of the area for the last 43 years (1974 to 2016)

and mean monthly rainfall for the years 2005, 2006, 2007, 2014 and 2016 for the

Debre Sina area.


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4.4.5.2 Spatial distribution of δ18O

The δ18O of groundwater samples and the altitude from which they are collected,

are inversely related and the regression line on the cross plot indicates a depletion

rate of −0.1‰/100 m for δ18O (Fig. 4.10b). A similar altitude/pseudo-altitude effect

in the Rift valley was also reported by Kebede and Travi (2012). At lower altitudes

clouds are usually higher above the ground level than at higher altitudes.

Therefore, the evaporative enrichment during rainfall is larger at low altitudes,

which is called a pseudo-altitude effect (Kebede and Travi, 2012; Girmay et al.,

2015). Therefore, the depletion in δ18O of the shallow groundwater in the highland

plateau of the study area (Fig. 4.10b) is due to recharge from already depleted

precipitation reaching these elevated ground areas due to the altitude effect.

However, the enrichment of shallow groundwater in the middle and lower altitudes

(Fig. 4.10b) of the area can also be attributed to recharge from already enriched

local rainfall. For river water, as it moves towards the eastern lower altitudes,

progressive enrichment is also likely due to continuous evaporation as the surface-

water flows downstream and further evaporation of the shallow groundwater that

feeds the base flow in the discharge areas in the lower parts of Dem Aytemashy,

Robi, Majete, and Shenkorge Rivers and their tributaries (Fig. 4.4). This altitude-

isotopic composition relationship can contribute to understanding the origin and

flow paths of water within a slope. The EC versus oxygen isotope (δ18O) also shows

a strong correlation (Fig. 4.10c) which can indicate the dominance of locally

recharged shallow groundwater flow system in the area.

4.4.6 Vertical Electrical Sounding

Geophysical studies were carried out at two selected sites, namely, Yizaba and

Armaniya. Vertical electrical soundings (VES) were conducted along the deep-

seated landslide in Yizaba and along the shallow to intermediate landslide in

Armaniya in order to trace the orientation and location of the faults and geological

contacts, which can have considerable effect on the groundwater circulation, as

well as to map the various aquifer systems. These investigations were also aimed

at determining the thickness of the overburden materials, to characterize the


123

vertical distributions of subsurface layers, to estimate the depth to the water table,

to identify probable aquifer beds and to characterize the nature of the bedrock. The

VES interpretation of the three profiles is based on lithological outcrops, surface

observations and the overall geological setup of the area.

4.4.6.1 Profile line–1

This profile line is about 368 m long and comprising two VES points: VES–1 and

VES–2 (Fig. 4.3 shows profile line-1, and Fig. 4.12 shows the VES points). This

profile shows a subsurface represented by six distinct major lithological units. The

upper layer with a thickness range of 0.947–2.97 m and resistivity range of 40.2–

47.4 Ω-m is interpreted as rhyolite. The second layer, having a resistivity response

range of 19.1–23.7 Ω-m and a thickness range of 2.5–14.4 m, is interpreted as

highly weathered pyroclastic sediment. The third layer, which has shown a

resistivity range of 64.6–98 Ω-m, is the highly weathered porphyritic basalt. The

fourth layer, with a resistivity of 10.8 Ω-m and 10.5 m thickness, is attributed to

the highly weathered pyroclastic sediment, which is inferred to be highly

saturated. The fifth layer, with a resistivity of 207 Ω-m and has 25.9 m thickness,

is attributed to ignimbrite.

In the locality around VES–1, the beds with relatively high resistivity rest on a

formation characterized by a low-resistivity response (1.72 Ω-m). This low

resistivity can be explained by the highly fractured and saturated nature of the

fractured basalt. Depth to the compact rock varies due to the presence of deep-

seated geological structures. The layers from fourth to sixth have vanished in VES–

2 because of the normal fault. The locality around VES–2 is observed to be

vulnerable to sliding, which is possibly favoured by increased groundwater

pressure within the fault zone. The measured resistivity values of the different

geological units coincide with the previously obtained resistivity values in the

literature (Keller and Frischknecht, 1966).


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profile line–1.


The pseudo-depth section and resistivity section for profile line-2, displayed in Fig.

4.13, allowed qualitative data to be interpreted, based on both lateral and vertical

resistivity variations in the subsurface. This line is about 300.5 m long and it

comprises three VES points: VES–3, VES–4, and VES–5 (Figs. 4.3 and 4.13). The

profile covers the landslides in the Armaniya area (Fig. 4.3). This section displays

four layers that have distinct resistivity values reflecting variation in grain size,

moisture content and weathering condition of the underlying rocks. The top thin

layer, with a resistivity range of 14.5–25 Ω-m and thickness range of 0.4–2.92 m,


125

represents dry silty clay soil. The second layer is characterized by very low

resistivity values (4.3–5.99 Ω-m) and its thickness varies from 1.5 m to 14.8 m; it

is interpreted to be saturated sandy clay/silt soil. The third layer has shown a

resistivity range of 10.37–19.67 Ω-m and 8.0–104 m thickness range. It is

associated with the highly weathered porphyritic basalt.


profile line–2.


126

The fourth layer, which has low resistivity (1.13–2.49 Ω-m), is the response of the

highly fractured and saturated nature of the highly to completely weathered

scoriaceous agglomerate basalt. Great thicknesses of disturbed and sliding soils

are located in the localities around VES–3 and VES–4. This could be due to the

presence of thick saturated soils and the high degree of weathering and fracturing

and the saturated nature of the underlying rocks.


Profile line–3 is about 260.5 m long and it comprises three VES points: VES–6,

VES–7, and VES–8 (Figs. 3 and 14). This section reveals six layers with a

resistivity variation between 6.7 Ω-m and 28.9 Ω-m. Accordingly, the top thin layer,

with resistivity values between 9.0 and 14.1 Ω-m and 0.43–4.23 m thickness range,

is associated with the upper poorly sorted colluvial deposit. The second layer,

having a resistivity range of 6.7–12.2 Ω-m and 1.75–5.48 m thickness is associated

with the highly weathered pyroclastic sediment (tuff). The third layer, with a

resistivity range of 10.1–12.6 Ω-m and 7.72–19.2 m thickness range, is attributed

to the highly weathered porphyritic basalt. The fourth layer, having a relatively

low resistivity (7.12–7.26 Ω-m) and thickness range of 12.8–18.1 m, is interpreted

as pyroclastic sediment; this low resistivity within the profile is possibly due to the

intensive degree of weathering and saturated nature of the layer. The fifth layer,

which has a resistivity response of 20.6–28.9 Ω-m, is interpreted as moderately

weathered aphanitic basalt; this layer has a thickness ranging between 28.6–34.4

m. The bottom-most layer has a resistivity of 9.05 Ω-m, which is lower than that of

the overlying layer, is associated with the water-bearing fractured basalt. The

fragile state of the bedrocks accelerates the rock mineral weathering by facilitating

water ingress into the rock mass.

Generally, all the above interpreted geophysical data indicate that the area is

covered by unconsolidated sediments and highly decomposed and weak volcanic

rocks that are susceptible to sliding when they get moist. The heterogeneity of the

geological materials and the presence of relatively impermeable layers embodied

within the highly permeable volcanic rocks can result to the build-up of high water

pressure at the interface between the contrasting permeability layers, which can


127

trigger landslides. On the other hand, the intense fracturing in the tilted basalt

and ignimbrite beds can create weak zones that accelerate the infiltration of water

which can be responsible for the build-up of high hydrostatic pressure, resulting in

lowering of the effective normal stresses in the rock mass, giving rise to landslides.


profile line–3.


128

4.5 Conclusions

The hydrogeology in the volcanic areas in the western part of the study area is very

complex as the lithology is disrupted by cross-cutting faults and interrupted by

volcanic structures. As can be seen from the chemical characterization, shallow to

intermediate aquifers cause groundwater flow into the landslide mass, influencing

long-term groundwater-level fluctuations underneath the landslide and, as a

consequence, its stability. The low TDS and bicarbonate types (Ca–Mg–HCO3 and

Ca–HCO3) of groundwater chemistry indicate a fast hydrogeological regime

receiving a relatively high amount of precipitation with infiltrated water flowing

in the fissured and disturbed aquifers developed in various volcanic rocks and

intercalated sediments. The slight rock-water interaction has shaped the

groundwater chemistry, as shown by the ionic ratios and Gibbs plots. A Piper plot

depicts that groundwater types of the study area are Ca–Mg–HCO3, Ca–HCO3,

Ca–Mg–Cl–SO4 and Na–HCO3. Groundwater shows a systematic change in

hydrochemical facies along the groundwater flow direction from the highland area

towards the lowland area. The dominantly depleted isotopic signatures in the

study area indicate that the high amount of precipitation in the cool and humid

highlands is the main source of both the groundwater and surface water in the

area. In the highland areas, the groundwater storage and flow are predominantly

in fault zones and joints, resulting in little contact between the groundwater and

the geological materials. Two groundwater flow systems (shallow/local and

intermediate-deep) are identified in the study area. The chemical and isotopic

characterization indicates that shallow to intermediate aquifers cause

groundwater flow into the landslide mass, influencing long-term groundwater-

level fluctuations underneath the landslide and, as a consequence, its stability.

The VES investigation in the Armaniya and Yizaba areas indicates that the

landslide is a deep-seated feature incorporating both bedrock and surficial

deposits. There are many springs and seepage zones along the contact between the

basalt and ignimbrite beds with the pyroclastic sediments (volcanic ash). The

heterogeneity of the geological materials and the presence of relatively

impermeable layers embodied within the highly permeable volcanic rocks can


129

result in the build-up of high water pressure at the interface between the

contrasting permeability layers, which can trigger landslides. On the other hand,

the intense fracturing in the tilted basalt and ignimbrite beds can create weak

zones that accelerate the infiltration of water, which can be responsible for the

build-up of high hydrostatic pressure resulting in lowering of the effective normal

stresses in the rock mass giving rise to landslides. Furthermore, the concave shape

of a terrain can enhance the convergence of groundwater flow into the landslide

area since groundwater levels are relatively high in such terrains. In general, the

main triggering factors for landslide problems in the area are the intensive

weathering of the rocks; the prominent geological structures; steep slope and

gradient; heavy rainfall; the groundwater pressure developed during precipitation;

and the presence of low-permeability beds which force the percolating water to flow

laterally. This study has provided a good level of understanding of the effect of the

hydrogeologic environment on landslide triggering.

Acknowledgements



supported by the Ruhr University Research School PLUS, funded by Germany’s

Excellence Initiative (DFG GSC 98/3). We highly appreciate Isodetect


Earth Sciences at the Technical University of Darmstadt for analysing the water

samples (stable isotopes δ18O and δD). We also thank the three anonymous

reviewers and the editors of Hydrogeology Journal for their constructive comments

and suggestions.


130

Chapter 5

5 Slope stability analysis of deep-seated

landslides using Limit Equilibrium and Finite

Element methods under static and seismic load

in Debre Sina area, Ethiopia

This chapter is based on Tesfay Kiros Mebrahtu, Thomas Heinze, Stefan Wohnlich,

Michael Alber (submitted). Slope stability analysis of deep-seated landslides using

Limit Equilibrium and Finite Element methods under static and seismic load in

Debre Sina area, Ethiopia.

Abstract

Slope failure is a very common phenomenon and critical issue in the western

margin of the Main Ethiopian Rift in the Debre Sina area. In order to minimize the

damage caused by failure events, a detailed investigation of landslide-prone areas

using numerical modelling plays a crucial role. The main aim of this study is to

evaluate and compare safety factors calculated by the different available numerical

methods. Stability analyses of slopes prone to different types of failures were

performed with different techniques. The stability was assessed for slopes of

complex geometry composed of aphanitic basalt, porphyritic basalt, tuff and

colluvium (poorly sorted clayey sand to silty sand) using the limit equilibrium

method and the shear strength reduction method based on finite elements.

Furthermore, numerical analysis was done under static and pseudo-static loading

using the horizontal seismic coefficient to model their stability during a seismic

event. The slope stability analysis indicates that the studied slopes are unstable,

and any small scale disturbance will further reduce the factor of safety and cause

failure. The critical strength reduction factors from the finite element method are

significantly lower than the factor of safety from the limit equilibrium method in

all studied scenarios. The slope stability of landslide prone hills in the study area

strongly depends on the saturation conditions and the seismic load.


131

5.1 Introduction

Slope failure is a common phenomenon around the world, often severely affecting

human lives. Slope failures are one of the deadliest and destructive natural

hazards, as they can cause substantial human causalities and property damage

(Keefer, 1984; Varnes, 1984; Nadim et al., 2006; Clague and Roberts, 2012). The

International Association of Engineering Geology (IAEG) database indicates

approximately 14 % of injuries and deaths from natural catastrophes are caused

by slope failures (Aleotti and Chowdhury, 1999). Slope instability has always been

a major problem along the hilly and mountainous terrains of the highlands of

Ethiopia. The Debre Sina area comprises one of the largest mountain chains with

undulating topography and a bi-modal monsoonal season. The area is one of the

largest deep-seated landslides in Ethiopia superimposed by several shallow

landslides. The fragile and highly deformed nature of rocks makes this area

vulnerable towards different types of slope failures. These failures cause a

considerable loss of life and property, lead to displacement of people and also

seriously impact agricultural land, dwellers and infrastructure. Almost 60 % of the

total population in Ethiopia lives in the highland areas (Ayalew, 1999) which is

characterized by high relief, complex geology, high rainfall, rugged morphology,

very deep valleys and gorges with active river incision.

Mechanical analysis of slope stability provides knowledge of parameters

controlling landslides, entirely removing the guesswork (Cruikshank and Johnson,

2002). In the 1990s most rock slopes were evaluated solely with stereographic

projections using kinematic analysis to identify potential geological structures or

structure sets that may induce sliding or toppling failures (Bar et al., 2019). In the

2000s, slope stability modelling techniques evolved to more complex two-

dimensional limit equilibrium analysis and numerical modelling for isotropic rock

masses (Bar et al., 2019). Various numerical tools such as the limit equilibrium

method (LEM), finite difference method (FDM), boundary element method (BEM),

and finite element method (FEM) have been used by researchers for analysis of

slope stability problems (Griffiths and Lane, 1999; Jing, 2003; Cheng et al., 2007;

Sarkar and Singh, 2008; Esteban et al., 2015; Liu et al., 2015; Stianson et al., 2015).


132

The characterization of the dynamics of complex rock masses, especially fractured

rock slope deformation and failure using numerical models, has become a major

challenge in the field of engineering geology and rock mechanics in present day

(Kundu et al., 2016). The LE methods remain popular because of their simplicity

and the reduced number of parameters they require, which are slope geometry,

topography, geology, static and dynamic loads, geotechnical parameters and

hydrogeologic conditions (Ayob et al., 2019). However, they do not take into account

the ground behavior and the safety factors are supposed to be constant along the

failure surface. In recent years, the finite element method has gained its popularity

due to its robustness in arbitrary boundary and interface condition, complex

problem solving capacity, free from presumption of critical slip surface and

elimination of assumptions regarding the inclinations and locations of interslice

forces (Hammah et al., 2004). The technique addresses complexities regarding

geometry, non-linear deformability, material heterogeneity, complex boundary

conditions, in-situ stresses and gravity in addition to several coupled processes

such as pore pressure and seismic loading (Ayob et al., 2019). The application of

FEM in slope stability analysis has become more common and is able to describe

progressive failure.

The FEM is increasingly applied to slope stability analysis. One of the most

popular techniques for performing FEM slope stability analysis is the shear

strength reduction (SSR) approach (Griffiths and Lane, 1999). The SSR

systematically reduces the shear strength envelope of a material by a factor and

computes FEM models of the slope until deformations are unacceptably large or

solutions do not converge. The factor is termed as strength reduction factor (SRF).

The SRF is said to be critical when the finite element model does not converge to a

solution or in a simple term, the system becomes unstable (Hammah et al., 2005,

2007). The FEM as a part of continuum analysis method, is based on the principle

of dividing the whole domain into finite elements, where each element shares its

nodes with neighboring elements, typically based on triangulation. This technique

allows only small displacements forbidding large dislocations and complete

detachment of elements (Kundu et al., 2016). An advantage of the FEM is that

joints can be incorporated into the model by the implementation of fracture


133

elements, which consider spacing, aperture, infilling and continuity (Jing and

Hudson, 2002). In case of blocky rock masses, the displacement is very small until

its failure. Rocks exhibit considerable variations in strength and deformability

both spatially and inherently, because they are formed as a result of various

previous geological and tectonic processes (Sari, 2019). However, the interaction

between discontinuities and the intact rock mass mostly defines the stability when

the natural equilibrium is disturbed. If the mechanical model is close to reality, a

proper stress-strain analysis can be done and it is possible to understand the most

probable failure mechanism (Elmo, 2010). For rock mass, it is necessary to use joint

patterns to have an adequate representation of realistic rock configurations. The

mechanical properties of joints are functions of the physical properties in

discontinuities that affect the mechanical behavior by friction, compressive

strength, weathering, and filling. Also, in comparison to the slice method used in

LEM, the SSR does not require a pre-defined failure surface or the search for a

minimum failure surface as the failure plane is an output of the SSR method (Sari,

2019).

A number of studies have previously compared the results of slope stability

analyses using the LE and FE methods (Griffiths and Lane, 1999; Hammah et al.,

2004; Khabbaz et al., 2012; Vinod et al., 2017; Zein et al., 2017), achieving a good

agreement between LEM and FEM for homogenous material and simple

geometries but revealing an overestimation of the slope stability using LEM for

complex geometry and heterogeneous material. However, Hammah et al. (2004)

recommended adopting the FE method using shear strength reduction factor as an

additional robust and powerful tool for design and analysis. As Sari (2019)

mentioned, it is necessary to conduct different analysis methods in rock slope

stability studies, since the use of a single method may not always produce

satisfactory results regarding the stability condition of rock slopes. Review of

previous studies in the study area (Schneider et al., 2008; Woldearegay, 2008),

revealed that the failure mechanisms of medium to large rock slides were assessed

based on the distribution of rock and soil masses and field observations of features

that indicated mass movements. In order to improve the understanding of the

triggering of such large rock slides, numerical models of slope stability with LEM


134

and an elasto-plastic FEM using the SSR technique with the Mohr-Coulomb failure

criteria were carried out in the Debre Sina area, particularly in slope sections of

Shotel Amba, Yizaba, Nib Amba and Wanza Beret. These areas were severely

affected by landslide incidences in recent years. In this research, the SSR

technique was used to determine an SRF or factor of safety (FS) value that is

associated with a slope at the verge of failure. Based on the model outcome, future

slope performances are evaluated and various combinations of loading conditions

in the natural environment that may affect the area in the future are studied. The

factor of safety is a very useful index for determining how close or far away a slope

is from failure. Stability was revised in static and seismic conditions. In addition,

comparisons were made between four different methods of calculating the FS using

the LEM.

5.2 Geology of the area

Ethiopia is located close to the active East African Rift System (EARS) which

results in numerous landslides in many parts of the country. The Ethiopian

highlands are susceptible to various types and sizes of landslides due to their

variable topography and geology. The study area is located in the central-western

highlands of Ethiopia north of Addis Ababa, which is part of the northwestern

Ethiopian plateaus. Debre Sina lies within a tectonically active region along the

rift escarpment border. The area is characterized by undulating topography and

the presence of alternating hills and valleys. The drainage pattern in the area

represents parallel to sub-parallel dendritic patterns developed along faults and

master joints in the hard rocks (Mebrahtu et al., 2021).

The exposed rocks are highly jointed aphanitic basalt-porphyritic-agglomerate,

ignimbrite-tuff-volcanic ash, porphyritic basalt-scoriaceous agglomerate,

Tarmaber basalt and upper ignimbrite of Tertiary age (Fig. 5.1) and

unconsolidated deposits (colluvial and alluvial deposits) (Mebrahtu et al., 2020a).

These rocks of the area are of volcanic type and contain geological structures such

as faults, lineaments and fractures/joints. Intense fracturing, columnar jointing

and spheroidal weathering are very common features. The ignimbrite found in the

study area is associated with vitric tuff. It is characterized by both vesicular and


135

massive variety. The rock fragments include pumice, older ignimbrite, vesicular

basalt, fine-grained glass material, and volcanic ash.


2020a).


136

In particular, the hummocky topography of the Yizaba–Shotel Amba area covered

by colluvium materials, locally including pyroclastic sediments (Mebrahtu et al.,

2020b). These deposits consist of unsorted to poorly sorted loose soil sediments

(clayey sand to silty sand) and matrix-supported rock fragments, with large blocks

of basalt toppled from upslope cliff faces (Mebrahtu et al., 2020b). The alluvial

deposits are composed of unconsolidated sediments ranging in size from fine clayey

sand, sub-rounded to rounded pebbles, cobbles and boulders. The unit is

characterized by highly variable width, thickness and composition, both along and

across the sequence. The most significant deposit occurs along the Dem

Aytemashy, Robi and Shenkorge rivers, northern, northeastern and southeastern

part of the study area where there is relatively flat topography (Fig. 5.1).

Figure 5.2: Panoramic view of the main Yizaba and Shotel Amba landslides from

east with examples of characteristic geodynamic features within the main

landslide body and its surroundings (modified from Mebrahtu et al., 2021): (a)

rotational slide, (b) rock slide, (c) debris/earth slide, (d) debris flow, (e) earth flow,

and (f) a quasi-rotational slide widening retrogressively with ponded spring water

at the toe of the Wanza Beret landslide.


137

The main geological structures identified in the present study area are faults,

lineaments and fractures/joints. The study area has been affected by several

normal faults and lineaments, which mainly strike N–S, E–W, NE–SW NNE–SSW

and NW–SE major trends of faulting and lineaments (Fig. 5.1). Various types of

slope failure have affected most parts of slopes. According to Varnes (1978)

classification, the most common types of landslides in the study area are rotational

slides, translational slides, rockfalls and toppling, debris slides, and debris and

earth flows (Figs. 5.1 and 5.2). The most prominent landslide phenomena were

observed at Shotel Amba, Yizaba, Nib Amba, and Wanza Beret areas (Fig. 5.1).

5.3 Methods and materials

The rock samples from different zones of the slope were collected for determination

of the geo-mechanical parameters of the failure criterion. The rock block samples

were collected from each category of the rocks and their respective geotechnical

properties measured experimentally. The representative rock samples were cored

in the laboratory using diamond drilling and tested to determine the

geomechanical parameters according to ISRM (1981) specifications. The laboratory

testing included index properties tests, uniaxial compressive strength (UCS),

triaxial compressive strength (TCS), and Brazilian tensile strength (BTS) in the

rock mechanics laboratory of Ruhr University of Bochum, Germany (Fig. 5.3). The

engineering properties of the different rock masses were estimated from field tests

(Schmidt Hammer Rebound) and geological strength index (GSI). RocData 5.0

software from Rocscience was used to determine the equivalent rock mass

properties using the GSI. Also, the residual values of cohesion and friction angle

were taken as 25 % and 20 % of the peak values for the slope stability analysis,

respectively. The mean value of the tested samples was taken for the numerical

solution of the slope stability analysis (Table 1). Minimum three specimens have

been tested from each unit.


138

(a) Pre-failure of

aphanitic basalt-UCS

(b) Post-failure of

aphanitic basalt

(c) Pre-failure of

aphanitic basalt-BTS

(d) Post-failure of

aphanitic basalt

(e) Pre-failure of

porphyritic basalt-TCS

(f) Post-failure of

porphyritic basalt

(g) Pre-failure of

porphyritic basalt-

BTS

(h) Post-failure of

porphyritic basalt

(i) Pre-failure of tuff-

UCS

(j) Post-failure of

tuff

(k) Pre-failure of tuff (l) Post-failure of

tuff

Figure 5.3: Specimens prepared and tested under uniaxial, triaxial, and tensile

loading.

5.3.1 Model generation

After all the necessary data has been collected and potential slope failure locations

have been identified, slope profiles were defined from geological data and

topographic maps. The potential landslide locations selected in Figs. 5.4–5.7 were

taken as a domain for the present stability analysis. The model development for

the landslide slope sections includes: selecting and defining the problem geometry,

assigning the appropriate material model and properties, followed by applying

boundary conditions. The LE methods are based on force and moment equilibrium,


139

while the FE methods use the stress-strain relationships to determine the behavior

of the model. For this study, slope analyses were performed using Rocscience

SLIDE2 and RS2 for LE and FE methods, respectively. The geometry was

implemented for the different slope sections and each rock strata in every slope

section was assigned with the corresponding properties for the rock mass and joints

accordingly in the RS2 model as given in Table 5.1. Then the boundary conditions

were assigned to the slope model (Figs. 5.4a–5.7a). The boundary at the base of the

FE model was fixed for displacement in the x and y-directions, while the vertical

side boundaries were fixed for displacement in the x-direction (Figs. 5.4b–5.7b).

The Mohr-Coulomb criteria is used to define the intact rock and joint strength

characteristics for an elastic-perfectly plastic behavior.

The rock slopes are discretized into deformable six-nodded graded triangular finite

elements with increased density near the faults surface as shown in Figs. 5.4b–

5.7b. Gravitational stress field with horizontal to vertical in-situ stress ratio of

unity was adopted. According to Eberhardt et al. (2003), stress was initialized

assuming horizontal to vertical stress ratio of 0.5. The mesh was made up from

approximately 434,318 nodes and 216,659 elements for Shotel Amba, 435,318

nodes and 217,160 elements for Yizaba, 138,547 nodes and 68,774 elements for Nib

Amba and 138,274 nodes and 66,485 elements for Wanza Beret slope sections (Figs.

5.4b–5.7b), though the results have been shown to be independent of the mesh

density. The hydraulic condition was determined based on the spring locations and

assumed to be in a steady state. Groundwater is a crucial factor in landslide

initiation. Groundwater level increase is often the critical factor for slope failure

because it induces high pore-water pressure which reduces the frictional strength

of slopes. An increase of pore-water pressures due to the flow of groundwater is an

important factor in the development of slope failures and the occurrence of

landslides. In particular, the presence of groundwater under pressure often

facilitates severe slides of the flow type.

5.3.2 Limit equilibrium analysis

The slope stability analysis is performed with the limit equilibrium methods based

on assumptions about the sliding surface shape. Limit equilibrium (LE) methods


140

are used extensively for slope stability analysis and use the Mohr-Coulomb failure

criterion to determine the shear strength along a slip surface (Deng et al., 2015).

A state of limit equilibrium exists when the mobilized shear strength is expressed

as a fraction of the shear stress. The slope is primarily considered to fail along an

assumed slip surface. At failure, the shear strength is fully mobilized along the

critical slip surface. The shear stress at which a soil fails in shear is defined as the

shear strength of the soil. In saturated soils, the Mohr-Coulomb shear strength for

an effective stress analysis is usually expressed in a linear form as follows Eq. (5.1).

τ = c′ + σ′tan′ (5.1)

Fs =S

τ=

c′+σ′tan′

τ (5.2)

where, = shear stress (kPa), c'= effective cohesion (kPa), '= effective normal stress

on the surface of rupture, (kPa), '= the effective angle of internal friction (°) and

Fs= factor of safety.

The FS for a slope failure is calculated as the ratio of the available shear strength

to the mobilized shear strength (Eq. (5.2)). The FS is a very common method for

evaluating the stability of slopes. In theory, a FS of 1 means the driving and

resisting forces are at equilibrium. Limit equilibrium methods are relatively simple

in their application, compared to numerical methods (Eberhardt et al., 2003; Tang

et al., 2017). SLIDE2 (Rocscience Inc. 2018) is used for the stability of slip surfaces

using vertical slices. In this method, various forces responsible for driving the rock

mass and resisting forces are evaluated for discrete slices along the profile

constrained by the surface and the failure plane. The ratio of resisting forces to

driving forces at equilibrium defines the FoS. LE methods have been broadly

applied since they produce satisfactory FoS results that can be corroborated from

its basic ideologies. However, the LE methods are rather basic in their form as they

do not fully consider the stress-strain relationship of the soil, which is also

essential for slope stability evaluation. Limit equilibrium analysis is used to give

an estimate of the FS and does not manifest information regarding the

deformations associated with failure.


141

A number of LE methods have been developed for solving the force and moment

equilibrium equations of the sliding body in circular and non-circular failure

surfaces. The circular and non-circular limit equilibrium methods considers the

equilibrium of the total failing mass only and therefore the internal equilibrium of

the sliding mass is not considered. In this study, four well-known techniques,

namely Bishop’s simplified method (Bishop, 1955), Janbu’s simplified method

(Janbu, 1968), Spencer’s method (Spencer,1967) and Morgenstern-Price method

(Morgenstern and Price, 1965) were employed for locating critical slip surfaces in

heterogeneous rock mass conditions using the grid search and auto refine search

tools provided by the program. Each point in the slip center grid represents the

center for rotation of a series of slip circles. These methods are commonly used due

to relatively adequate accuracy while calculating the FS and for establishing a

common platform for conducting the comparative study between LE and FE

methods.

5.3.3 Finite element analysis

Finite element analysis was performed using RS2 v.9.0 software (Rocscience Inc.,

2020). The application of FEM can overcome limitations in LEM, because instead

of just the FS, the maximum shear strain, total displacement, and yield elements

of the slope can be evaluated. The SSR method is commonly used in FE slope

stability analysis to calculate the critical SRF. In this approach, the strength

parameters are incrementally reduced by a certain factor (SRF) until failure

occurs. This approach is best explained for slope materials characterized by Mohr-

Coulomb strength parameters (Hammah et al., 2007). The SSR technique involves

reducing the Mohr-Coulomb strength parameters cohesion (c) and angle of friction

() by the SRF until non-convergence occurs within a specified number of iterations

and tolerance (Zhou et al., 1994; Griffiths and Lane, 1999; Gover and Hammah,

2013). Non-convergence occurs when there is unsolved force and displacement

induced at a node of a finite element model (Kainthola et al., 2012). The reduction

factor that causes the FE model not to converge is called critical reduction factor.

The critical SRF value that brings the slope to failure is taken as the slope’s factor

of safety. The SRF that corresponds to the last convergence state is equivalent to


142

the safety factor. The SRF parameters are given as follows (Eq. (5.3) and Eq. (5.4)).

The FS is obtained by dividing the base strength by the lowest strength at which

the slope is stable.

𝐶𝑟 =𝐶

SRF (5.3)

tan(𝑟) =

tan

SRF (5.4)

where, c is the cohesion, is the angle of internal friction, Cr and r are the reduced

shear strength parameters and SRF is the shear strength reduction factor.

Table 5.1: Material parameters for rock used in LE and FE models.

Material parameters Colluvium

Porphyritic

basalt Tuff

Aphanitic

basalt

Value Value Value Value

Unit weight (MN/m3) 0.021 0.0225 0.0199 0.0275

Peak cohesion (MPa) 0.107 13.54 5.30 38.17

Peak friction angle (°) 35 56.68 53.02 62.23

Peak tensile strength (MPa) 0.0028 8.10 3.55 18.87

Residual cohesion (MPa) 0.08025 10.155 3.975 28.6275

Residual friction angle (°) 28 45.344 42.416 49.784

Residual tensile strength (MPa) 0 0 0 0

Young's modulus (MPa) 20 10400 612 64121

Posisson's ratio 0.25 0.11 0.16 0.17

Numerical techniques are also useful to see the effect of the variation of the input

parameters on the overall response of the rock structures. Geology, discontinuities,

material properties (e.g., normal stiffness, shear stiffness, shear strength, and

deformability), constitutive equations, failure criterion, groundwater pressure,

external loads, in-situ stresses are taken as input parameters based on the

requirements of the methodology deployed. The rock structure and patterns of the

joints are represented by using the Mohr-Coulomb constitutive model. The

discontinuities present in the rock mass play a significant role in controlling the

strength and deformational characteristics. The behavior of the discontinuities is

usually defined in the form of normal and/or shear stiffness. Joint stiffness

parameters describe the stress-deformation characteristics of the joint and are

fundamental properties in the numerical modelling of jointed rock. Among others

Barton (1972) suggested the following Eq. (5.5) for the estimation of the peak


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normal stiffness (MPa/m). The faults and interfaces also follow Mohr-Coulomb

failure criterion in order to evaluate the possibility of slipping failure along the

faults (Table 5.2). The normal stiffness of joints (Kn) can be estimated from rock

mass modulus, intact rock modulus and joint spacing (Eq. (5.5)), whereas the shear

stiffness of joints (Ks) were taken as Eq. (5.6). The rock mass modulus for normal

stiffness estimated using the Hoek-Brown criterion and the GSI for different rock

types were determined.

𝐾n =𝐸i𝐸m

L(𝐸i−𝐸m) (5.5)

where, Em= rock mass modulus, Ei= intact rock modulus, Kn= joint normal

stiffness, L= mean joint spacing, Ks= joint shear stiffness

𝐾s = 0.1 ∗ 𝐾n (5.6)

To obtain realistic results from the FE analysis, some researchers strongly

suggested to include the effect of discontinuous media in the analysis (e.g., Styles

et al., 2011; Agliardi et al., 2013; Satici and Unver, 2015). In addition, some studies

showed that defining the joints in the FE models could produce an altered failure

surface (Hammah et al., 2008; Fu and Liao, 2010). In low stress environments such

as slopes, discontinuities exert a greater influence on the rock mass behavior than

intact rock properties do. Besides, the occurrence of discontinuities changes the

stress distribution around the rock mass. The presence of discontinuities has

significant impact on the mechanical behavior of the rock mass, as the joint

intersections are often areas of high stress as well as deformation counters, damage

and failure (Barton and Choubey, 1977).

Table 5.2: Geomechanical parameters used for faults.

Fault parameters Shotel Amba – Yizaba Nib Amba

Fault 1 Fault 2 Fault 1

Normal stiffness, Kn (MPa/m) 6180 6260 276.29

Shear stiffness, Ks (MPa/m) 618 626 27.63

Peak friction angle (°) 38 38 38

Peak cohesion (MPa) 0 0 0

Peak tensile strength (MPa) 0 0 0

Residual friction angle (°) 28 28 28


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along the Shotel Amba section.


along the Yizaba section.


145


along the Nib Amba section.


along the Wanza Beret section.

5.4 Results and discussion

5.4.1 Limit equilibrium analysis

The model outputs of the SLIDE2 2018 program for the Mohr-Coulomb material

types using grid search methods are shown in Figs. 5.8–5.10. The calculated FS


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using the LE modeling are presented in Table 5.3. Between the four LE methods,

MPM resulted in the highest and JSM resulted in lowest FS but the differences

are marginal.

Table 5.3: Calculated FS using LE and FE methods without horizontal seismic

coefficient (h= 0).

Slope section

LEM Min. FS FEM

BSM JSM SM MPM Critical

SRF

Max. total displacement

(m)

Shotel Amba 2.26 2.16 2.27 2.27 1.51 23.2

Yizaba 2.59 2.52 2.59 2.59 1.87 15.6

Nib Amba 11.31 10.46 11.31 11.17 1.95 0.52

Wanza Beret 2.67 2.53 2.67 2.67 2.01 0.84

LEM= Limit Equilibrium Method; BSM= Bishop’s Simplified Method; JSM=

Janbu’s Simplified Method; SM= Spencer’s Method; MPM= Morgenstern-Price

Method; FEM= Finite Element Method

The calculated FS without seismic load ranged between a minimum of 2.16 and a

maximum of 11.31 (Table 5.3) and between a minimum of 0.92 to a maximum of

4.76 with seismic load for the various cases evaluated in this study (Table 5.4).

This clearly confirms that there is possibility of a sliding circular (rotational) slip

failure in the studied slopes. From the Shotel Amba to Wanza Beret cross-sections

(Figs. 5.4–5.7), the Shotel Amba cross-section attained the lowest FS of 0.92 with

seismic load and 2.16 without seismic load in the 2D limit equilibrium analysis as

shown in Fig. 5.8. It can be seen from Table 5.4 that the conventional method for

calculating the FS with seismic load is the JSM, which has a value of 0.92 in Shotel

Amba slope section. The highest FS is produced by MPM with a FS value of 1.02.

While BSM and Spencer methods produce almost similar FS values and failure

surfaces than the MPM, the JSM results in lower FS values and different failure

surfaces. The fact that JSM generates lower FS values is compulsory, since this

method is simpler than the BSM and Spencer methods, which are both more

rigorous methods. This leads to some differences in the FS values, though the

rigorous methods are often considered more reliable (Duncan and Wright, 1980).

From Table 5.3, it is observed that Nib Amba was the only slope section with very

high FS using the four LE methods. Among the slope sections, the Nib Amba slope


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section differs very high from the other sections. The likely reason for that is when

computing the FS, the LEM didn’t consider joints act as initiation points of the

failure plane. In the slope section of Nib Amba there are weakness zones such as

faults and joints distributed along the lithologic boundary between the porphyritic

basalt and tuff.

Figure 5.8: 2D cross-section result from slide along the Shotel Amba section.

Figure 5.9: 2D cross-section result from slide along the Yizaba section.


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Figure 5.10: 2D cross-section result from slide along the Wanza Beret section.

The FS values computed without seismic load via BSM, JSM, SM, and MPM are

all independently stable, while the FS values computed with seismic load indicate

slopes close to failure. This shows that the stability of the area becomes unstable

when there is a combination of saturation and seismic load. The upper part of the

slopes is dominated by colluvial deposits or intensively fractured rocks. In this

case, circular failure surfaces are expected.

5.4.2 Finite element analysis

The landslide stability of the selected slope sections was analyzed for shear strain,

slope stability and total displacement using the FE method (Figs. 5.11–5.14). The

RS2 software was utilized for constructing the slopes in Shotel Amba, Yizaba, Nib

Amba and Wanza Beret (Rocscience Inc. 2020). The models were divided into two

cases: (i) case 1 is the initial state in a static condition before the earthquake and

(ii) case 2 is the dynamic slope model incorporating a seismic event to simulate the

influence of an earthquake event on the selected slope sections (Figs. 5.11–5.14).

Based on the seismicity and the knowledge of the geology and tectonics, the region

can be broadly divided into three seismic sources (Mammo, 2005) namely the Afar

depression, the escarpment and the Ethiopian Rift system which are very near to

the study area. The study area is situated along the western margin of the Main

Ethiopia Rift (MER) which is tectonically active. The earthquake coefficient of

Peak Ground Acceleration (PGA) values for the Afar area ranging from 0.16g


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(EBCS, 1995) to 0.75g (RADIUS, 1999) for the 0.01 annual probability. The inertia

forces due to earthquake shaking are represented by a constant horizontal force

(equal to the weight of the potential sliding mass multiplied by a coefficient) and

are commonly referred to as pseudo-static analysis (Pyke, 2002). In this work, the

horizontal earthquake coefficient of h= 0.3 is adopted for the pseudo-static slope

stability analysis as an average value for Shotel Amba and Wanza Beret and h=

0.2 for Yizaba and Nib Amba slope sections, as for the slope sections of Yizaba and

Nib Amba with the horizontal earthquake coefficient of h= 0.3 the FE model does

not converge. A slope is considered unstable in the SSR technique when its FE

model does not converge to a solution (within a specified tolerance). This shows

that the slopes are unstable as the seismic load increases.


analysis for total displacement with maximum total displacement of 225 m along

the Shotel Amba section.


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The slope sections have been analyzed under pseudo-static loading conditions

along with the gravitational forces using FEM-based RS2 software to model their

stability under a seismic event (Figs. 5.4–5.7 and Figs. 5.11–5.14). The RS2

program, on the other hand, accurately captured the complex failure pathway that

partly lay along the fault surfaces and partly pass through the intact rock (Figs.

5.11–5.14).

The slope failures in the rock masses in the FEM are more complex than in the

LEM and mainly controlled by joints and develop across surfaces formed by one or

more joint planes (Wyllie and Mah, 2004). This is in agreement with previous

studies (e.g., Styles et al., 2011; Agliardi et al., 2013; Satici and Unver, 2015).

Results for the dynamic case are presented in Figs. 5.11–5.14, the total

displacement and maximum shear strain contours exhibit a potential shear failure

surface and the critical SRF or FS for the cases evaluated in this study range

between a minimum of 0.58 and a maximum of 1.03 (Table 5.4). This indicates that

the slope stability is low due to the weak rock mass. On the other hand, for the

results from the static case the critical SRF ranged between a minimum of 1.51

and a maximum of 2.01 (Table 5.3). The slope at Shotel Amba showed that the SRF

of the slope without seismic load is 1.51, whereas the SRF of slope with seismic

load is 0.58 (Fig. 5.11). The results from this simulation are shown in Figs. 5.11–

5.14 which depicts the total displacement and maximum shear strain contours. In

comparing the resulting slope failure surfaces using the LE and FE analysis, a

large difference in the shape of the failure surface was noted. However, the slope

failure surfaces of the Wanza Beret slope section resulting from the LE methods

are best-matched to the critical failure surfaces resulting from the FE method

(Figs. 5.10 and 5.14). The critical failure surfaces resulting from the LE analyses

were perfectly circular due to the selected search criterion (Figs. 5.8–5.10), the FE

method produces a near-circular zone of failure surfaces near the toe of the slope

(Figs. 5.12 and 5.14). The FE analyses produce a better-defined failure path than

the LE analyses. The FEM be able to automatically locate the failure regions,

thereby not requiring the prior assumptions on the failure surface as compared to

the LEM.


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Table 5.4: Calculated FS using LE and FE with horizontal seismic coefficient (h=

0.2 and 0.3).

Slope

section

Seismic

load LEM Min. FS FEM

h BSM JSM SM MPM Critical

SRF

Max. total displacement

(m)

Shotel Amba 0.3 0.98 0.92 1.01 1.02 0.58 225

Yizaba 0.2 1.06 1.02 1.07 1.07 0.89 36.1

Nib Amba 0.2 4.76 4.39 4.75 4.62 1.03 1.3

Wanza Beret 0.3 1.23 1.15 1.25 1.24 0.84 1.9

LEM= Limit Equilibrium Method; BSM= Bishop’s Simplified Method; JSM=

Janbu’s Simplified Method; SM= Spencer’s Method; MPM= Morgenstern-Price

Method; FEM= Finite Element Method

Further, using the FE method, it is possible to compute the total displacement of

rock and soil from the input data. Figures 5.11–5.14 show the critical failure

surface and shear strain developed in the slope at the time of failure. The areas of

maximum shear strain indicate the likely failure pathway that would develop

through the modeled rock mass. The maximum shear strain along the critical

failure surface is found to be 18.6 in Shotel Amba and 2.35 in Yizaba slope sections.

The maximum total displacement in Shotel Amba and Yizaba are found to be

225 m (Fig. 5.11) and 36.1 m (Fig. 5.12), respectively. This is in agreement with the

findings of (Kropáček et al., 2015), who found that topographic profiles show that

the maximum estimated thickness of the active main landslide body is about 150–

200 m. The maximum total displacement in the Shotel Amba and Yizaba slope

sections coincide with the results of the topographic profiles (Figs. 11 and 12). The

maximum stress concentration along with the critical FS presumes the possibility

of a middle zone collapse and subsequent failure of the slope.


152



the Yizaba section.

In the slope section of Yizaba, shear elements are noticeable along the circular

failure surface originating from the tension zone and continuing till the toe of the

slope (Fig. 5.12). For the FS, the largest displacements are in red and the minor

ones in blue. It can be deduced that the mechanism of failure is a block sliding with

a displacement produced by the low strength of colluvial deposits (poorly sorted

clayey sand to silty sand) and tuff. In the slopes Shotel Amba and Yizaba the

middle zone of the slope moves to the right and a crack is formed in the middle of

the slope, showing that deep-failure mechanism (Figs. 5.11 and 5.12). It is also

observed that the failure surface and the location estimated by the numerical

method match with the field observation (Fig. 5.2f). There are many tension cracks

which are developed in the area, and these tension cracks are indications for

probable some more slides to occur in the near future (Fig. 5.2). A factor of safety


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of 1.03 has been attained through the finite element analysis of the Nib Amba slope

with seismic load (Fig. 5.13).



the Nib Amba section.

The computed FS with seismic load for the cases evaluated in this study is less

than one except for Nib Amba (Fig. 5.13) but which is very close to one. So, there

is a risk of failure of all slopes during an earthquake. The calculated FS of Nib

Amba slope section using FEM becomes less than one when the seismic load

increases. The findings indicated that differences between the LE and FE methods

are evident when a heterogeneous and complex slope geometry is analyzed. The

calculation of the FS shows a significant difference between the two approaches.

The FS from LEM with seismic load ranged between 0.92 and 4.76, while critical

SRF from finite element analysis were between 0.58 and 1.03 for 2D FEM

simulations. Among the slope sections, the Nib Amba slope section differs very


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strongly from the other sections. From this perspective, the case study shows that

tectonic activity plays a well-defined role in promoting landslides with seismicity

as a possible predisposing factor or even a trigger, as the area is in close proximity

to the most seismically active regions in the country (Mebrahtu et al., 2020b). The

critical SRF from FEM are significantly lower than the FS from LEM. The low

value of SRF emphatically suggests that the rock slope can be vulnerable to failure

under influence of any triggering force.



the Wanza Beret section.

The results of the FEM showed that from the static and dynamic analysis, it has

been concluded that the slope sections have the lowest FS compared to the

calculated FS using the LEM. At the critical SRF, the slope material undergoes

significant strength degradation as the failure surface is formed by coalescing

fractures, which converge with the kinematically feasible release plane. Any


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further increment in SRF value accelerates the slide mass movement which causes

an increased total displacement of slope mass material.

The lithologic units of the study area are anisotropic in their behavior and their

stress-strain behavior is quite variable, due to the presence of volcanic ash. The

elasto-plasticity behavior of the ash due to the presence of water poses a serious

threat during the periods of prolonged and heavy rainy seasons. The SRF due to a

rise in pore-water pressure might lead to slope failure. The slope stability analysis

using the LE and FE methods indicate that the area is highly susceptible to sliding

when they get moist. The springs observed at the base of the slope provide the pore

pressure regime, which determines critical conditions for the study area slope

stability. This results in the build-up of high hydrostatic pressure and results in a

reduction of effective normal stresses in the rock mass giving rise to landslides.

Moreover, the concave shape of a slope can enhance the convergence of

groundwater flow into the landslide area since water tables are relatively high in

such slopes (Mebrahtu et al., 2021).

5.5 Conclusions

In this paper, a study on both LE and FE methods for slope stability is performed.

The results of the stability analysis show that the slope stability of landslides in

the study area strongly depends on the saturation conditions and seismic load.

While the saturation conditions were assumed constant in the shown simulations,

it is evident that a pore pressure increase would further reduce the FS. The

importance of faults in hydraulic gradient of the groundwater and their influence

on the slope stability in the study area has previously been discussed extensively

(Mebrahtu et al., 2020a; Mebrahtu et al., 2020b; Mebrahtu et al., 2021). The

strength reduction technique within a framework of elasto-plasticity allows to

magnify the deformations and failure mechanisms to emerge in a natural way. The

FS obtained using LE and FE methods without seismic load ranged between a

minimum of 2.16 and a maximum of 11.31 and between a minimum of 1.51 to a

maximum of 2.01, respectively. Whereas the calculated FS using LE and FE

methods with seismic load ranged between a minimum of 0.92 and a maximum of

4.76 and between a minimum of 0.58 to a maximum of 1.03, respectively. In


156

particular, the slope sections after the earthquake event were considered as

unstable with a SRF between 0.58 and 1.03.

It was found that the FS results from LE and FE methods are significantly

different. Since the LE and FE methods employ completely different numerical

techniques, they did not produce very similar FS values and failure surfaces for

the studied slopes. According to the results of the LE analyzes, the slopes are

classified as stable, but as a result of the FE analysis, the slopes are unstable. From

the FE analysis, it can be inferred that the study area is critically unstable and

any small scale disturbance will further reduce the FS and cause failure,

particularly when the area experience heavy rainfalls or earthquakes. The

presence of multiple sets of joints and faults in rock masses and intense rainfall

further accelerate the slope failure. The result of the stability analysis shows that

the slope stability of landslides in the study area strongly depends on the

saturation conditions and seismic load. Hence, the major landslide that occurred

in September 2005 is most probably triggered by such combination of saturation

and earthquake activities.

From the stability analysis results, both the LE and FE methods produce deep-

seated failure surfaces through the foundation of colluvial deposits. It is also

observed that the failure surface and location estimated by the numerical method

match with the field observation. The strength reduction method is well-suited for

complex geometry and further representative normal stress distributions,

subsequent generating consequential FS values. The FEM generates a FS values

less than 1 for almost all the slope sections, although those are over 1 for the LEM.

In general, the FS computed from the stress-strain behavior of the rock and soil

using FE method is more realistic and provides even more accurate FS results

when a heterogeneous and complex slope geometry is analyzed. From the finding,

it is concluded that the studied slope stability evaluation methods should be

obtained collectively as part of a larger slope stability analysis to determine the

resultant FS. Furthermore, this study states that differences between the FE and

LE methods are evident when a heterogeneous slope is analyzed. Especially when

faults are involved and act as initiation points of the failure plane, the LEM tends


157

to overestimate the slope stability significantly. Both methods show a significant

decrease in the obtained FS values for the study area when the seismic load is

considered but LEM always results in higher FS values. Though, for the static

assessment without a seismic effect, the interpretation of LEM and FEM might not

vary so much, as both methods indicate the slopes as stable with FS values around

two. Therefore, it is a useful practice to employ different tools to reach a conclusion

about the stability state of slopes presenting with highly variable rock mass

conditions. In general, field investigations indicate that several different failure

mechanisms are superimposed on the deep-seated Debre Sina landslide. The

laboratory tests reveal that the lowest value of peak strength is from less

compacted tuff and prone to sliding. The tuff layers with low peak strength are

initiation points for the sliding surfaces. The FEM is found more applicable for

stability assessment because of the complex geometry, heterogeneous material and

the failure-dominating faults in the study area. The studied slopes are initially

close to failure and increased pore pressure or seismic load are very likely triggers.

Authors’ contributions

TKM as a first author carried out the fieldwork, collected rock samples, did the

laboratory analysis and interpretation of the data, conducted the numerical

simulations and wrote the manuscript while taking comments from TH, SW and

MA, and finalized the manuscript. TH and MA have been involved in a detailed

review of the manuscript prior to submission. All authors gave their approval of

the final manuscript to be published.

Acknowledgments




Excellence Initiative (DFG GSC 98/3).

Summary and future research perspectives

158

Chapter 6

6 Summary and future research perspectives

6.1 Summary

Landslides are a common phenomenon in the central highlands and Rift

escarpments of Ethiopia, which brought a heavy impact on agricultural land,

dwellers and infrastructure, and often lead to the displacement and death of

people. The Debre Sina area is one of the most tectonically active areas located

along the western margin of the Afar depression, which is frequently affected by

landslides. Despite that, urban and rural development is currently active in almost

all constricted valleys as well as on the imposing cliffs without prior site

investigation and thereby exposing these areas to landslide problems. In this

chapter the findings obtained in the four chapters (chapters 2 to 5) and their

implications are discussed and summarized. The main results and conclusions

related to the overall controlling parameters of deep-seated landslides, the

processes leading to the triggering of a landslide, and the failure mechanisms in

the study area are summarized below.

Geologically, the study area is represented by aphanitic basalt-porphyritic-

agglomerate, ignimbrite-tuff-volcanic ash, porphyritic basalt-scoriaceous

agglomerate, Tarmaber basalt, upper ignimbrite, colluvial and alluvial deposits.

The presence of highly fractured porphyritic-agglomeratic basalt, highly shattered

ignimbrite and volcanic ash, which are all prone to water absorption and

susceptible to slaking, was identified as one of the reasons for a high concentration

of landslides and main triggering factors of reactivation in the observed cases. The

results obtained during this study show that the inherent variation in the physical

property of the lithologic sequence and their structures influence slope stability.

The study area experiences high tectonic activity with intense fracturing due to its

location at the western margin of the Main Ethiopian Rift. The area has been

affected by N–S, E–W, NNE–SSW, NE–SW, NW–SE, NNW–SSE, and WSW–ENE


159

major trends of faults. Among these, the N–S and E–W trends are widespread in

the area with the highest frequency. The N–S trending normal faults are arranged

in a stepwise system towards east, which controls the morphology of the slope of

the Rift margin escarpment. The landslides displacement is orthogonal to the

NNE–SSW, and N–S striking normal fault systems that are affected by NW and

NE striking trans-tensional components. The interaction of these fault systems

produced a complex displacement across and along the escarpment, manifesting

oblique continental rifting. The intense fracturing and presence of faults favor an

easy movement along existing fault planes during saturation of the rocks or soils

and during seismic events or a combination of both conditions. The kinematic

evaluation of the faults data and slope faces using Dips v 7.0 software revealed

that the major landslide that took place on 13 September 2005 in the Yizaba Wein

locality was mainly controlled by NNE–SSW, NNW–SSE, and WSW–ENE

trending faults. The kinematic analysis shows that the rock slope has a higher

probability of failure in the wedge sliding failure mode (66.67%) compared to

planar sliding (33.33%). The faults in the area are not only weak zones, but also

mostly characterized by deeper weathering and higher potential for concentrated

groundwater flow, which can act as a lubricant and produce water pressures

causing landslides.

The geomorphological survey shows that various landforms have been identified in

the area, including long term tectonic, deep-seated slope failures, deep fluvial

dissection and slope forms. The morphology of the area is especially conditioned by

numerous rotational slides, translational slides, rockfalls and toppling, rock slides,

debris slides, and earth flows. It is observed that the morphology in combination

with the tectonic assemblage and the intense weathering processes strongly favors

the mass movement. The geomorphological analysis also showed that the complex

landslide sloped surfaces in the detachment zone are associated with a hummocky

and step-like morphology as a result of successive or retrogressive sliding. Overall

assessment of the morphometric analysis revealed that the slopes ranging from 10°

to 40°, with an elevation of 1800–2500 m and aspect to east and southeast, are

highly prone to sliding.


160

The depth of investigations from the presented seismic refraction survey was

attained with a maximum of 75 m. The seismic refraction study revealed zones of

overburden material from top to bottom, consisting of: clay, loosely cemented

colluvial sediments and highly weathered material (Vp 1000 m/s) with a

thickness of 7–15 m, highly weathered agglomeratic basalt (1000–1500 m/s) up to

30 m thick, highly to moderately fractured porphyritic basalt, ignimbrite,

rhyolite/trachyte and volcanic ash (1500–2500m/s) with a maximum 30 m thick,

moderately to slightly fractured ignimbrite, rhyolite/trachyte and basalt (2500–

3500m/s) 40–50 m thick and very strong, massive, fresh rock/ bed rock (Vp 3500

m/s). The depth to very fresh sound bedrock ranges between 45 m and 75 m. These

units are highly susceptible to sliding when it gets moist, because the volcanic

ashes are prone to slaking and acts as lubricant material. The seismic refraction

data shows that the currently active landslide is superimposed by a larger slide

including parts of the bedrock. The slip surface generally coincides with the 2000

m/s isoline due to the presence of highly fractured and saturated nature of the

underlying rocks. The upper layers of unconsolidated deposits and porphyritic-

agglomeratic basalt rocks experiences significant water transit towards the deeper

layer of ignimbrite-volcanic ash. However, the pyroclastic sediments are impeding

the vertical percolation of rainwater due to their low permeability and hence force

the rainwater to flow laterally. The study shows that tectonic activity plays a well-

defined role in promoting landslides with seismicity as a possible predisposing

factor and by determining the lines of weakness along which the landslides may

have developed. In general, the geophysical data indicates that the area is covered

by unconsolidated sediments and highly decomposed and weak volcanic rocks

which are susceptible to sliding when they get moist.

The hydrogeological conditions of the terrains are generally favourable for the

development of seepage forces within the pyroclastic sediments (tuff and pumice

horizons) and unconsolidated deposits during periods of rainfall. The residence

time is short, the soil/rock-water interaction is low and the water is barely

mineralized in the highland and intermediate regions. Therefore, it is possible to

conclude that the main cause of the landslide is not because of active soil/rock-

water interaction. It is rather because of the steep slope topography and the


161

pressure formed during precipitation, which leads to an increase in the weight of

the loose and weathered materials (increasing its shear stress) and loses its shear

resistance which finally results in mass failure or landslide.

Three groundwater groups have been identified from the preliminary HCA based

on major-ion chemistry (Na+, K+, Mg2+, Ca2+, HCO3–, SO4

2–, F–, Cl–) of the water

samples collected in this study. Group I samples are collected from basaltic and

scoriaceous aquifers in the highlands bounding the rift valley areas are

characterized by low EC (72–222 µS/cm) and low concentrations of all the major

ions. This indicates that groundwater in the highland areas is getting recharge

from rainwater. Group II samples with a similar lower EC (115–465 µS/cm) and

low concentrations of all the major ions are collected from highly fractured and

shattered ignimbrite, rhyolite, trachyte associated with basalt close to the

escarpments. However, they have relatively higher concentrations of Na+, K+, Cl–

and SO42– as compared to group I, which is mainly related to the solution or

interaction between water and secondary minerals or clay that precipitate into

faults. Group III samples are collected from the lower altitude areas (below 1500

m asl) in the eastern and northeastern parts of the study area which is extensively

covered with volcanic ash-dominated units and sporadic colluvial-alluvial deposits.

In these litho-units, the groundwater movement is slow, which together with the

presence of soluble minerals, enhances the effects of rock-water interaction giving

rise to relatively higher concentrations of Na+, K+, Cl– and SO42–. The EC values of

the groundwater samples within this group is between 215 and 573 µS/cm and

increase towards the Shewa Robit valley. It indicates that there is intermediate to

deep groundwater circulation and relatively higher residence time of the

groundwater.

The lithostratigraphic, geomorphologic, isotopic and hydrochemical evidences have

indicated that two groundwater flow systems (shallow/local and intermediate-

deep) exist in the study area. The shallow groundwater flow is mainly localized to

the highland areas and adjacent escarpments and its water table is a subdued

replica of the surface topography. The intermediate-deep groundwater flow is

strongly influenced by the lithostratigraphy and the major faults in the area rather


162

than the surface geomorphology. There are four major groundwater types in

general hydrochemical facies of the Debre Sina area identified as Ca–Mg–HCO3,

Ca–HCO3, Ca–Mg–Cl–SO4 and Na–HCO3. The low TDS and bicarbonate types (Ca–

Mg–HCO3 and Ca–HCO3) of groundwater chemistry indicate a fast hydrogeological

regime receiving a relatively high amount of precipitation with infiltrated water

flowing in the fissured and disturbed aquifers developed in various volcanic rocks

and intercalated sediments. The chemical and isotopic characterization indicate

that shallow to intermediate aquifers cause groundwater flow into the landslide

mass, influencing long-term groundwater level fluctuations underneath the

landslide and, as a consequence, its stability.

The cross plot of δ18O and δ2H values of the water samples shows that local

precipitation is the major source of recharge to the aquifers of the area. The stable

isotope results indicate that rainfall is one of the main triggering factors of the

slope instability in the area associated with degrading rock mass strength and

increase of the weight of the slope mass, i.e. increasing the pull of gravity. The EC

versus oxygen isotope (δ18O) also shows a strong correlation which can indicate the

dominance of locally recharged shallow groundwater flow system in the area. The

hydrogeological conditions of the terrains are generally favourable for the

development of seepage forces within the pyroclastic sediments (tuff and pumice

horizons) and unconsolidated deposits during periods of rainfall. This indicates

that precipitation is one of the potential triggering factors for the slope failure in

the Debre Sina area. Besides, the concave shape of the terrain is enhancing the

convergence of groundwater flow into the landslide area. The relationship between

rainfall and landslide events indicates that deep-seated failures prevail after a

longer period of intensive rainfall. The presence of multiple sets of joints in rock

masses and intense rainfall further accelerate the slope’s failure.

It is evident that a pore pressure would further reduce the FS and influence the

slope stability. The calculation of the FS shows a significant difference between the

LE and FE methods. The FS from LEM with seismic load ranged between 0.92 and

4.76, while critical SRF from FEM were between 0.58 and 1.03 for 2D FEM

simulations. The minimum FS is calculated to be 0.58 for the Yizaba slope section


163

using the FEM for the saturated condition with seismic load, which has a similar

condition to the massive landslide of September 2005. This shows that the stability

of the area becomes unstable when there is a combination of saturation and seismic

load. This means that tectonic activity plays a well-defined role in promoting

landslides with seismicity as a possible predisposing factor or even a trigger, as the

area is in close proximity to the most seismically active regions in the country. The

critical SRF from FE analysis are significantly lower than the FS from LE analysis.

The low value of SRF emphatically suggests that the rock slope can be vulnerable

to failure under influence of any triggering force. Both methods show a significant

decrease in the obtained FS values for the study area when the seismic hazard is

considered but LEM always results in higher FS values. From the FE analysis, it

can be inferred that the study area is critically unstable and any small scale

disturbance will further reduce the FS and cause failure, particularly when the

area experiences heavy rainfalls or earthquakes. The numerical analysis showed

that the presence of joints in rock masses considerably affects the slope safety

factors. From the finding, it is concluded that the FEM-based slope stability

analysis with the SSR technique is found more applicable for stability assessment

because of the complex geometry, the heterogeneous material and the failure-

dominating faults in the study area. In general, the main controlling factors for

landslide problems in the area are the intensive weathering of the rocks; the

prominent geological structures; steep slope-gradient; the groundwater pressure

developed during precipitation; and the presence of low permeability beds which

force the percolating water to flow laterally.

6.2 Future research perspectives

The following recommendations are made for further action and landslide research

in the Debre Sina area in particular and the highlands of Ethiopia in general. Since

rainfall and earthquake-induced landslides are the major failure triggers in the

northwestern rift escarpment of Ethiopia, monitoring rainfall (such as intensity,

duration and antecedent) and groundwater fluctuation, is vital for proper landslide

hazard prediction and prevention. This could be used to develop early warning

system for mass movement hazards. For the future, comprehensive data


164

generation using borehole drilling, periodic monitoring of the rate of movement

and additional geophysical data collection using different techniques must be

conducted to have a wider and more comprehensive view of the area from a

mechanical and geological points of view. In order to improve understanding of

triggering factors of large-scale rock slides, further research on slope stability

analyses should be performed using field monitoring and numerical modelling

based on 3D finite elements. The study area is in close proximity to one of the most

seismically active regions in the world. Therefore, evaluating seismic activity and

its concomitant impacts on deep-seated landslides to obtain a complete view of the

area from a tectonic perspective should be further explored. As there are intensive

urban and infrastructural developments taking place along hills and rugged

mountains in the study area and surroundings, serious attention is also required

to consider the seismic hazard in the area during planning, design, and

construction phases. Much of the mountainous terrains of central highlands of

Ethiopia remain highly fragile in terms of mass movements; thus, any external

factors such as heavy rainfall or excavation could lead to slope failure. Prior to any

development planning, it is advisable to undertake proper landslide hazard

assessment and risk analysis in certain areas. Landslide and landslide-related

hazards are one of the major natural hazards causing tremendous losses in the

country but no attention is given at the moment. So far, the economic, social and

environmental impacts of mass movement hazards in Ethiopia have not been

widely recognized as a problem of national concern. Thus, a continuous research

work on landslide and related hazards in Rift margin and highland terrains is

highly recommended to increase the level of understanding on both local and

regional scales. This would be aid to finally reduce the damages and risks

prevailing on intense environmental degradation and failure of major

infrastructures. There are no historical records of landslides that indicate time of

occurrences, its magnitude, travelling distances, triggering factor, and associated

damage. It is, therefore, crucial to establish a landslide inventory data at least in

the major areas known for their landslide hazard in the country to properly address

and predict the time of occurrences, expected magnitude and travel distances as

well as associated damages.

Declaration of authorship

165


Chapter 2

Citation: Tesfay Kiros Mebrahtu, Bedru Hussein, Andre Banning, Stefan

Wohnlich, 2020a. Predisposing and triggering factors of large-scale landslides in

Debre Sina area, central Ethiopian highlands. Bull of Eng Geol Environ. 80:1–19.

DOI: 10.1007/s10064-020-01961-1.

Declaration of authorship: Tesfay Kiros Mebrahtu (TKM) as first author, was

responsible for the whole process including fieldwork, data collection, database

preparation, data analysis and interpretation. TKM also wrote the manuscript

including comments from Bedru Hussien (BH), Andre Banning (AB) and Stefan

Wohnlich (SW) on the data interpretation and presentation. AB was involved in a

detailed review of the manuscript prior to submission. TKM finalized the

manuscript for journal submission after a consensus is reached with BH, AB and

SW. All authors gave their approval of the final manuscript to be published.

Chapter 3

Citation: Tesfay Kiros Mebrahtu, Michael Alber, Stefan Wohnlich, 2020b. Tectonic

conditioning revealed by seismic refraction facilitates deep-seated landslides in the

western escarpment of the Main Ethiopian Rift. Geomorphology 370, 107382. DOI:

10.1016/j.geomorph.2020.107382.

Declaration of authorship: TKM conducted the field investigations, data analyses

and interpretation and wrote the manuscript while receiving comments from

Michael Alber (MA) and SW, and finalized the manuscript for journal submission

after a consensus was reached with MA and SW. All authors gave their approval

of the final manuscript to be published.

Chapter 4

Citation: Tesfay Kiros Mebrahtu, Andre Banning, Ermias Hagos, Stefan Wohnlich,

2021. The effect of hydrogeological and hydrochemical dynamics on landslide


166

triggering in the central highlands of Ethiopia. Hydrogeol J 29, 1239–1260. DOI:

10.1007/s10040-020-02288-7.

Declaration of authorship: TKM conducted the fieldwork, collected water samples,

developed the method and obtained all results in consultation with all authors. The

manuscript was written by TKM and reviewed by AB, Ermias Hagos (EH) and SW.

AB, EH and SW were involved in a detailed review of the manuscript prior to

submission. All authors gave their approval of the final manuscript to be published.

Chapter 5

Citation: Tesfay Kiros Mebrahtu, Thomas Heinze, Stefan Wohnlich, Michael Alber,

(submitted). Slope stability analysis of deep-seated landslides using Limit

Equilibrium and Finite Element methods under static and seismic load in Debre

Sina area, Ethiopia.

Declaration of authorship: TKM carried out the fieldwork, collected rock samples,

did the laboratory analysis and interpretation of the data, conducted the numerical

simulations and wrote the manuscript while taking comments from Thomas

Heinze (TH), SW and MA, and finalized the manuscript. TH and MA have been

involved in a detailed review of the manuscript prior to submission. All authors

gave their approval of the final manuscript to be published.

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Appendix

189

Appendix

Appendix A: Enhanced Landsat image mosaics

Appendix

190

Appendix

191

Appendix

192

Appendix

193

Appendix B: Structural data

SID X_coord Y_coord Dazim Damt SID X_coord Y_coord Dazim Damt

1 583856 1089472 325 80 52 583930 1091433 120 55

2 583825 1090016 240 85 53 583988 1093087 105 88

3 584464 1088136 160 80 54 582794 1092931 130 85

4 585171 1088105 320 87 55 584196 1094688 325 45

5 582289 1094649 30 55 56 584160 1094871 200 55

6 592804 1097482 225 60 57 584264 1095291 180 78

7 592704 1097519 350 84 58 584463 1095581 320 85

8 592672 1097519 250 85 59 584476 1095695 300 80

9 592905 1097405 180 78 60 584448 1096052 105 88

10 592353 1097378 290 10 61 584546 1096165 340 55

11 591283 1097384 135 55 62 584550 1096264 40 60

12 591174 1097349 210 15 63 584907 1096746 310 88

13 591141 1097355 50 85 64 584996 1096940 80 89

14 591136 1097346 205 82 65 584935 1096961 75 87

15 590947 1097255 215 87 66 585015 1096994 290 60

16 590927 1097240 330 85 67 583222 1092728 50 85

17 590865 1037103 105 25 68 583448 1093144 290 10

18 590594 1097085 340 85 69 583751 1093984 30 70

19 590193 1097100 210 75 70 583861 1093797 30 55

20 589878 1097159 150 75 71 585172 1094387 70 87

21 589874 1097082 270 70 72 585389 1094485 180 80

22 589841 1097124 145 35 73 585840 1094453 135 55

23 589711 1097139 292 64 74 586086 1094471 310 88

24 580192 1095213 110 53 75 585909 1093746 215 20

25 579726 1095808 60 75 76 585830 1093735 180 79

26 579669 1095746 80 25 77 585790 1093774 205 83

27 584389 1094372 170 60 78 585741 1093782 50 86

28 584781 1094877 320 54 79 585668 1093802 290 15

29 584398 1094472 40 85 80 583988 1093708 320 88

30 584063 1093380 20 70 81 583735 1093384 300 80

31 590927 1097240 320 85 82 583681 1093275 250 85

32 590865 1037103 180 80 83 583657 1093220 80 70

33 589841 1097124 145 35 84 583578 1092958 350 85

34 577765 1085761 292 64 85 583171 1092704 340 85

35 580192 1095213 20 60 86 583161 1092715 350 83

36 579669 1095746 50 25 87 584021 1093346 270 70

37 592507 1097337 320 30 88 584222 1092779 120 55

38 592818 1097424 160 50 89 584072 1092636 100 60

39 590615 1092747 330 80 90 584026 1092510 310 65

40 588079 1091626 310 87 91 584032 1092458 135 55

41 587865 1091812 325 78 92 583939 1092355 310 88

42 587295 1092135 180 80 93 583792 1092244 180 80

43 586863 1092516 70 87 94 583570 1092154 50 85

44 588452 1091706 45 83 95 583749 1091840 120 55

45 584557 1088333 330 78 96 583421 1091780 215 82

46 587326 1091693 70 85 97 583435 1091731 315 70

47 584175 1094249 80 89 98 583439 1091732 15 65

48 583727 1094316 350 80 99 583483 1091709 295 40

49 582911 1092704 150 75 100 583422 1091400 330 80

50 583450 1092652 270 70 101 583431 1091350 88 30

51 584504 1092783 180 78 102 583445 1091293 40 50

Appendix

194


103 583478 1091290 320 55 157 586212 1086818 270 65

104 583489 1091278 80 45 158 586192 1086731 70 85

105 583521 1091260 350 80 159 587078 1084806 50 70

106 583539 1091259 40 52 160 587018 1084677 270 78

107 583560 1091239 270 80 161 586988 1084658 285 55

108 583586 1091202 70 55 162 587064 1084709 80 85

109 583680 1091106 20 60 163 587116 1084626 345 55

110 583711 1090897 340 75 164 587502 1084487 5 65

111 583839 1090858 15 80 165 587412 1084394 215 88

112 583749 1090680 350 70 166 587258 1084308 355 50

113 583732 1090622 290 50 167 587065 1084281 20 55

114 583868 1090040 285 45 168 586652 1084174 330 54

115 583862 1089948 288 80 169 586451 1084155 25 60

116 583978 1089874 20 56 170 586283 1083986 42 70

117 583975 1089817 324 60 171 585831 1083777 290 80

118 583981 1089783 40 55 172 592662 1100334 40 65

119 583967 1089616 345 85 173 592254 1100254 20 55

120 587635 1089894 60 65 174 591908 1100232 282 85

121 586759 1090489 140 80 175 591705 1100195 340 80

122 586727 1090483 70 55 176 591699 1100226 180 80

123 586701 1090529 32 50 177 591703 1100238 335 50

124 586687 1090643 340 75 178 590582 1092559 180 75

125 586710 1090675 335 85 179 590598 1092571 80 87

126 586541 1090691 290 50 180 590614 1092574 95 25

127 586531 1090820 30 65 181 590662 1092567 50 85

128 586699 1091164 350 83 182 590583 1092353 215 88

129 586577 1091199 70 55 183 590224 1092258 350 84

130 586339 1091177 75 70 184 589702 1092192 205 82

131 586204 1091375 88 89 185 589518 1092141 290 10

132 586120 1091540 180 55 186 589288 1092083 170 80

133 586035 1091692 280 50 187 588941 1091768 10 60

134 585886 1091899 40 70 188 588673 1091550 350 84

135 585860 1091908 320 85 189 582664 1090472 300 50

136 585790 1091802 20 55 190 582626 1091083 290 75

137 585726 1091686 356 50 191 582616 1091104 170 80

138 585730 1091621 57 55 192 583200 1088773 355 85

139 585864 1091022 325 60 193 584771 1087835 20 58

140 586968 1089068 25 55 194 584767 1087644 270 75

141 584955 1088260 280 30 195 584777 1087562 82 88

142 584986 1088260 15 60 196 585053 1087234 303 68

143 584990 1088233 290 50 197 585018 1087134 343 50

144 585026 1088195 18 56 198 584873 1087345 272 78

145 585038 1088168 280 55 199 585305 1087410 302 52

146 585059 1088150 305 70 200 583260 1088296 270 75

147 585074 1088144 84 52 201 577015 1084216 210 10

148 585174 1088105 284 70 202 572969 1081414 300 80

149 585506 1087783 25 50 203 585944 1094741 235 85

150 585541 1087738 70 50 204 584017 1093702 180 78

151 585519 1087640 180 60 205 584249 1092485 315 87

152 585330 1087565 35 56 206 577571 1080734 277 15

153 585748 1087212 180 50 207 587176 1084862 30 60

154 585935 1087197 80 87 208 586831 1084831 215 87

155 586072 1087247 350 40 209 585434 1085307 340 64

156 586196 1086918 50 55 210 584243 1085739 285 55

Appendix

195


211 584394 1094469 105 85 265 590676 1092544 350 84

212 584079 1093388 325 60 266 584869 1087769 28 35

213 583900 1093843 15 75 267 584873 1087709 323 60

214 583531 1091939 320 60 268 584873 1087709 82 88

215 583531 1091939 10 60 269 584873 1087709 330 48

216 583575 1091916 320 60 270 584873 1087709 272 78

217 583575 1091916 15 65 271 585421 1089917 87 40

218 583575 1091916 350 85 272 585386 1089961 350 30

219 583575 1091916 10 65 273 585332 1089992 75 45

220 583537 1091500 320 55 274 585347 1090052 20 52

221 583631 1091466 324 60 275 585389 1090198 280 45

222 583631 1091466 350 70 276 585440 1090383 10 55

223 583631 1091466 290 45 277 585479 1090466 25 55

224 583679 1091409 290 60 278 585553 1090503 10 52

225 583679 1091409 10 60 279 585559 1090512 300 60

226 585047 1088467 20 60 280 585587 1090465 40 50

227 585047 1088467 322 60 281 585904 1091883 320 60

228 585079 1088467 55 65 282 598205 1082170 250 60

229 585119 1088402 20 55 283 598150 1082208 215 70

230 585119 1088402 10 60 284 597951 1082365 10 58

231 585166 1088351 275 60 285 597973 1082475 290 62

232 585166 1088351 20 56 286 597928 1082538 300 60

233 585166 1088351 304 45 287 597831 1082671 340 55

234 585166 1088351 298 52 288 597505 1082785 235 60

235 585266 1088312 50 40 289 597380 1082789 240 57

236 585266 1088312 30 50 290 597155 1082706 240 62

237 585266 1088312 60 45 291 597093 1082693 295 60

238 585266 1088312 335 60 292 596975 1082647 255 68

239 585266 1088312 75 60 293 596892 1082612 185 70

240 585266 1088312 284 70 294 595881 1082996 210 55

241 585422 1087772 270 65 295 594999 1083288 135 65

242 585422 1087772 70 85 296 594785 1083301 235 85

243 587081 1084865 20 60 297 594734 1083325 250 68

244 587157 1084916 60 58 298 594659 1083352 220 60

245 587157 1084916 320 60 299 594593 1083376 240 70

246 591797 1100402 34 53 300 594559 1083381 210 60

247 591797 1100402 322 65 301 594447 1083466 240 60

248 591797 1100402 346 52 302 594447 1083478 210 60

249 591797 1100402 30 60 303 594276 1083564 215 55

250 591797 1100402 300 62 304 594269 1083568 220 60

251 591797 1100402 315 65 305 594188 1083681 215 65

252 591797 1100402 40 60 306 596847 1082586 110 83

253 591795 1100445 310 57 307 596778 1082556 340 85

254 591795 1100445 285 50 308 596631 1082491 65 83

255 591795 1100445 56 62 309 596543 1082533 344 85

256 591795 1100445 76 50 310 596457 1082566 347 87

257 590675 1092766 336 55 311 596368 1082706 58 86

258 590675 1092766 278 50 312 596349 1082753 295 85

259 590598 1092571 340 50 313 596166 1082858 293 88

260 590598 1092571 40 65 314 596153 1082910 328 86

261 590662 1092567 322 57 315 596135 1082953 5 80

262 590662 1092567 65 55 316 596070 1083004 270 90

263 590676 1092544 35 53 317 595955 1082994 15 85

264 590676 1092544 314 45 318 595929 1082986 50 80

Appendix

196

SID X_coord Y_coord Dazim Damt

319 595712 1083025 306 70

320 595638 1083002 345 65

321 595579 1083023 80 56

322 595518 1083084 290 86

323 595472 1083118 340 80

324 595338 1083204 64 75

325 595271 1083219 275 85

326 595140 1083212 10 50

327 594074 1083725 60 65

328 594010 1083756 22 75

329 594002 1083860 20 80

330 593985 1083880 306 85

331 593944 1083910 70 88

332 593876 1084022 310 85

333 593731 1084106 320 75

334 593676 1084157 60 60

335 593558 1084241 46 70

336 593552 1084281 290 60

337 593497 1084377 355 65

338 593424 1084464 272 86

339 593364 1084439 335 85

340 593249 1084478 15 84

341 593190 1084545 350 83

342 593049 1084535 340 85

343 592972 1084442 95 25

344 592806 1084554 230 85

345 592790 1084599 250 85

346 592652 1084629 300 80

347 592547 1084681 210 10

348 592503 1084725 180 80

349 592422 1084829 210 75

350 592448 1084926 170 80

351 592433 1084962 290 10

352 592223 1085013 50 85

353 592099 1085190 205 82

354 592055 1085270 215 88

355 591971 1085393 20 55

356 591844 1085402 330 85

357 591727 1085412 310 88

358 591657 1085478 210 15

359 591619 1085640 180 80

360 591524 1085754 250 68

Curriculum Vitae

197

Curriculum Vitae

Personal details

Name Tesfay Kiros Mebrahtu

Address Institute of Geology, Mineralogy and Geophysics

Universitätsstraße 150

44801 Bochum, Germany

Email: [email protected]/[email protected]

Education

Sep. 2009 – Jul. 2011 Master of Science (MSc) in Geological Engineering

Mekelle University, Ethiopia

Sep. 2004 – Jul. 2008 Bachelor of Science (BSc) in Applied Geology

Mekelle University, Ethiopia

Professional experience

Oct. 2015 – Present Research Assistant (PhD candidate) in Applied Geology

Department (Hydrogeology working group) at the

Institute of Geology, Mineralogy and Geophysics

Ruhr-Universität Bochum, Germany

Jul. 2011 – May 2015 Academic staff at Department of Geology

Addis Ababa Science and Technology University

Addis Ababa, Ethiopia

Oct. 2008 – Aug. 2009 Junior Hydrogeologist

Relief Society of Tigray (REST), Mekelle, Ethiopia

mailto:[email protected]

mailto:[email protected]

Publications

198

Publications

Mebrahtu, T.K., Hussien, B., Banning, A., Wohnlich, S. (2020a). Predisposing and

triggering factors of large-scale landslides in Debre Sina area, central Ethiopian

highlands. Bulletin of Engineering Geology and the Environment.

Mebrahtu, T.K., Alber, M., Wohnlich, S. (2020b). Tectonic conditioning revealed by

seismic refraction facilitates deep-seated landslides in the western escarpment of

the Main Ethiopian Rift. Geomorphology. 370. 107382.

Mebrahtu, T.K., Banning, A., Hagos, E., Wohnlich, S. (2021). The effect of

hydrogeological and hydrochemical dynamics on landslide triggering in the central

highlands of Ethiopia. Hydrogeology Journal.

Mebrahtu, T.K., Heinze, H., Wohnlich, S., Alber M. (submitted). Slope stability

analysis of deep-seated landslides using Limit Equilibrium and Finite Element

methods under static and seismic load in Debre Sina area, Ethiopia.

Scientific Meetings

Mebrahtu, T.M., Wohnlich, S., Alber, M., Hussien B., Banning, A. (2019).

Integrated approach to unravel mechanics of slope failure inducing landslides in

Debre Sina area, central highlands of Ethiopia. 2018 WMESS, Prague, Czech

Republic (Presentation).

Kiros, T., Wohnlich, S., Alber, M., Hussien, B. (2018). Oblique divergence

activating large-scale rainfall induced landslides: Evidence from Tarma Ber,

Northwestern Plateau of Ethiopia. – AGU Fall Meeting 2018, Washington, D.C,

U.S.A (Poster).

Mebrahtu, T., Wohnlich, S. (2018). The effect of groundwater and rainfall on

landslide triggering in the central Highlands of Ethiopia: the case of Debre Sina

Area. – 26. Tagung der Fachsektion Hydrogeologie (FH-DGGV), Bochum, Germany

(Poster).

Publications

199

Kiros, T., Wohnlich, S., Hussien, B. (2017). Large-scale landslide triggering

mechanisms in Debre Sina area, Central Ethiopian Highlands at the western Afar

rift margin. – AGU Fall Meeting 2017, New Orleans, U.S.A (Poster).

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