UNDERSTANDING TALUS SLOPE GEOMORPHOLOGY FROM PERMAFROST DISTRIBUTION USING GEOPHYSICAL METHODS AND DETAILED GEOMORPHOLOGICAL MAPPING
COL DU SANETSCH, SWITZERLAND
Word count: 25817
Joke Laporte
Stamnummer: 01411341
Promotor: dr. Amaury Frankl – UGent – physical geography Copromotor: Prof. dr. Reynald Delaloye – Université du Freibourgh Begeleider: drs. Hanne Hendrickx – UGent – physical geography
Masterproef voorgelegd voor het behalen van de graad master in de richting Geografie
Academiejaar: 2018 - 2019
I PREFACE
This thesis describes the results from a study about permafrost in the European Alps. Because of my passion for nature, mountains and my concerns about the current climate change I wanted to get involved and contribute to this scientific fields. My interest helped me when I had hard times and finishing this dissertation shows me that even though something looks impossible, if I work hard I’m able to finish it. However it wouldn’t be possible without the help and support from others. Therefore I want to thank everyone who supported me along this process. First of all my supervisor Hanne Hendrickx and promotor Amaury Frankl. It was never a problem to pass by when I was struggling and they provided me with valuable tips, insights and feedback. Furthermore, my field research wouldn’t be possible without them, my fellow student Ewout and the other visitors who passed by in Switzerland. Geophysical measurements and drone campaigns can’t be done on your own and are quit hard in a challenging environment such as the Alps. So thank you for all your help and support!
II POPULARIZING TEXT
Climate change is a hot topic, global temperatures are rising, precipitation patterns are changing. In the Alpine environment, permafrost, ground that remains at or below 0°C for at least two consecutive years, is present but degrading. As a result the environment will experience important modifications. Because of the causal relationship between the thermal regime and geohazards in mountain permafrost regions, scientific attention raises. This research forms a case study about the influence of permafrost on the geomorphology of a talus slope, how are they related, which trends can be found? Based on drone imaginary a geomorphological map of the study area was made. By combining temperature data, information about the horizontal displacement of the surface layer, measurements of permafrost probability and the geomorphological map, the interrelationships are explored and linked to the existing literature and theories.
POPULARISERENDE TEKST
Vandaag de dag staat klimaatsverandering hoog op de agenda. De temperatuur neemt toe en neerslag patronen veranderen. In Alpiene gebieden resulteert dit in een degradatie van permafrost, grond dat gedurende minimaal twee jaar een temperatuur lager dan 0°C heeft. Aangezien er een causaal verband is tussen permafrost degradatie en de toename aan geohazards in periglaciale gebieden is het belangrijk deze processen goed in kaart te brengen en te begrijpen. Dit onderzoek is een case studie waarin nagegaan wordt wat de invloed is van permafrost op de geomorfologie van een talus slope. Door het combineren van geomorfologische data, permafrost distributie, meteorologische data en horizontale oppervlakte snelheden, worden de relaties bestudeerd en gerelateerd met bestaande literatuur en theorieën.
ABSTRACT
Within this research we try to understand the influence of permafrost on talus slope geomorphology.
Based on temperature data, measurements about the surface velocity and geophysical transects the permafrost distribution on the talus slope Col du Sanetsch is explored and explained. This information is used to interpret the geomorphology of the talus slope and its interrelationship with permafrost existence. A geomorphological map was developed by the combination of a high resolution DEM, resulting from an UAV campaign in 2018, and field observations. We concluded that the talus slope on Col du Sanetsch is a good example to explain the evolution, concepts and processes which influence the geomorphology of talus slopes. Furthermore, this map suggest permafrost by the existence of a rock glacier and protalus rampart, gelifluction and stress related landforms, such as transverse ridges.
Nivation processes are related to the long lasting snow in this periglacial environment. According to our geophysical measurements a possible or probable existence of permafrost is present in a coarse grained area on the talus slope, the rock glacier and the landslide. Temperatures from springs at the foot of the talus slope and the Winter Equilibrium Temperature (WEqT) can be used as indicators for permafrost distribution. However, when using the WEqT as an indicator it is important to incorporate the context. When you only have data from one year, the intra-site differences of WEqT are equally important as the absolute value to indicate permafrost. Furthermore, the presence of unfrozen water
III can have a positive effect on the WEqT due to the latent heat released while freezing. The permafrost distribution is influenced by the temperature. The Mean Annual Ground Surface Temperature (MAGST) is lowest in areas were permafrost exist. A long lasting snow layer and a coarse grained surface negatively influence this temperature and will favour permafrost existence.
One of the main landforms on the talus slope is a landslide. Due to gelifluction, this landform is slowly moving downwards. A network of 28 points is measured every summer since 2011 to monitor the displacement rates. High horizontal surface velocities are measured within areas and years with high WEqT. The intra-site variations mainly depend on the difference in moisture content and the presence of an impermeable layer resulting from permafrost or seasonal frost. The several transverse ridges which can be found on the landslide are proves of this displacement.
In the summer of 2018, a webcam was installed to observe the evolution of the snow layer through the year. This will partly close the knowledge gap resulting from the absence of precipitation and snow data and make it possible to further explain and investigate the observed patterns.
SAMENVATTING
Deze studie tracht de invloed van permafrost op de geomorfologie van de talus slope op Col du Sanetsch te begrijpen. Een geomorfologische kaart is ontworpen door de combinatie van een hoog resolutie DEM en veldwerk. Het DEM resulteert uit luchtfoto’s verkregen tijdens een UAV campagne in de zomer van 2018. De permafrost distributie op Col du Sanetsch werd in kaart gebracht met behulp van geofysische metingen en indicatoren zoals de Winter Equilibrium Temperature (WEQT) en de temperatuur van de verschillende bronnen aan de voet van de talus slope. Aan de hand van meteorologische en topografische data wordt getracht de distributie te begrijpen en uit te leggen. Deze informatie wordt op zijn beurt gebruikt om de geomorfologie en de relatie met de permafrost distributie te interpreteren.
De talus slope op Col du Sanetsch is een goed voorbeeld om de evolutie, concepten en processen die aan de basis liggen van de geomorfologie van een talus slope uit te leggen. Daarnaast wordt de aanwezigheid van permafrost gesuggereerd door het voorkomen van een rotsgletsjer, gelifluctie en landvormen, zoals transversale ruggen, die gerelateerd worden met stress en interne deformatie. Verder zijn ook nivatie processen aanwezig in deze periglaciale omgeving. Geofysische metingen bevestigen de mogelijke en waarschijnlijke aanwezigheid van permafrost in de rotsgletsjer, landslide en de gebied op de talus slope dat bedekt wordt door grotere rots blokken. De distributie van permafrost wordt beïnvloed door de temperatuur, de Mean Annual Ground Surface Temperature (MAGST) is het laagst in gebieden waar permafrost voorkomt. Een sneeuwlaag die lang blijft liggen en de aanwezigheid van een oppervlakte laag met grote rots fragmenten beïnvloeden de grond temperatuur negatief en hebben op deze manier invloed op de distributie van permafrost.
IV Een opvallende landvorm op de talus slope is de landslide. Deze verplaatst zich langzaam richting de vallei via gelifluctie. Op deze landvorm werden 28 punten gemarkeerd die sinds 2011 ieder jaar opnieuw opgemeten worden. Dit maakt het mogelijk om de evolutie en ruimtelijke variatie in deze verplaatsing te analyseren. Een exponentiële relatie werd gevonden tussen de WEqT en de horizontale snelheid. De intra-site variaties worden voornamelijk gelinkt met de verschillende vochtigheidsgraad van de bodem en de aanwezigheid van een ondoordringbare laag, als resultaat van permafrost of seizoensgebonden vorst. De vele transversale ruggen die aanwezig zijn op de landslide getuigen van deze verplaatsing.
In de zomer van 2018 werd een webcam geïnstalleerd. Deze maakt het mogelijk om de evolutie van de sneeuwlaag doorheen het jaar te observeren. Op deze manier kan dieper ingegaan worden op de rol van de sneeuwlaag in de verschillende processen en trends.
V TABLE OF CONTENT
1. Introduction ... 1
2. High Alpine talus slopes and their interaction with permafrost distribution ... 2
2.1 Geomorphology of talus slopes ... 2
2.1.1 Formation of talus slopes ... 2
2.1.2 The cross section of a talus slope ... 3
2.1.3 Geomorphological transport processes and their resulting landforms on talus slopes ... 3
2.2 Distribution and evolution of mountain permafrost ... 4
2.2.1 Definition of mountain permafrost ... 4
2.2.2 Understanding permafrost distribution on mountain slopes ... 5
2.2.3 Mountain permafrost maps and modelling ... 9
2.2.4 Permafrost degradation ... 9
2.3 Geomorphological processes and dynamics on talus slopes as impacted by permafrost degradation ... 11
3. Study objectives ... 12
4. Study site ... 12
5. Material and methods ... 15
5.1 Topographic survey based on field observations and UAV ... 16
5.1.1 Field observations ... 17
5.1.2 Data acquisition with UAV ... 17
5.1.3 Real-Time Kinematic GNSS (RTK-GNSS)... 17
5.1.4 Data processing in Agisoft PhotoScan ... 18
5.1.5 Geomorphological mapping based on DEM interpretation ... 20
5.1.6 Defining topographic roughness ... 20
5.2 Surface velocity measurements on the landslide ... 22
5.3 Temperature measurements... 24
5.4 Permafrost mapping based on geophysical measurements ... 25
5.4.1 VES ... 27
5.4.2 ERT ... 28
5.4.3 Transects ... 31
5.4.4 Defining permafrost probability classes ... 34
VI
5.5 Statistical analyses ... 34
5.5.1 Analysis of the meteorological trends ... 35
5.5.2 Analysing temporal and spatial differences in surface velocity and their relation with temperature, permafrost distribution and geomorphology ... 35
5.5.3 Analysing the distribution of surface roughness and the interrelation with permafrost distribution ... 36
5.5.4 Comparing the measured permafrost distribution and permafrost probability maps ... 36
6. Results ... 37
6.1 Geomorphological map ... 37
6.1.1 Gravitational processes ... 37
6.1.2 Snow, frost and (peri)glacial landforms ... 37
6.2 Temperature measurements... 42
6.3 Geophysical survey ... 44
6.3.1 VES ... 44
6.3.2 ERT transects ... 46
6.4 Annual surface velocity of the landslide ... 49
6.5 Relations between geomorphological characteristics, permafrost distribution, meteorological parameters and surface velocity on the landslide... 50
6.5.1 Surface velocity related to meteorological factors, permafrost distribution and topographical characteristics ... 50
6.5.2 Surface roughness as an explanatory factor of permafrost distribution ... 53
7. Discussion ... 55
7.1 Geomorphological mapping ... 55
7.1.1 Gravitational processes ... 55
7.1.2 Snow, frost and (peri)glacial landforms ... 56
7.2 Meteorological analysis ... 59
7.3 Permafrost distribution indicated by geophysical and temperature measurements ... 61
7.3.1 Permafrost distribution on the talus slope ... 61
7.3.2 Permafrost distribution on the rock glacier ... 62
7.3.3 Permafrost distribution on the landslide ... 64
7.3.4 Differences in apparent resistivity between landslide and talus slope ... 65
VII
7.3.5 Permafrost probability maps ... 65
7.4 Annual surface velocity ... 69
7.5 Possibilities for future research ... 70
8. Conclusions ... 70
9. References ... 71
Literature ... 71
Maps and data ... 80
10. Attachments ... 81
10.1 Attachment 1: VES field template: Wenner array ... 81
10.2 Attachment 2: VES field template: Schlumberger array ... 82
LIST OF FIGURES Figure 1: Surface temperature anomalies, relative to the period 1961 – 1990, in the Swiss Alps (Säntis, Lugano and Zürich) compared to the global anomalies. ... 1
Figure 2: resistivity of soil, rock and minerals ... 7
Figure 3: The chimney effect ( ... 8
Figure 4: Permafrost degradation ... 10
Figure 5: General meteorological context: long-term means of monthly mean temperature, monthly maximum and minimum temperature as well as monthly precipitation sums in Sion ... 13
Figure 6: Installation and measurement of GCPs. a) a fixed GCP, b) a ‘cloth’ GCP ... 18
Figure 7: a) Workflow in Agisoft to build a DEM based on Hendrickx et al. (2019). b) A model after the first alignment. In the upper zone, different gaps are visual, here we added more photographs. ... 19
Figure 8: a) Flow chart to calculate SDrestopo. b) Example of the residual topography, the difference between the LiDAR DTM and mean DTM results in the residual topography. c) The influence of the moving-window size on the surface roughness... 21
Figure 9: Topographic roughness classes based on photographs ... 22
Figure 10: Overview data points: temperature and surface velocity (source: Google earth, 2016) ... 23
Figure 11: Principles of ERT prospecting. a) Relation between the resistivity and the temperature. b) An example of the electrical field created by inserting I into the quadripole. c) The most used ERT electrical arrays ... 27
Figure 12: Relative position of the electrodes in a Wenner (a) and Schlumberger array (b), VES ... 28
Figure 13: concept of the Wenner-Schlumberger array ... 29
VIII Figure 14: Relative differences in apparent resistivity for transect CdSLS1a. Bad data points (green) can
be selected and removed. ... 30
Figure 15: RMS error statistics of CdSLS1a after a preliminary inversion. By moving the green line, bad data points can be selected and removed. ... 31
Figure 16: Overview geophysical measurements ... 33
Figure 17: Debris flow deposits and gullies ... 40
Figure 18: Landslide ... 40
Figure 19: Panorama of the talus slope ... 40
Figure 20: Rock glacier (dotted lines = ridges) ... 41
Figure 21: Nivation hollow (dotted lines = ridges) ... 41
Figure 22: Nivation zone on the talus slope, small depressions followed by an asymmetric ridge ... 41
Figure 23: Ridges in the nivation zones visualised using the DEM-derivative ‘aspect’ on a fine resolution DEM ... 41
Figure 24: Lateral arcuate ridge on the rock glacier, as seen from the UAV ... 41
Figure 25: Climatological graphs based on the output of 3 t-loggers located on the landslide at Col du Sanetsch (August 2014 – August 2018). ... 43
Figure 26: Results from the VES measurement (Schlumberger array) on landslide position 1 (t-logger 017 – down) (transect 1) ... 45
Figure 27: Results from the VES measurement (Schlumberger array) on landslide position 2 (t-logger 010 – mid) (transect 4) ... 45
Figure 28: Results from the VES measurement (Schlumberger array) on the talus slope, (t-logger AT – GST – 1, down) (transect 10) ... 45
Figure 29: Results from the VES measurement (Schlumberger array) on the talus slope, (t-logger AT – GST – 2, up) (transect 11) ... 45
Figure 30: Results from the VES measurement (Wenner array) on landslide position 2 (t-logger 010 mid). 0 = t-logger, ‘-20’ means the position 20 m from the t-logger, direction NE. The red line indicates the location of the ERT profile. (transect 7) ... 46
Figure 31: ERT measurements on landslide position 1 (t-logger 017 – down) ... 47
Figure 32: ERT measurements on the landslide position 2 (t-logger 010 - mid) ... 47
Figure 33: ERT measurements above the landslide, position 3 (t-logger 005 - up)... 48
Figure 34: Lateral ERT profile rock glacier ... 48
Figure 35: Vertical ERT profile rock glacier ... 49
IX Figure 36: Annual horizontal surface velocity on the landslide. Mean of a set of points selected in several
sections of the moving landform ... 49
Figure 37: Exponential relation between the annual horizontal displacement and the WEqT. ... 50
Figure 38: Annual and winter precipitation (mm) Sion and the annual horizontal displacement on Col du Sanetsch (m) ... 51
Figure 39: Relation between the displacement rate and slope (light blue = area above the landslide, dark blue = on the landslide) ... 52
Figure 40: Grain size distribution on the talus slope ... 53
Figure 41: Grain size distribution on the landslide ... 53
Figure 42: Landslide on Col du Sanetsch ... 57
Figure 43: Representation of an idealized landslide ... 57
Figure 44: Rooting zone of the landslide ... 58
Figure 45: Gully filled with materials from the landslide ... 59
Figure 46: Mean MAGST and average zero curtain period for the different t-loggers ... 60
Figure 47: Borehole temperatures at approximately 20 m depth compared to the WEqT and MAGST of Arp - 010 - mid ... 61
Figure 48: T-logger located below the front of the rock glacier (GST – 13 – RG) ... 64
Figure 49: Temperature of t-loggers GST-9-RG and GST-13-RG in 2016 - 2017 ... 64
Figure 50: Permafrost distribution: Geophysical measurements compared to the Swiss Potential Permafrost Distribution Map ... 67
Figure 51: Permafrost distribution: Geophysical measurements compared to the APIM ... 68
LIST OF TABLES Table 1: overview of some typical geophysical outcomes ... 6
Table 2: Possible influences of permafrost degradation on geomorphological processes ... 12
Table 3: Overview data ... 16
Table 4: Technical table UAV survey ... 20
Table 5: Topographic roughness classes ... 22
Table 6: Technical specifics of the t-loggers ... 24
Table 7: Overview of the t-loggers on Col du Sanetsch ... 25
Table 8: Metadata for the different transects ... 32
X Table 9: Permafrost probability classes ... 34 Table 10: Mean WEqT (Bolt = permafrost existence probable) ... 42 Table 11: Colour legend of permafrost classes, ERT profiles on the landslide ... 48 Table 12: Permafrost distribution from geophysical and temperature measurements related to topographic factors on the talus slope ... 62 Table 13: Permafrost distribution from geophysical and temperature measurements related to topographic factors on the rock glacier ... 63 Table 14: Permafrost distribution from geophysical and temperature measurements related to topographic factors on the landslide ... 65
LIST OF MAPS
Map 1: Location study site: Col du Sanetsch ... 14 Map 2: Location talus slope and different landforms ... 14 Map 3: Geomorphological map Col du Sanetsch ... 39 Map 4: Thematic map - slope, permafrost distribution and annual surface velocity (1 = slow moving frontal zone, 2 = fast moving median zone) ... 52 Map 5: Thematic map surface roughness and permafrost distribution ... 54
LIST OF ABBREVIATIONS
APIM = Alpine Permafrost Index Map BTS = Bottom Temperature of Snow DEM = Digital Elevation Model
ERT = Electrical Resistivity Tomography GCP = Ground Control Point
GPR = Ground Penetrating Radar GST = Ground Surface Temperature MAAT = Mean Annual Air Temperature
MAGST = Mean Annual Ground Surface Temperature PERMOS = Swiss Permafrost Monitoring Network SDrestopo = Standard Deviation of Residual Topography SRT = Seismic Refraction Tomography
UAV = Unmanned Aerial Vehicle VES = Vertical Electrical Sounding WEqT = Winter Equilibrium Temperature
1 1. INTRODUCTION
Since the late 19th century, the climate has changed rapidly because of anthropogenic impacts (Pachauri
& Meyer, 2014). Global temperatures are rising, precipitation patterns are changing. There have always been periods with increasing or decreasing temperatures, but since the 1970s this warming is particularly marked (Stocker et al., 2013, Haeberli & Beniston, 1998). Following the IPCC (Pachauri &
Meyer, 2014) it is 95 % certain that this current global warming is caused by humans. As can be seen in Figure 1, this temperature increase is more marked in the Alps, where the average temperature has risen twice as fast as the global average (Gobiet et al., 2014; Beniston, 2006). As this warming will continue during the 21st century (Stocker et al, 2013), the Alpine cryosphere will experience important modifications (Deluigi et al., 2017). Permafrost, ground that remains at or below 0°C for at least two consecutive years (Dobinski, 2011), degrade due to this climate change (Etzelmüller & Frauenfelder, 2009). Because of the causal relations between the thermal regime and geohazards in mountain regions, mountain permafrost has been receiving an increased scientific attention (Etzelmüller, 2013).
Examples of research topics related to permafrost warming are the influence on rock fall activity (e.g.
Gruber and Haeberli, 2007; Ravanel et al., 2010), rock glacier acceleration (e.g. Kääb et al., 2007; Roer et al., 2008; Delaloye et al., 2010) and sediment transfer (e.g. Lane et al., 2007; Kobierska et al., 2011).
The aim of this study is to understand the influence of permafrost distribution on talus slope geomorphology. The first section gives the state-of-the-art on the current knowledge and advances related to the scientific research on permafrost degradation and talus slope geomorphology. The second part is the core of the study.
Figure 1: Surface temperature anomalies, relative to the period 1961 – 1990, in the Swiss Alps (Säntis, Lugano and Zürich) compared to the global anomalies. (Source: Beniston, 2006)
2 2. HIGH ALPINE TALUS SLOPES AND THEIR INTERACTION WITH PERMAFROST DISTRIBUTION In order to analyse and understand the relations between talus slope geomorphology and permafrost distribution, it is important to have a theoretical background. Providing this theoretical background is the purpose of the first chapter. We first focus on the talus slope geomorphology and more specific the formation and transport processes. The second part elaborates on the distribution and evolution of mountain permafrost. In the last part, these topics are combined and we discuss possible influences of permafrost on high alpine talus slope geomorphology.
2.1 Geomorphology of talus slopes
Talus slopes, or scree slopes, are formed by unconsolidated clasts of different sizes accumulating at the foot of a rock cliff (Müller et al., 2013; Gutiérrez & Gutiérrez, 2016). They are one of the dominant types of debris storage and has a high potential of debris release and mass movements (Phillips et al., 2009). On an idealized periglacial mountain slope, a sequence of characteristic landforms can be found:
an upper convex segment, with the headwall, a talus slope, which normally has a constant slope angle, and a concave segment at the base. Depending on the actual situations, some of these components may not be present or negligible (Ritter et al., 1995; Huggett, 2017; Müller et al., 2014).
2.1.1 Formation of talus slopes
Talus slopes are the result of active sedimentation of materials destabilized on the headwall (Schoeneich et al., 2011; Müller et al., 2014). The destabilization of these materials depends on the shear stress and shear strength. Stress is formed by any force that tends to move the materials, for instance, gravity. Shear strength are the properties of the matter that resist the stresses generated by gravitation. Many different factors influence the shear strength and shear stress, for example: freeze- thaw cycles, the slope, roughness of the plane, size and shape of the particles and the cohesion between the particles. When the shear stress exceeds the shear strength the slope is unstable, and a small trigger may initiate a mass movement (Huggett, 2017; Ritter et al., 1995). The travelling distance of a boulder is a balance between the energy gain and loss. As a boulder moves downslope, he will gain energy due to his movement in the gravity field. Collisions and friction will result in a loss of energy.
When shear strength increases, by for instance an increase of surface roughness, or the shear stress decreases because of a less steep slope, this may result in a higher loss/ smaller gain of energy and the boulder may halt (De Blasio & Sæter, 2015; Ritter et al., 1995).
Two important processes for the development of a talus slope are frost weathering and paraglacial adjustment (Schoeneich et al., 2011; Ballantyne, 2002). These processes decrease the shear strength of the headwall and so they contribute to the production of debris. Frost weathering, or frost shattering, breaks off small grains and large boulders from the bedrock. It depends on freeze-thaw cycles, and is a function of freeze-thaw frequency, moisture content and the tensile strength of the bedrock. More information about this process can be found in other literature (e.g. Ritter et al., 1995; Hugget, 2017).
Another process that can be related to the origin of a talus slope is paraglacial adjustment. When a bedrock is covered by ice, this overlying ice induces internal stress levels that are much higher than
3 those that might be expected from the loading alone. Glacier downwastage result in a relaxation of tensile stresses within this rock mass. This stress release results in propagation of the internal joint network, a loss of cohesion and a reduction of internal locking stresses. This may lead to immediate or delayed rock-slope failure, deformation or rock fall, depending on the lithology and structure (e.g. joint density and weaknesses) of the rock mass (Ballantyne, 2002). The block and grains that breaks off, will move downslope by bouncing, rolling or sliding (Gutiérrez & Gutiérrez, 2016; Huggett, 2017; Ritter et al, 1995).
2.1.2 The cross section of a talus slope
Talus slope typically has a straight longitudinal profile, with a slope angle of 30° - 40° and a basal concavity (Gutiérrez & Gutiérrez, 2016; Luckman, 2013). The slope angle is determined by the friction angle, the angle at which a block will begin to slide, of the cohesionless debris (Gutiérrez & Gutiérrez, 2016; Ritter et al., 1995). On most talus slopes, an increasing grain size can be found downslope (Lambiel & Pieracci, 2008; Ritter et al., 1995; De Blasio & Sæter, 2015; Gómez et al., 2003). This process, known as fall sorting, is most marked at the bases of slopes. It can be explained by a combination of different mechanisms. As a large boulder has a greater momentum, it has the possibility to travel a greater distance. Furthermore, the frictional resistance of the surface depends on the relationship between the size of a moving boulder and the irregularities of the surface. In that way, boulders will only come to rest in areas with boulders of more or less the same size (Luckman, 2013).
De Blasio & Sæter (2015) did an experimental study of the behaviour of a single grain falling and travelling on a homogeneous granular bed. Small grains, which fall on larger grains, will stay close to the point of falling. Large grains who fall on a bed of smaller grains may behave in two manners. Or they stop immediately, in a self-created crater, or they roll down slope. Their abundance peak around the fall zone and further downslope. The degree of sorting depends on different aspects, length of the slope, cliff height and the size and shape of the dominant particles (Luckman, 2013).
2.1.3 Geomorphological transport processes and their resulting landforms on talus slopes
In addition to this ‘primary’ processes, which are responsible for the formation of the talus slope, different other geomorphological processes can be found. They will commonly rework the initial structure of a talus slope. In periglacial conditions processes related to frost, thaw and snow are highly active (Huggett, 2017). On the talus slope, snow avalanches and permafrost creep, for instance rock glacier creep, dominate (Müller et al., 2014). Other processes that occur are debris flows, solifluction, landslides and dry ravel (Ritter et al., 1995; Huggett, 2017).
Solifluction is a slow transport process related to freeze-thaw actions. Depending on the temperature of the soil, different solifluction processes may occur. A plug-like flow, which result from thawing ice lenses produced by an upward frost penetration, will only occur in cold permafrost regions (Matsuoka, 2001).
Other solifluction processes, such as diurnal frost creep, annual frost creep and gelifluction may influence slopes underlain by both cold or warm permafrost or even non-permafrost zones. All of those, slow, mass wasting processes result from freeze-thaw actions in fine-textured soils. They are controlled
4 by freeze-thaw cycles and the depth and thickness of ice lenses (Matsuoka et al., 2001). The specific form, frost creep, is a result of the expansion of the soil during freezing and a contraction during the thawing. This expansion is perpendicular on the slope, the contraction is vertical, because of the gravity.
The result is a net downslope moment of the materials (Huggett, 2017). Gelifluction is a very slow downslope movement of a water-saturated soil over the frozen ground during the summer months (Huggett, 2017). In this season ground become saturated because of the restricted drainage associated with a seasonally frozen water table or permafrost and the moisture delivered by thawing. This increased moisture may also intensify other mass movement processes (Hugget, 2017; Ritter et al., 1995). As the moving soil overturned the superficial soil, due to a reduced velocity, solifluction lobes, large tonguelike masses of surface debris, may be produced. Those are usually poorly sorted, and the form of these features depend on the texture, gradient and soil moisture. A fine-grained soil layer overriding a coarse sediment, will also produces solifluction lobes and sheets (Matsuoka, 2001; Huggett, 2017).
Fast acting processes that rework the structure of a talus slopes, are for example debris flows, landslides and dry ravel. Debris flows consist of a fast-moving body of sediment particles (70 – 90%) with water that moves downslope. An abundant source of moisture, fine grained sediments and relatively steep slopes are required conditions. They typically form in small gullies, where run off can be concentrated (Luckman, 2013). Landslides are rapid mass movements along clear-cut shear planes. Dry ravel is the movement of individual particles by rolling, bouncing of sliding (Huggett, 2017). All those processes will influence the form and profile of the talus slope. For instance, debris flows may gully the upper slope and produce debris cones where the transported materials accumulate (Luckman, 2013).
Snow and related processes may also influence talus slopes. As the frictional properties of snow- covered talus are quite different, large boulders can be trapped or slide over the surface. This result in another distribution of grain sizes (Luckman, 2013). Nivation hollows may also be formed. This are depressions, associated with late-lying or perennial snow patches, where frost action and intensified weathering may increase erosion (Gutiérrez & Gutiérrez, 2016). Nivation hollows can be associated with protalus ramparts. These ridges, at the downslope margin of a snow patch, consist of unsorted gravel and boulders. Rockfall debris from the cliff, rolls and slides down over the snow and accumulate at the foot, were it builds a ridge (Gutiérrez & Gutiérrez, 2016; Shakesby 2004). However, some other authors see protalus ramparts as embryonic rock glaciers (e.g. Scapozza et al., 2011; Haeberli, 1985). We decided to use the morphological definition. Furthermore, snow avalanches may erode loose, fine materials and transport them downslope, where they are deposited on coarser materials.
2.2 Distribution and evolution of mountain permafrost 2.2.1 Definition of mountain permafrost
The definition of a permafrost body is based on the following thermal criteria: ‘ground that remains at or below 0°C for at least two consecutive years’ (Dobinski, 2011; Gruber & Haeberli, 2009, Harris et al., 2009). It refers to a thermal state of the lithosphere, so it is not a material phenomenon (Dobinski, 2011).
Permafrost results from a negative energy balance and depends on incoming solar radiation and
5 sensible heat flux (Schoeneich et al., 2011). Mountain permafrost is found in mountain areas and will be influenced by the characteristics of topography, such as altitude and aspect (Gruber & Haeberli, 2009). This results in a complex, discontinuous permafrost distribution in the European Alps (Pieracci et al., 2008; Deluigi et al., 2017; Lambiel & Pieracci, 2008). Most of the mountain permafrost in the European Alps is temperate or warm, which means it will have a temperature close to 0° C (Etzelmüller
& Frauenfelder, 2009).
2.2.2 Understanding permafrost distribution on mountain slopes
Permafrost is invisible, which makes it hard to know its exact distribution on mountain slopes.
Theoretically, the lower limit of permafrost in the Alps is 2500 m.a.s.l. (Deluigi et al., 2017). In reality, this is influenced by many topoclimatic factors, such as altitude, slope, aspect, Mean Annual Air Temperature (MAAT), snow and solar radiation (Harris et al., 2009; Deluigi et al., 2017). All those factors can give a first indication were permafrost may be present. Furthermore, the presence of mountain permafrost can be indicated by thermal data, geomorphological phenomena and geophysical characteristics.
2.2.2.1 Thermal indicators
As permafrost is a thermal state, a first, but not absolute, hypothesis of the existence of permafrost can be made based on the temperature (Scapozza et al., 2011). A MAAT of -3°C is the theoretical threshold value that can be used to identify altitudinal bands in the European Alps that contain a significant amount of permafrost (Gruber & Haeberli, 2009). The MAAT is influenced by both the altitude and aspect of the slope. As a result, permafrost on northern slopes will occurs at lower altitudes than on southern slopes (Schoeneich et al., 2011). A commonly used thermal indicator for permafrost distribution is the Winter Equilibrium Temperature (WEqT). The WEqT is the mean temperature over a period of 30 days during which the snow layer is stable. This period ends 14 days before the start of the melting period. As snow is a good insulator, the WEqT depend on the ground thermal regime, and thus the presence or absence of permafrost (Schoeneich, 2011). The snow cover must have a depth of 80 – 100 cm depth and be established since the beginning of the winter, in order to eliminate short term influences of the air and surface (Hoelzle et al., 1999; Schoeneich, 2011). With WEqT lower than -3°C, permafrost is probable, permafrost is unlikely if the temperature is higher than -2°C. Between -2°C and -3°C only marginal permafrost in thick active layers may eventually occur (Hoelzle et al., 1999; Schoeneich, 2011). During the melting period, the so called ‘zero curtain period’, the temperature is constant around 0°C.
2.2.2.2 Geomorphological indicators
Not only temperature data may indicate permafrost. Rock glaciers are commonly used as geomorphological indicator of permafrost presence (e.g. Deluigi et al., 2017). If active or inactive rock glaciers are present, permafrost can be expected at the slope. The presence of relict rock glaciers suggests the absence and degradation of permafrost conditions (Deluigi et al., 2017). Another possible indicator of permafrost is solifluction. However, this one is not so clear as the presence of rock glaciers.
Only if plug-like flow exists, the presence of permafrost can certainly be expected. The characteristics
6 of the other kinds of solifluction will differ between cold, warm or non-permafrost zones. For instance, in areas with cold permafrost, the surface velocity will rarely exceed 5 cm year-1. This velocity may rise and reach a maximum (> 10 cm year-1) towards regions underlain by warmer permafrost or those lacking permafrost. Furthermore, the maximum depth of the movement, which is reflected by the minimum frontal height of the solifluction lobe, may decrease toward warmer permafrost. In non-permafrost areas, this depth ranges from a few centimetres up to 50 cm (Matsuoka, 2001). However, there seems to be insignificant difference in solifluction between zones of warm permafrost and seasonal frost (Matsuoka, 2001).
2.2.2.3 Geophysical measurements as indicator
Last, but not least, as geophysical properties of the soil alter significantly with a phase change of water, those properties can also be used as an indicator for permafrost presence (Harris et al., 2009). Based on characteristics as the electrical resistance, propagation speed of georadar waves and velocity of seismic waves, frozen ground can be distinguished from unfrozen ground (Harris et al., 2009).
Commonly used methods are Vertical Electrical Soundings (VES) (e.g. Lambiel & Pieracci, 2008), Electrical Resistivity Tomography (ERT) (e.g. Kenner et al., 2017; Hauck et al., 2003; Kneisel et al., 2000, Scapozza et al., 2011), seismic refraction tomography (SRT) (e.g. Kenner et al., 2017; Hausmann et al., 2007) and Ground Penetrating Radar (GPR) (e.g. Hausmann et al., 2007). SRT may be used complementary to resistivity techniques, to divide air-filled porous layers from ice-filled porous layers (Kenner et al., 2017). GPR is commonly used in high latitude regions, for instance Alaska, but can also be used to determine the internal structure of a rock glacier (Hausmann et al., 2007). Table 1 gives an overview with some typical values for every material and method. The values are based on the earlier mentioned researches. It is important to keep in mind, that the exact values may differ depending on the host materials, ice content, temperature of the ice and impurities (Hauck & Vonder Mühll, 2003). In Figure 2 the resistivity of different rocks, soils and minerals are listed (Loke, 2002).
VES ERT SRT GPR
Debris / open-work
surface layer 6 – 300 kΩ.m 10 – 100 kΩ.m 950 m.s-1
Ice rich permafrost 8-300 kΩ.m > 50 kΩ.m 3300 m.s-1 0,14 – 0,15 m.ns-1
Bedrock < 10 kΩ.m
Depend on the material e.g.
Dolomites: < 8 kΩ.m Serpentinite 18 – 25 kΩ.m
4100 m.s-1
Unfrozen materials/
low ice content 1 – 20 kΩ.m < 10 kΩ.m 1000 – 2000
m.s-1
Table 1: overview of some typical geophysical outcomes (Based on: Lambiel & Pieracci, 2008 (VES); Hauck et al., 2003 (ERT); Scapozza et al., 2011 (ERT); Kenner et al., 2017 (ERT, SRT);
Hausmann et al., 2007 (GPR))
7 Figure 2: resistivity of soil, rock and minerals (source: Loke, 2002)
2.2.2.4 Distribution of mountain permafrost on talus slopes
Within the belt of discontinuous permafrost, which is above the 2200 – 2300 m.a.s.l, it is common to measure highest resistivity and coldest temperatures in the lower part of the debris accumulation. In the higher parts, temperatures increase and resistivity decreases (Lambiel & Pieracci, 2008). This suggest a higher probability of permafrost in the lower parts and a warming or absence of permafrost in the higher parts of a talus slope. Different case studies confirm this unexpected contrast (e.g. Pieracci et al., 2008; Scapozza et al., 2011; Delaloye & Lambiel, 2005). The indicated theoretical distribution of permafrost in a talus slope is the result of three main controlling factors: grain size, presence of snow and in some cases a chimney effect (Lambiel & Pieracci, 2008; Pieracci et al., 2008; Scapozza et al., 2011). These are complex processes, so the exact impact of every single process is difficult to disentangle.
In coarse, blocky materials, which dominate at the foot of the talus slope, a negative temperature anomaly exists (e.g. Harris & Pederson, 1998; Lambiel & Pieracci, 2008). There are different theories to explain this temperature anomaly: the Balch effect, the water retention capacity, the continuous air exchange with the atmosphere and the chimney effect. The Balch effect is based on the fact that cold air is denser and displace warmer air. This process can only occur where there are large connecting
8 spaces between the blocks (Harris & Pedersen, 1998). Additionally, a continuous air exchange, which is possible where there is no continuous snow cover, will favour ground cooling in blocky areas (Lambiel
& Pieracci, 2008). The difference in water retention capacity is also a possible explanation. Coarse grained soils have a lower water retention capacity, which result in less water. Since there is a smaller amount of water, less latent heat will be released while freezing and there will be a faster cool down during autumn (Kenner et al., 2017).
The chimney effect, a mechanism of air circulation, is another explanation. This process does not influence all talus slopes. The efficiency depends on the macro-porosity of the soil, the temperature contrast between the inside and outside air, and the snow depth (Phillips et al., 2009; Pieracci et al., 2008). In small grained formations, or formations completely sealed with ice, circulation of air is far more difficult, and the chimney effect may be less important, or non-existent (Pieracci et al., 2008). In Figure 3 the chimney effect is visualized. It exists of two different phases: winter and summer. In winter, the air inside the slope is warmer, and thus lighter, than outside. This air rises inside the slope and expel at the top. In the lower parts, this causes an aspiration of cold air into the talus. In summer the process reverses and there will be a gravitational discharge of cold, heavy air at the base of the talus slope. This process causes an overcooling at the base of the slope and a positive temperature anomaly in the upper parts.
In winter, the ascent of warm air, can create a basal melting of the snow cover and eventually snowmelt windows are formed (Delaloye & Lambiel, 2005).
Figure 3: The chimney effect (Source: Wicky & Hauck, 2016)
A last factor which may explain the existence of permafrost in the lower parts of the talus slope is the snow distribution. Usually as a result of avalanches or the chimney effect, snow will be redistributed.
Long lasting snow deposits can be found at the bases of talus slopes (Pieracci et al., 2008; Kenner et al., 2017; Lambiel & Pieracci, 2008; Millar et al., 2014). This will result in a delayed snowmelt and, because of the insulating properties of snow, a relative cooling of the ground. Kenner et al. (2017) hypothesize that this ‘insulation theory’ has a minor influence on the presence of permafrost, and that the thermal characteristics of the ground are more important.
9 2.2.3 Mountain permafrost maps and modelling
Since permafrost is invisible, this makes it very difficult to get to know the exact distribution. Boreholes and geophysical data can tell us if permafrost is present, but only at the site itself. To get an area covering map, models are developed (e.g. Deluigi et al., 2017). This models are based on topoclimatic factors and case studies. Mostly they are calibrated for a specific site or regions and do not cover the whole Alpine region. The Alpine Permafrost Index Map (APIM) is the first map that covers the whole European Alps. This map is created by Boeckli et al. (2012) and can be found as a kml-file or downloaded as a georeferenced png-format. The index is an indicator for the probability for permafrost occurrence, but cannot allocate the extent and thickness of the permafrost as this depend on various local and regional processes. A gradation between ‘permafrost in nearly all conditions’ and ‘permafrost only in very favourable conditions’ is made. These conditions refer to the topographical and ground characteristics. For instance: permafrost will not be expected in very fine materials or solid rocks (Boeckli et al., 2012). The map of potential permafrost distribution is another map, which concentrate on the distribution in Switzerland. It incorporates different parameters such as, altitude and solar radiation (www.bafu.admin.ch, 11/5/2018).
2.2.4 Permafrost degradation
Permafrost is a thermal state and will therefore be influenced by the climate and climate change. The thermal evolution of permafrost depends on air temperature and snow cover. An increase of sensible heat flux can result in a general warming of ground temperatures, which result in a degradation of permafrost (Dobinski, 2011; Deluigi et al., 2017). Not only a decrease in the extent of permafrost, but also an increase in the temperatures, will occur (Dobinski, 2011). Figure 4 shows the degradation processes in a permafrost body.
10 Figure 4: Permafrost degradation (Source: Dobinski, 2011)
As a result of this degradation, there will be more temperate permafrost areas (Gruber & Haeberli, 2009) and the altitudinal distribution of frost types and permafrost will shift towards higher regions. Sporadic permafrost disappears from the middle altitudes and the lower permafrost limit shift towards higher altitudes, where continuous permafrost is replaced by discontinuous permafrost (Schoeneich et al., 2011). The thermal regime of permafrost also depends on the snow cover, which act as an insulator.
Most climate models predict an increase of winter precipitation (Schoeneich et al., 2011). Depending of the trend in timing and the thickness of the snowpack, this will result in warming or rather a cooling. An early or thick snow cover prevent cooling during the winter, but a late melt prevents warming in the spring. Snow cover will only impact in areas of moderated slopes, where the snow cover can develop and stay (Schoeneich et al., 2011). The last decades, there has been an increasing interest in permafrost research. This resulted in an augmentation of monitoring sites and research. An example is the Swiss Permafrost Monitoring Network (PERMOS), which has as main goal to document the state and temporal variations of permafrost in the Swiss Alps on a long-term basis (Vonder Mühll et al., 2008).
All the observed elements (e.g. MAGST, borehole temperatures and ERT) showed a warming trend since the year 2009. This trend is likely to be a cumulative effect of continuously warm climate conditions (Nöztzli et al., 2016).
11 2.3 Geomorphological processes and dynamics on talus slopes as impacted by permafrost degradation
This paragraph provides a non-exclusive overview of recent research about the impacts of permafrost degradation on geomorphological processes. The focus will be on geomorphological processes that can occur on talus slopes. Harris et al. (2008) identifies the change in active layer thickness as a key response of permafrost on climate change. The thickness of this layer depends on interannual variations in temperature. This is also stated by other researcher and case studies (e.g. Guo & Wang, 2017; Deline et al., 2015). The rate of increase, and in general the influence of a changing temperature on the geomorphology in periglacial environments, depend on the ice content. In ice-rich permafrost, the latent heat released during thawing, will reduce the thermal conductivity. As a result, there may be a delayed warming. Therefore, one could say that ice-rich permafrost will be less sensitive to climate change, from a thermal point of view (Deline et al., 2015; Gruber & Haeberli, 2005; Schoeneich et al., 2011; Phillips et al., 2009). Furthermore, the altitudinal range, with the highest frequency of freeze-thaw cycles is likely to move to higher altitudes. As a result, the geomorphological activities in the middle altitudes may decrease. Contrary, an increase is expected at the lower limit of the continuous permafrost belt (Schoeneich et al., 2011). Most geomorphological processes with permafrost occurrence are connected to freeze-thaw cycles and active layer thickness. Therefore, the changes mentioned above, together with the increasing temperatures of the permafrost, may influence the nature, intensity and frequency of these processes (Harris et al., 2008; Dobinski, 2011; Deluigi et al., 2017).
Table 2 provides an overview of possible geomorphic reactions on climate change, related to permafrost degradation. One of the possible geomorphological processes reactions is thermokarst. The loss of ice can lead to thermokarst phenomena (Schoeneich et al., 2011). Thermokarst phenomena may tell something about the history and formation of the permafrost. Kenner et al. (2017) identified, for instance, a thermokarst depression in a lobate structure on the Flüelapass. These observations were used to reconstruct the history and origin of the geomorphology and the presented permafrost. Another possible influence is the change in solifluction rate. If there is no frozen layer left, the soil drains and the solifluction will stop (Schoeneich et al., 2011). However, there was no empirical evidence of this phenomenon on the different case studies in action 5.3 from permaNET (Lieb & Kellerer-Pirklbauer, 2011). Rock glaciers has different reaction on permafrost degradation. In general, an increased creep rate can be observed. However, if most of the ice is melted out, the creep rate could decrease, because of the increased friction and eventually they may collapse (Harris et al., 2009). Other permafrost creep phenomena will also show a non-linear increase in velocity because of the decrease in viscosity of ice (Deline et al., 2015). At altitudes with an increased frequency of freeze-thaw cycles, the frost weathering and talus production may intensify. An increased instability of steep bedrocks, and following rock falls, may be the result. Those rock falls may impact lower situated talus slopes (Schoeneich et al., 2011).
Furthermore, the talus production provides a higher amount of available, erodible material for debris flows. Together with the changing precipitation regimes, this may lead to an increasing frequency of debris flows (Marchi et al., 2009). However, if there is no permafrost less, the ground drains and the threshold value of precipitation to start a debris flow may increase (van Steijn, 1996)
12
Geomorphological process Examples of studies
Changes in frost weathering and talus production Kenner et al., 2017 Stoffel & Huggel, 2012
Thermokarst Kenner et al., 2017
Changes in the rate of rock glacier creep
Ikeda & Matsuoka, 2002 Bodin et al., 2017 Giaccone et al., 2016 Changes in solifluction rates Matsuoka et al., 2005 Changes in the frequency and magnitude of mass
movement events
Stoffel & Huggel, 2012 (rock fall) Marchi et al., 2009 (debris flow)
Changes in the volume of unstable materials
Stoffel & Huggel, 2012 Ravanel & Deline, 2011
Table 2: Possible influences of permafrost degradation on geomorphological processes 3. STUDY OBJECTIVES
The purpose of this study is to understand the influence of permafrost distribution on talus slope geomorphology. The key question of this research is ‘How does permafrost distribution influence the talus slope geomorphology on the Col du Sanetsch, Switzerland?’. To accomplish this, we defined some more specific objectives: First of all, a geomorphological mapping is made. In addition to this, measurements of talus slope dynamics, ground surface temperature and the current permafrost distribution are done. Afterwards, we will try to relate those factors to explain the permafrost distribution and related processes. In the last phase the geomorphological and permafrost mapping are combined to answer the research question. Out of the literature review and the research objects, some hypothesis are set to be tested. First off all permafrost is expected in the rock glacier, the landslide and eventually in blocky surface layers along the talus slope. Furthermore, we assume that the displacement of the landslide will be related to the distribution of permafrost and the related factors such as GST and moisture content.
4. STUDY SITE
The case study that will be examined is Col du Sanetsch (46°20’22.60” N, 7°18’41.02” E), located in the western Swiss Alps, at the northern side of the Rhone valley, nearby Sion (Commune Savièse – VS) (Map 1). The slope is oriented towards the northwest with an elevation ranging from 2340 to 2700 m.a.s.l. The talus slope, which mainly consist of limestones (Federal Office of Topography, SGTK, 2012), is located at the base of the rock wall between the Arpelistock (3036 m.a.s.l) and Arpelihore (2921 m.a.s.l) in the Sanetsch – Wildhorn Massif. In the south, the border of the talus slope is defined by the highly erodible badlands. Below the talus slope Late – Pleistocene, Holocene moraine deposits can be found (Federal Office of Topography, 2017). The average inclination of the talus slope is
13 approximately 30 – 40 °, with a basal concavity. In the upper region slopes up to 45° can be found, while at the base the inclination approaches 25°. The 0°C isotherm in this area is situated around 2300 m.a.s.l (Lambiel, 2006). In this context, the talus slope is located within the belt of discontinues permafrost (Scapozza et al. 2011). A series of interesting landforms, such as a landslide, rock glacier and debris flows can be found on the site (Map 2). The tong-shaped rock glacier is formed by two convex ridges and covered with a mix of medium sized to coarse grained boulders. Its present suggest the existence of discontinues permafrost. The landslide is covered by a mix of soil and small to large boulders. To determine the surface velocity and deformation, a network of measuring points was installed in 2011.
Since 2014, continuous ground surface temperature (GST) data is available. However, there is no weather station where long term temperature and precipitation data can be obtained. The weather station ‘Les Diablerets’ is located at 6 km, but in a different context, at 2964 m.a.s.l. on a glacier, and without long term data. The weather station in Sion, 14 km to the south, is located in the valley at 482 m.a.s.l. and is therefore also not completely representative for the study area. At this station the mean annual precipitation from the period 1981 – 2010 is 603 mm (MeteoSwiss, 2016). The mean annual temperature is 10,1°C. Figure 5 visualise the general meteorological context.
Figure 5: General meteorological context: long-term means of monthly mean temperature, monthly maximum and minimum temperature as well as monthly precipitation sums in Sion (MeteoSwiss, 2016)
Map 1: Location study site: Col du Sanetsch
Map 2: Location talus slope and different landforms
46°21’00” N
17°18’29” E 7°18’55” E
46°20’08” N 46°20’34” N
7°18’03” E 7°17’36” E
5. MATERIAL AND METHODS
The purpose of this research is to understand talus slope geomorphology, and more specific the geomorphology of the landslide in relation to the permafrost distribution. This chapters provides an overview of the data and the applied methods. Most of the data was collected during a field work campaign in August 2018. The gathered data can be divide into three groups: geomorphological, geophysical and meteorological data. The geomorphological data exist of UAV (Unmanned Arial Vehicle) photographs used to build a Digital Elevation Model (DEM), field observations and a geodetic network to measure the surface velocity. The temperature data exist of hourly measurements of the ground temperature. Those different data sources (Table 3) are brought together to analyse eventual relationships and explain the talus slope geomorphology.
16
Description Period
DEM
LiDAR – 2m – CH1903+/ LV95 2013
UAV imagery – 6,6 cm – CH1903+/ LV03 Summer 2018
UAV imagery – 10 cm – CH1903+/ LV03 Summer 2017
Differential GNSS
27 points + 4 control points on landslide – CH1903+/ LV03 2011 – 2018 (every year)
Meteorological data
Tloggers_DebrisFlow 01/09/2016 – 12/08/2017 (every hour)
Tloggers_RockGlacier 01/09/2016 – 23/08/2018 (every hour)
Tloggers_TalusSlope 12/08/2017 – 1/07/2018 (every hour)
Tloggers_Landslide (PERMOS) 18/08/2014 - 18/08/2017 (every hour)
Temperature and precipitation data Sion 01/2011 – 11/2018 (every month) Geophysical data
Lateral and vertical ERT profiles on rock glacier August 2016 Lateral and vertical ERT profiles on three locations along
landslide
August 2018
Two Schlumberger VES profiles in blocky surface layer on the talus slope
August 2018
Schlumberger and Wenner VES profile on landslide August 2018 Potential permafrost distribution maps Alpine Permafrost Index Map (APIM) Boeckli et al., 2012 2012 Map of Potential permafrost distribution (FOEN, 2005) 2005 Table 3: Overview data
5.1 Topographic survey based on field observations and UAV
The first important purpose is the creation of a geomorphological map (Map 3). The different features were mapped in Qgis 3.4.5. The visualisation is done in Arcmap 10.6.1. The mapping itself is a combination of two different methods: Field observations and a DEM created by photographs from an UAV. Articles of Hendrickx et al. (2019) and Westoby et al. (2012) provide more detailed information about the principles of this so-called, ‘Structure-from-Motion’ photogrammetry.
17 5.1.1 Field observations
Equipped with a handhold GPS (Garmin eTrex 20), the landslide and talus slope where explored. During this field campaign, waypoints of interesting features, e.g. ridges, snow patches, proofs of sorting etc.
where mapped. A description was written down and additionally pictures and sketches were made. In order to make it possible to compare pictures and get an idea of the grain sizes, the same pen is visible on every picture. This information is complementary to the DEM and provides information to interpret or clarify observations on the DEM.
5.1.2 Data acquisition with UAV
5.1.2.1 UAV
To obtain a high resolution DEM of the study site, we did an UAV survey. The photographs were taken using a 16MP Panasonic Lumix DMC-GM5 with a 12-32 mm F3.5 – 5.6 lens. The focal length was fixed on 20 mm, but there was also a flight on 12 mm. The shutter speed varied between 1/400 and 1/800 second, the ISO between 400 and 800 depending on the light conditions. The camera was fixed on a custom-made Hexacopter DJI F550 with a Pixhawk flight controller. The flight speed was 4 – 6 m.s-1. To establish sufficient overlap between the pictures, the acquisition interval was set on 1 second. The ground sampling density was 1,64 cm.pixels-1. Eleven flights were flown, from which eight were used to create the DEM.
The preparation of this survey consist of two different tasks. The first one to be done was the flight planning. Therefore, flight lines were drawn in QGIS. These flight lines are parallel to the contour lines of the area. In this way, the flight height, which is approximately 90 m, is adjusted to the topography.
The flight lines were translated into a waypoint-file, with flight direction, using a Python script. In the field, these files are imported in the program ‘APM Mission Planner’ and written to the UAV. The other preparation task is done in the field. Six ground control points (GCP’s) are distributed over the talus slope and landslide. They are needed to reference the model to a real-world system. It is important that these GCP’s are easy to identify on the pictures, so they need a strong contrast and clear centre (Figure 6) (Hendrickx et al., 2019). On the lower places we used plates (Figure 6a). These ground control points are fixed on an iron pin, which is drilled on a stable rock. The clear centre of these plates is an advantage, but the weight of 1 kg makes it inefficient to use them on higher locations. Furthermore, in the upper part of the talus slope, there are no stable rocks where we could install the plates. As a result we used cloths (Figure 6b). The centre of these GCPs is less clear, but we can carry them in our backpacks and they weigh less. Since the accuracy of the GCP’s defines the accuracy of the final DEM, it is important that the GCP’s are measured with a high precision. We used a RTK-GNSS to perform these measurements.
5.1.3 Real-Time Kinematic GNSS (RTK-GNSS)
An RTK-GNSS improves the accuracy of the measurements by using two different receivers. One fixed receiver, which is used as reference, and assumed to be motionless and a second receiver, the rover,
18 which is mobile. The first receiver determines the difference between the received position and its’
known position. These differences will be used to correct the measurements from the second receiver.
The rover, second receiver, measures the position of the different points on the landslide. These receivers communicated permanently and the reference receiver sent the correction immediately to the rover (Wee, 2018). To control the accuracy of the measurements, the average of five to ten measurements is retained (Wee, 2018). This method allows to reach a precision up to 3 cm. However, these final measurements need to be corrected. The first correction converts the altimetry data from the Bessel ellipsoidal reference to the geoidal reference. This is done by the Swisstopo platform (https://www.swisstopo.admin.ch/fr/cartes-donnees-en-ligne/calculation-services/reframe.html). In addition, by measuring a geodetic point in the area, we found out that the location of the base station, the first receiver, is not exact. Therefore, a second correction is needed. To get the correct absolute coordinates following corrections are done: easting (-1,438), northing (-1,117) and the altitude (+2,976).
Figure 6: Installation and measurement of GCPs. a) a fixed GCP, b) a ‘cloth’ GCP
5.1.4 Data processing in Agisoft PhotoScan
The obtained UAV photographs were imported into the program Agisoft PhotoScan 1.2.6 to compose the final DEM. This program uses the principles of the Structure-from-Motion photogrammetry, a technic which builds 3D structures from a series of overlapping images. Contrary to the conventional photogrammetry, there is no need to specify a priori network of targets with known 3D-positions. To estimate the camera positions and object coordinates, matching features on multiple images are tracked. More information about this technic can be found in the article of Westoby et al. (2012). The workflow followed to create the DEM can be found in Figure 7a. On the first try, one out of four photographs were imported. Afterwards, we added more photos in the zones were there where gaps
a)
b)
19 (Figure 7b) or where the alignment did not work. We only managed to align eight of the eleven flights.
As a result, the upper part of the talus slope is not covered. Furthermore, there is also a small data gap in the southern part of the landslide. Several runs were executed to investigated the problem. These runs varied in the amount of photographs they included or the settings used. There were tries in which, all photographs were imported within one single chunks and runs with one chunk for every flight or even flight line. Even though these problems could not be encountered. As we used the same methodology on the whole area, the question rise what the difference is between the upper region and the other areas. Between the flight lines, there is not only a horizontal displacement, but the UAV also ascends.
We presume that the overlap problem can be related to this ascent and has its origin in the steepness of the area. Too much ascending, resulting from the steepness in the area, between the flight lines can cause an insufficient overlap. Adding some flight lines is something which can be tried during the next field work.
Following Hendrickx et al. (2019) we run everything on medium quality. In this article, the performance of ten medium and high quality run are compared. The differences were very small, but the variability for the medium quality run is more predictable. Furthermore, the variations in DEM and curvature also show a better performance. Considering those facts and the computational effort calculation time, a medium quality run is more efficient. In the final model the GCPs have a RMS error of 10,9 cm. The used coordinate system is CH1903+/ LV95 (EPSG: 2056). The resolution of the achieved DEM is 6,6 cm. In Table 4 the technical details of our UAV survey and the resulting DEM can be found.
Figure 7: a) Workflow in Agisoft to build a DEM based on Hendrickx et al. (2019). b) A model after the first alignment. In the upper zone, different gaps are visual, here we added more photographs.
a) b)
20
Number of flights 8 Number of points 1 779 247
Flying altitude 90 m Ground sampling density 1,64 cm.pixel-1
Number of images 890 RMS of GCPs 10,9 cm
Covered area 0,408 km² RMS of check points 84,9 cm
Coordinate system CH1903+/LV95 (EPSG: 2056)
Resolution DEM 6,6 cm
Table 4: Technical table UAV survey
5.1.5 Geomorphological mapping based on DEM interpretation
The geomorphological map is based upon the DEM. As we did not want to visualise every single small rock, we first applied a moving window (15x15) on the DEM. In the resulting ‘smoothed DEM’, every cell represents the mean height of a square of 99 x 99 cm (Arcmap > focal statistics). From this layer we derived aspect, slope, surface roughness and tangential curvature (Qgis > Saga). Each of these layers give some extra insight in the geomorphology of the landslide and talus slope. By combining the layers, different features were highlighted and identified. The northern part of the talus slope was not covered by our DEM, in this area the DEM from 2017, made by Hanne Hendrickx was used. In the area above the landslide, where there is no DEM from 2017 or 2018, LiDAR data was used to close the data gap. It is important to keep this in mind as the parts mapped on the LiDAR data are less detailed e.g. on the rock glacier it was not possible to map the small scale geomorphology. A hillshade and contour lines derived from LiDAR data provide the background of the map. The focus of the geomorphological map lies within the explanation of the different processes and displacement. This formed the starting question to analyse the data and map all the different landforms. Source areas, accumulation zones and proves of (earlier) displacement are visualised.
5.1.6 Defining topographic roughness
Topographic roughness is calculated as it can be used as a proxy for grain size (Otto et al., 2012).
Grohmann et al. (2011) compares six different methods to define topographic roughness within the field of geomorphology. In this research, we choose the standard deviation of residual topography (SDrestopo).
This method is also applied in other researches in mountain areas (e.g. Otto et al., 2012; Haneberg et al., 2005) and filters out the large-scale topography, which is an advantage. Furthermore, it is derived from the DEM itself, so small errors in the DEM will not be enhanced, as can be the case if we use a derivative such as the standard deviation of slope (Grohmann et al., 2011; Haneberg et al., 2005).
Finally, it is also easy to produce the SDrestopo using raster GIS algebra (Haneberg et al., 2005).
Calculating the roughness index is done in three different steps (Figure 8a). In the first step, a moving- window of 15 x 15 is applied on the high resolution DEM. This results in a ‘smoothed DEM’ in which every cell gives the average height over approximately 1 m². This resolution was chosen as a compromise. If we used a smaller resolution, more big boulders would be seen as normal topography and only the edges would be highlighted. By using a higher resolution small depressions and ridges could be smoothed, which would result in a topographic roughness value higher than the effective
