RECENT FAULTING AND ACTIVE SHORTENING OF THE MIDDLE ATLAS
MOUNTAINS, MOROCCO, WITHIN THE DIFFUSE AFRICAN-EURASIAN
PLATE BOUNDARY
_______________________________________
A Thesis presented to the Faculty of the Graduate School
University of Missouri-Columbia
_______________________________________________________
In Partial Fulfillment
Of the Requirements for the Degree
Master of Science
_____________________________________________________
by
Michael Rigby
Dr. Francisco Gomez, Thesis Supervisor
MAY 2008
The undersigned, appointed by the Dean of the Graduate School, have
examined the thesis entitled:
RECENT FAULTING AND ACTIVE SHORTENING OF THE MIDDLE ATLAS
MOUNTAINS, MOROCCO, WITHIN THE DIFFUSE AFRICAN-EURASIAN
PLATE BOUNDARY
Presented by Michael Rigby
A candidate for the degree of Master of Science
And hereby certify that in their opinion it is worthy of acceptance.
Dr. Francisco Gomez
Dr. Robert Bauer
Dr. Brent Rosenblad
ACKNOWLEDGEMENTS
I would like to thank my thesis advisor, Dr. Francisco Gomez, for all of
his help throughout the research process. I would also like to thank my thesis
committee members, Dr. Robert Bauer and Dr. Brent Rosenblad, for their
assistance during the writing process. Also, I want to thank Yousuef Hahou,
Zakir Abdelali, Nacer Jabour, and Abdelali Fadil for their assistance during
the completion of field work. I would also like to thank the National Science
Foundation of the funding of this project (EAR-0408661) and Dr. Rick
Allmendinger for the use of his structural programs. Lastly, I would like to
thank my family and friends for their continuous support.
ii
TABLE OF CONTENTS
ACKNOWLEDGEMENTS................................................................................ ii
LIST OF FIGURES .......................................................................................... v
LIST OF TABLES .......................................................................................... vii
ABSTRACT ...................................................................................................viii
CHAPTER 1 INTRODUCTION AND TECTONIC BACKGROUND..................1
Introduction ..................................................................................................1
Objectives and Justification ......................................................................4
Tectonic Setting ...........................................................................................5
Tectonic Evolution of the Atlas Mountains ...................................................6
Triassic through Early Tertiary..................................................................6
Cenozoic Tectonics ................................................................................10
CHAPTER 2 - MIDDLE ATLAS STRUCTURE ..............................................19
Introduction ................................................................................................19
Geologic Mapping of the Folded Middle Atlas............................................22
Stratigraphy ............................................................................................23
Triassic ...............................................................................................23
Lower Liassic ......................................................................................23
Middle Liassic .....................................................................................24
Cretaceous .........................................................................................24
Neogene .............................................................................................24
Quaternary..........................................................................................24
Remote Sensing Analysis and Interpretation..........................................24
Data Types .........................................................................................24
Analyses .............................................................................................27
Stereo Photogrammetry ..................................................................27
Multi-spectral image processing and interpretation .........................27
Geologic Map of the Folded Middle Atlas ...............................................33
Structural Analyses ....................................................................................35
Balanced Cross-Section of the Folded Middle Atlas ..................................40
Discussion..................................................................................................42
CHAPTER 3 ACTIVE TECTONISM IN THE FOLDED MIDDLE ATLAS
MOUNTAINS .................................................................................................45
Longitudinal Stream and Stream Terrace Profiles .....................................46
Quaternary Fault Scarps ............................................................................49
Microtopographic Mapping .....................................................................49
Radiocarbon Age Dating ........................................................................57
Fault Scarp Degradation Modeling .........................................................58
Single Event Fault Scarp Degradation Model .....................................59
Composite Fault Scarp Degradation Model ........................................66
iii
Discussion..................................................................................................70
Conclusion .................................................................................................72
REFERENCES CITED ..................................................................................74
iv
LIST OF FIGURES
Figure
Page
1.1. Simplified tectonic map of the western Mediterranean. ...........................2
1.2. Topography map of the western Mediterranean. .....................................3
1.3. Tectonic evolution of the western Mediterranean from 200 - 30 Ma. .......7
1.4. Paleotectonic reconstructions of the western Mediteranean..................12
1.5. GPS velocity vector map. ......................................................................15
1.6. Seismicity Map. ....................................................................................16
2.1. Geologic map of the Middle Atlas Mountains..........................................20
2.2. ASTER band combination image with bands 7-3-1. ..............................29
2.3. Shortwave infrared spectral reflectance curves .....................................30
2.4. ASTER image composed of principal components 4, 5, and 6..............31
2.5. Geologic map of the Middle Atlas study area. .......................................34
2.6. Bedding Measurements.........................................................................36
2.7. Exposure of the Ait Oufella Fault in a road cut. .....................................38
2.8. Fault kinematic data ..............................................................................39
2.9. Balanced cross section..........................................................................41
3.1 Longitudinal Stream and terrace profiles ...............................................48
3.2. Field photograph of the composite scarp...............................................51
3.3. Field photograph of the singe event scarp.............................................52
3.4. 1 meter digital elevation model of the composite scarp. ........................54
3.5. 0.5 meter digital elevation model of the single event fault scarp............56
3.6. Single event fault scarp profile 1............................................................62
v
3.7. Single event fault scarp profile 2............................................................63
3.8. Single event fault scarp profile 3............................................................64
3.9. Single event fault scarp profile 4............................................................65
3.10. Root mean square analysis for D variable. ..........................................68
3.11. Composite fault scarp degradation model ...........................................69
vi
LIST OF TABLES
Table 2.1 Performance parameters for the ASTER radiometer. ...................26
Table 3.1 Radiocarbon age data ..................................................................57
vii
RECENT FAULTING AND ACTIVE SHORTENING OF THE
MIDDLE ATLAS MOUNTAINS, MOROCCO, WITHIN THE
DIFFUSE AFRICAN-EURASIAN PLATE BOUNDARY
Michael Rigby
Dr. Francisco Gomez
Thesis Supervisor
______________________________________________________________
Abstract
The intracontinental Atlas Mountains of northwest Africa are prominent
tectonic elements in the broader, diffuse plate boundary zone between the
converging African and Eurasian plates in the western Mediterranean.
Cenozoic geology, as well as present-day seismicity, document co-eval
tectonism across a zone more than 500 kilometers in width. Hence,
quantifying the rates of deformation in the Atlas Mountains is necessary in
understanding the distribution of crustal deformation across the entire plate
boundary zone. This study focuses on constraining the recent shortening rate
across the NE-SW striking Middle Atlas Mountains, which are obliquely
oriented to the NNW-SSE direction of Late Cenozoic plate convergence.
Both strike-slip and contractional structures are exhibited with apparent slip
partitioning: folding and thrusting is concentrated in the Folded Middle Atlas,
whereas strike-slip dominates in the Tabular Middle Atlas. In the
southwestern part of the Folded Middle Atlas, fault scarps of Quaternary
alluvium, including a 4 meter composite scarp and a 1 meter single event
scarp, attest to recent faulting along the mountain front. Detailed topographic
viii
mapping of the scarps provides a basis for geomorphic analysis and
degradation modeling. Furthermore, the reconstruction of longitudinal stream
terrace profiles helps constrain a long term deformation history. Radiocarbon
dates provide age constraints on the faulted surfaces. To place these active
tectonic observations in a larger context, the fault and fold geometry has been
assessed by completing a 12.5 km structural transect across the frontal
thrust, providing basis for the construction of a balanced cross-section. By
combining the structural geometry with the uplift rate, a minimum estimate of
the rate of horizontal shortening in the Middle Atlas is evaluated. Degradation
modeling results suggest the Middle Atlas Mountains are shortening at a rate
of 0.03 – 0.05 mm/yr, which is lower than the geodetic estimates. This implies
that shortening at a regional scale is being accommodated by an alternative
mechanism such as strike-slip motion.
ix
Chapter 1 Introduction and Tectonic Background
Introduction
The Atlas Mountains of northwest Africa are prominent tectonic
elements in the broader, diffuse plate boundary zone between the converging
African and Eurasian plates in the western Mediterranean. Cenozoic
geology, as well as present-day seismicity, document co-eval tectonism
across a zone more than 500 kilometers in width extending from southern
Iberia to the edge of the Saharan Platform (Figure 1.1). Hence, quantifying
the rates of deformation in the Atlas Mountains is necessary in understanding
the distribution of crustal deformation across the entire plate boundary zone.
In that spirit, this study aims to provide the first published constraints on the
recent shortening rates across the NE-SW striking Middle Atlas Mountains.
In tectonic terms, the Atlas Mountains are considered a prominent
example of an intracontinental mountain belt (e.g., Brede et al., 1992). In
contrast to the adjacent Rif and Tel Mountains of North Africa, which
comprise the corresponding interplate mountain ranges, the Atlas Mountains
are located at a considerable distance into the foreland from their
corresponding collision zone (Figure 1.1).
The Atlas Mountains span approximately 2000 kilometers from the
Atlantic coast of Morocco to the Mediterranean coast of Tunisia, and they
attain elevations greater than 4 kilometers (in comparison with maximum
altitudes of about 2 kilometers in the Rif and Tel Mountains) (Figure 1.2).
1
Figure 1.1. Simplified tectonic map of the western Mediterranean. Regions
surrounding the study area include TMA – Tabular Middle Atlas, HM – High
Moulouya, MB – Missour Basin, and GB – Guercif Basin. The area of Figure
2.1 is outlined by a black box. Figure adapted from Gomez et al. (2000).
2
Figure 1.2. Topography map of the western Mediterranean. The
intracontinental Atlas Mountains attain higher elevations than their
corresponding interplate mountains (i.e. the Rif and the Tel)
3
The main ranges that comprise the Atlas are the High Atlas and Middle Atlas
in Morocco and the Saharan Atlas in Algeria (Figure 1.1). The NE-SW
striking Middle Atlas Mountains are oriented oblique to the general ENEWSW strike of the High and Saharan Atlas.
The region adjacent to the plate boundary, including the Alboran Basin
and the Rif and Betic mountain ranges, has accommodated the majority of
the deformation since the early Miocene (Gomez, 2000). Inland from the
plate boundary, the High Atlas, oriented nearly perpendicular to the NNWSSE maximum Cenozoic plate convergence, represent a region of
compression and have been estimated to have accommodated 17-45% of the
overall plate convergence since the early Miocene (Gomez, 2000).
Objectives and Justification
Throughout the intracontinental Atlas Mountains, present day
convergence rates and structural geometries are not well constrained. The
focus region of this study, the Middle Atlas Mountains, is one key element of
foreland deformation occurring within the African- Eurasian plate boundary.
Evidence for recent faulting along the front of the Middle Atlas foreland fold
and thrust belt provide incentive for a neotectoinc investigation and an
opportunity to incorporate the active tectonic observations with a structural
study. In particular, what are the structural styles in the Middle Atlas
Mountains and what is the total Cenozoic shortening? Are there any along
4
strike variations in structural style or shortening? Also, what is the shortening
rate across the front of the Middle Atlas and how does it compare with recent
geodetic estimates? This study intends to evaluate these questions and their
implications on regional tectonics. Furthermore, these objectives not only
have tectonic implications, but also implications for the regional earthquake
hazard. In order to address these objectives, the structure and recent faulting
along the front of the Middle Atlas is evaluated.
Tectonic Setting
Adjacent to the Atlas ranges are two stable crustal blocks that have
experienced little deformation during the Cenozoic: the Moroccan Meseta
located northwest of the High Atlas and the High Plateau located to the
northeast (Figure 1.1) (e.g. Gomez, 2000). To the south of the Moroccan
Atlas are the Anti-Atlas, which are an exhumed Hercynian (late Paleozoic)
mountain chain (e.g. Africa’s Appalachian equivalent) and the stable Saharan
Plateau. Lying to the east and southeast of the Middle Atlas are the Guercif
Basin and Missour Basin. Along the Mediterranean Coast, north of the Atlas
ranges, are the western most Alpine orogenic belts: the Rif Mountains and the
Tell Mountains along northern Africa and the Betic Mountains of southern
Iberia. The Alpine mountain ranges in this area are divided into internal
allochthonous and external autochthonous zones. The Rif and Betic ranges
form an arcuate mountain belt surrounding the Alboran Basin (Figure 1.1).
5
Inherited geological structures likely influence the styles of deformation
in the Atlas Ranges. The present-day Middle Atlas Mountains reflect multiple
phases of deformation including components of extension, compression, and
strike-slip linked to the opening of the Atlantic Ocean and Tethys Sea in the
early Mesozoic and the reactivation of the failed rift basins during Cenozoic
convergence (e. g. Brede et al., 1992; Gomez, 1996). Hence, to place the
Atlas into a broader context, a summary of the Mesozoic and Cenozoic
tectonic evolution of the region follows.
Tectonic Evolution of the Atlas Mountains
Triassic through Early Tertiary
In the late Paleozoic, the supercontinent, Pangea, began breaking up
and forming rift basins as the plates separated. During this time, sedimentary
basins were formed within rift grabens throughout northern Africa (Brede et
al., 1992). This represents the first marine incursion of the Tethyan Ocean
over the margin of NW Africa, which resulted in the deposition of a carbonate
platform (Ait Brahim, 2002). The basins formed primarily along two axes
oriented at 30º - 45º or NE-SW and 70º or ENE-WSW, which are parallel to
the Atlantic and Mediterranean trends and correspond to the strike of the
future mountain ranges of the High Atlas and Middle Atlas (Figure 1.3) (Kanes
1973; Brede et al., 1992). Oriented similar to the basins, a system of
synsedimentary faults controlled the development of the basins. Both normal
6
Figure 1.3. Tectonic evolution of the western Mediterranean from 200 Ma to
30 Ma. Figure adapted from Brede et al., (1992).
7
and oblique slip is observed on the faults in these basins. The existence of
both normal and strike-slip components of motion on both fault sets
throughout the Triassic suggests a stress field with a least principal
extensional stress oriented at approximately NNW-SSE to NW-SE in
response to the rifting in the central Atlantic and the opening of the NeoTethys Sea to the east (Brede et al., 1992; Ait Brahim, 2002).
Through the Triassic and into the early Liassic (Lower Jurassic), rifting
continued in the basins associated with the High and Middle Atlas ranges.
The stable Moroccan Mesesta separated the Atlantic rifting system from the
Atlas rifting. During the early Jurassic or Liassic, the Atlantic rift system
became dominate and the Atlas rift system was aborted. Once aborted, the
intracontinental basins in NW Africa began to develop. As the central Atlantic
began active oceanic spreading, it initiated a large transform fault zone, the
Newfoundland-Gibraltar fault zone (NGF), due to differential spreading rates
of Africa and Iberia (Figure 1.3A). The northern Atlantic remained in a stage
of continental rifting causing Iberia to spread at a slower rate than Africa.
Uplift due to rifting diminished in northwestern Africa and the tensional forces
from spreading caused regional post-rift thermal subsidence. At this time,
Africa began to move eastward in relation to Iberia due to higher drift
velocities along the sinistral NGF.
Throughout the rest of the Jurassic and into the early Cretaceous,
Africa continued to move eastward in relation to Iberia. The African continent
remained in relative tectonic quiescence; however, as the continent moved
8
further east, the tensional stresses rotated slightly due to sinistral wrenching
along the NGF (Brede et al., 1992). The tensional stresses within the
continent were oriented at about 125º, which initiated normal faulting along
the 30 º - 45 º striking normal faults and transtensional movement along the
70º striking faults (Figure 1.3) (Brede, 1992).
In the Cretaceous, seafloor spreading began in the north Atlantic, and
Iberia began to slide eastward in relation to Africa causing dextral wrenching
on the NGF (Figure 1.3B) (Brede et al., 1992). In addition, a large fault splay
within the Newfoundland-Gibraltar fault zone, southwest of Iberia, propagated
to form the ALKAPECA microplate (ALKAPECA = Alboran-KabylianPeloritanian-Calabrian, Boulin et al., 1986).
This combined movement
induced dextral slip along the 70º faults and activated the 30º – 45º faults as
oblique reverse faults (Figure 1.3) (Brede et al., 1992). Paleostress analyses
estimate a maximum compressive stress oriented E-W to ESE-WNW, which
is congruent with the induced movement (Ait Brahim, 2002). This represents
the first component of compression within the Atlas; however, this is not
nearly as significant compared to the later compression during the Cenozoic.
From the Triassic through the early Tertiary, the differential spreading
rates of the Iberian and African plates created zones of weakness within the
African continent that would later be reactivated. The break up of Pangea
due to the spreading of the Atlantic Ocean and the opening of the Tethys Sea
created transtensional forces throughout northwestern Africa. The primarily
9
extensional deformation through the Mesozoic differs from the compressive
nature of Cenozoic deformation.
Cenozoic Tectonics
During the late Paleocene and early Eocene, the plates began to move
eastward at similar velocities (Ziegler, 1988; Brede et al., 1992). Later in the
Eocene, overthrusting began on the borders of the Alkapeca microplate and
the emplacement of the Alboren block began in response to the eastward
movement of Iberia and Africa (Dewey et al., 1989; Brede et al., 1992).
These tectonic processes caused the first major uplift in the High Atlas and
Middle Atlas that continued on through the early Oligocene. The movement
along these fault systems at this time was marked by the uplift and
denudation of footwall rocks and the deposition of syntectonic conglomerates.
In the Oligocene, convergence between the continents, linked to the
Alpine Orogeny, continued. The internal zones of the Rif, Betic, and Tell
Mountains were part of an uplifted region corresponding to an area of
thickened crust, at least as thick as 40 kilometers based upon gravity
anomalies (Platt, 1998).
Concurrently, tectonic forces were also transmitted
to inland Africa and uplift continued in the Atlas. However, the uplift occurring
in the Atlas during Eocene to Oligocene time was not the major orogenic
event. According to sedimentological studies of the Atlas region, major uplift
and denudation did not begin until the middle Miocene at the earliest (Gorler,
10
1988). Strata within the Missour Basin records uplift within the eastern High
Atlas and Middle Atlas as occurring from the middle Miocene through the
Quaternary (Gomez et al., 2000).
During the middle Miocene, the African and Iberian Plates not only
jointly moved eastward but they also converged toward each other (Figure
1.4) (Dewey et al., 1989). However, in contrast to overall plate movement,
the thickened crust of the Alboran region began to extend, which dispersed
the internal zones of the Rif, Betic, and Tell Mountains. This movement
initiated the nappe thrusting of the internal zones of the Rif-Betic Mountains
onto the external zones (Brede et al., 1992). At the same time, uplift
increased greatly within the High Atlas and Middle Atlas. The faults oriented
at 70º, corresponding to the High Atlas, experienced a large amount of upthrusting; while the faults oriented at 30º - 45º exhibited sinistral
transpression.
Since the Miocene, NNW-SSE African-Iberian plate convergence
causing thrusting in the High Atlas and transpressive forces within the Middle
Atlas. However, contrary to the Atlas, the thrusting occurring within the Rif
and Betic Mountains has been oriented WSW to WNW and concurrent
extension has been occurring within the Alboran Basin. This suggests forces
acting upon the region other than plate boundary forces. These geodynamic
forces are essential to the interpretation of the deformation occurring in this
diffuse plate boundary.
11
Figure 1.4. Paleotectonic reconstructions of the western Mediteranean
region. White arrows denote the directions of local tectonic transport. White
and black teeth denote inactive and active thrust faults, respectively. The
present day coastline is outlined in all time stages. Figure adapted from
Gomez et al. (2000).
12
Three major hypotheses have been suggested to explain the
configuration of the Rif-Alboran-Betic region: subduction rollback leading to
back-arc spreading (Lonergan, 1997), delamination of an overthickened
lithosphere (Seber, 1996, Platt 1998, Fadil, 2006), and convective removal of
mantle lithosphere (Platt, 1998). According to the subduction rollback
hypothesis, the oceanic lithosphere must be subducting at such a rate and
angle to induce back-arc spreading. In contrast, the delamination and roll
back of overthickened lithosphere induces extension through replacement of
dense lithospheric material with more buoyant asthenosphere. Lastly, the
active removal of mantle lithosphere below the Alboran region would create
high gravitational potential and induce spreading. All three hypotheses do
explain the thinning of the Alboran crust and the existence of extension;
however, new Global Positioning System (GPS) data has provided evidence
to support the roll back of a delaminated subcontinental lithospheric slab
(Fadil et al., 2006).
Based upon GPS data, the present day deformation occurring in the
region is asymmetric and appears to be more indicative of a component of
southward-directed delamination of lithosphere and associated N-S back arc
opening (Fadil et al., 2006). Since the crust of the Rif-Alboran-Betic region is
continental in nature (Platt, 1989), the present day slab is probably the mantle
part of the continental lithosphere, which has become detached from the crust
possibly due to the pull of an old slab (Faccenna et al., 2004, Fadil et al.,
2006). This is further supported by three-dimensional analysis of the upper
13
mantle where very low Pn velocities exist in the uppermost mantle below the
Alboran block and are underlain by a more dense higher velocity zone
(Seber, 1996; Calvert, 2000). The low velocity zone suggests the presence
of a more viscous material underlain by the detached higher velocity
lithosperic mantle. These geodynamic processes incorporated with the plate
boundary forces are responsible for the style of deformation occurring in
northern Africa and the active tectonism in Atlas ranges.
The active setting of the Atlas Mountains can be partially evaluated
through GPS measurements. The GPS measurements give an estimate of
an approximately 5 mm/yr convergence rate between the African and
Eurasian plates in the NW-SE direction (Figure 1.5) (Fadil et al., 2006). In
addition as much as 1 mm/yr convergence is occurring within the Middle
Atlas. However, uncertainties at 95% confidence are generally larger than
the velocities. Hence, GPS can, at best, place a maximum constraint on the
convergence across the Middle Atlas. Therefore, geologic estimates of
shortening across the Middle Atlas, which encompass longer time periods,
may be more appropriate.
In addition to GPS, recent seismicity throughout the Atlas Mountains
attests to the active nature of the mountain ranges and regional seismic
hazard (Figure 1.6). The most studied event in the Moroccan Atlas occurred
in the western High Atlas in the city of Agadir in 1960. The magnitude of the
event is estimated at mb = 5.9 with an intensity of Io = X for the majority of the
city (Medina et al., 1991; Alami et al., 2004). This event was responsible for a
14
Figure 1.5. GPS velocity vector map. Instantaneous maximum rate of
horizontal shortening of 1 mm/yr for the Middle Atlas Mountains (Reilinger,
unpublished).
15
Figure 1.6. Seismic map indicating earthquake localities with red circles and
determined fault plane solutions.
16
large amount of damage and the death of about 12,000 people. Although not
as destructive, recent events have also been recorded within the Middle
Atlas. In particular, an event estimated at mb = 4.0 occurred in 1979 along
the frontal thrust fault in the study area (Medina et al., 1991). In addition,
nearly 4000 events were recorded throughout northwest Africa and Iberia
from 1987 to 1994 including numerous smaller events throughout the High
and Middle Atlas (Alami et al., 2004).
In addition to seismicity, Quaternary volcanism is one of the more
prominent neotectonic features in the area. The timing of volcanism
coincides with the beginning of the present-day regional stress field and the
volcanic centers are approximately aligned NNW-SSE parallel to the
maximum convergence between North Africa and Iberia (Harmond and
Cantagrel, 1984; Ait Brahim, 1984; Morel, 1989; Gomez, 1996). Due to the
alignment of the volcanic centers, it has been suggested that they are related
to crustal scale extensional joints (Harmond and Moukadiri, 1986).
To address the active role of the Middle Atlas Mountains in the diffuse
plate boundary, a study location with identified fault scarps was chosen for an
integrated neotectonic and structural investigation. To begin, a structural
analysis along the front of the mountain range including the completion of a
geologic map and an approximately 12 km balanced cross-section provides
an estimation of shortening and geometry. Next, along the frontal thrust of
the range, fault scarps of Quaternary alluvium are evaluated through detailed
topographic mapping, which provides basis for a geomorphic analysis and
17
degradation modeling. Furthermore, the reconstruction of longitudinal stream
terrace profiles helps constrain a long term deformation history and
radiocarbon dates provide age constraints on the faulted surfaces. In order to
place the active tectonic observations into a larger context, the structural
geometry and the uplift rate are combined to give a minimum estimate of the
rate of horizontal shortening in the Middle Atlas.
18
Chapter 2 - Middle Atlas Structure
Introduction
In order to quantify the shortening across the southwest portion of the
Middle Atlas, it is necessary to constrain the spatial extent of geologic units
and understand the geometry of the structures in the study region. To
achieve these objectives, a general understanding of the Middle Atlas
structure is necessary along with a detailed geologic map and balanced
cross-section of the study area.
The Middle Atlas Mountains comprise three major geomorphic
provinces: the Tabular Middle Atlas (TMA), the High Moulouya/Missour Basin
area (HMMB), and the Folded Middle Atlas (FMA) (Figure 1.1 and Figure 2.1).
The TMA is the northeastern portion of the Moroccan Meseta consisting of
relatively flat lying Mesozoic strata and the majority of the Quaternary lava
flows. This region is dominated by sinistral strike-slip faulting including the
Tizi-n-Tretten Fault (Figure 2.1) (Gomez et al., 1996). The High
Moulouya/Missour Basin region is located to the southeast of the FMA at the
junction of the High and Middle Atlas and also contains flat lying strata. The
FMA, located between the TMA and the High Moulouya/Missour Basin, is the
focus region of the study and accounts for the large majority of the
compressive deformation in the Middle Atlas. To the northeast, the FMA get
as wide as about 55 kilometers corresponding to a major Jurassic rift basin.
Moving toward the southwest, the Folded Middle Atlas is only 20 to 25
kilometers wide.
19
Figure 2.1. Geologic map of the Middle Atlas Mountains. Labeled structures
include AOF – Ait Oufella Fault, TNTF – Tizi-n-Tretten Fault, and NMAF –
North Middle Atlas Fault. Study area is outlined by a black box. Figure
adapted from Gomez et al. (1998).
20
These three geomorphic provinces contain variable deformation styles.
The strain in the Middle Atlas Mountains is partitioned with strike-slip motion
concentrated to the northwest in the Tabular Middle Atlas and the majority of
the contraction found in the Folded Middle Atlas (Gomez et al., 1998). In
contrast, the HMMB constitutes the relatively undeformed footwall of the
Middle Atlas frontal thrust.
Bounding the Folded Middle Atlas to the north and separating the
region from the Tabular Middle Atlas is the North Middle Atlas Fault (NMAF)
(Figure 2.1). This fault exhibits primarily sinistral strike-slip motion with a
discontinuous braiding fault trace. Evidence for recent movement along the
fault is provided by faulted Quaternary lava flows and the faulting and folding
of Quaternary deposits (Fedan, 1989; Martin, 1981; Gomez et al. 1996). To
the south of the Folded Middle Atlas are the South Middle Atlas Fault and the
Ait Oufella Fault, which is the frontal thrust of the Cenozoic fold belt in this
region (Figure 2.1). The folds in the region trend primarily NE-SW, similar to
the fault traces, with a secondary E-W trend (Gomez et al., 1996).
Early studies on the development of the Middle Atlas involved wrench
faulting and flower structure (Fedan, 1989). More recent studies have
suggested thrusting in the Folded Middle Atlas linked to a decollement level
within the Triassic strata (Morel et al., 1993; Gomez et al., 1998). Based on
the structural work located northeast of the study area completed by Gomez
et al. (1998), there has been a total of 4.7 kilometers of total shortening
across Folded Middle Atlas. The majority of this shortening is accommodated
21
in the southeastern portion of the Middle Atlas toward the Ait Oufella Fault as
over-thrusting.
The study region, which is located in the narrow southwestern portion
of the Folded Middle Atlas, progresses from the northwest toward the
southeast across a series of anticlines and synclines leading up to the active
frontal thrust, the Ait Oufella Fault (Figure 2.1). Geologic mapping, fault
kinematic analysis, and the construction of a balanced cross section has been
completed in order to quantify the shortening across the study region.
Geologic Mapping of the Folded Middle Atlas
Geologic mapping of a transect was completed in order to understand
the structural styles in the southwestern Folded Middle Atlas and to provide
field basis for the completion of the balanced cross section. Two primary
methods were used in completing the map: 1) collection of structural
measurements and boundary contacts along a thin transect in the field, and
2) the processing and interpretation of remotely sensed imagery to
extrapolate the field data. Mapping of these data was completed digitally
using ArcGIS Deskstop.
Standard geologic field methods were used to determine orientations
of faults, bedding, and the location of fold hingelines. The Itzer, Morocco
1:50,000 scale topographic map was utilized for determination of
22
measurement localities in the field and as a reference during the completion
of the geologic map.
Stratigraphy
Recognizing stratigraphy and relative ages of the units in the field was
based on previous work in the region (Martin, 1981; Gomez et al., 1996;
Gomez et al., 1998; Itzer Map). The units are divided by age into Triassic,
Lower and Middle Liassic (both within the lower Jurassic), Cretaceous,
Neogene, and Quaternary deposits. Descriptions of these units as seen in
the study region follow.
Triassic - The unit includes red argillites, weathered basalts, and
volcaniclastic rocks. The red argillites were the most prevalent rock type of
the Triassic in the study area and served as a distinct marker.
Lower Liassic (Lower Jurassic) - The unit includes massive carbonates,
dolomitic carbonates, and multicolored marls (calcareous mudstone). The
cliff forming massive carbonates are very distinct and are prevalent
throughout the study region.
23
Middle Liassic (Lower Jurassic) – The unit includes sandy carbonates and
coral reef deposits. The reef deposits served as the distinctive rock type in
this unit.
Cretaceous – This unit lies unconformably on the Jurassic units. It includes
microconglomerate, marls and sandy fossiliferous carbonates. The sandy
carbonates were the most prevalent rock type of this unit and were typically
yellow in color.
Neogene – This unit includes conglomerates and sandy white lacustrine
limestones. Outcropping of this unit in the study region is almost completely
limited to the footwall.
Quaternary – These deposits include unconsolidated alluvium and colluvium.
Typically, these are found in stream channels, along side the streams as
terraces, or alluvial fan deposits.
Remote Sensing Analysis and Interpretation
Data Types
Two types of remote sensing data were utilized in completing photo
geologic mapping of the study region: aerial photographs and Advanced
Spaceborne Thermal Emission and Reflection Radiometer (ASTER) imagery.
24
These data were used for stereoscopic viewing and lithologic mapping of the
study region.
Aerial photographs at 1:60,000 scale were utilized prior, during, and
after field work was completed to assist in terrain interpretation and mapping.
The photographs were obtained in a 1973 flight mission and coverage was
limited to the front portion of the study region.
The ASTER sensor is one component of the Earth Observation
System (EOS) TERRA satellite that was launched in 1999 jointly by the
United States and Japan. The ASTER sensor records solar energy in 14
spectral bands through 4 different telescopes. The sensor measures three
bands in the visible near infrared (VNIR) with an additional back-looking
telescope for the generation of digital elevation models (DEM) at a 15 meter
spatial resolution. It measures six bands in the shortwave infrared (SWIR) at
30 meter resolution and five bands in the thermal infrared at 90 meter
resolution (Table 2.1). The VNIR and SWIR bands were used for this study.
The thermal bands were not used. PCI Geomatica OrthoEngine was used for
primary processing and generation of the DEM, while ER Mapper was utilized
for further processing. Table 2.1 outlines the performance parameters for the
ASTER radiometer taken from Fujisada (1995).
25
26
Table 2.1 Performance parameters for the ASTER radiometer. Table from
Fujisada (1995).
Analyses
Stereo Photogrammetry
Stereo viewing of aerial photographs was used to aid in the production
of the geologic map by extending rock unit contacts and interpreting
Quaternary deposits from above. In addition, it provided three-dimensional
perspective of the study area and an additional view of the landscape to
incorporate with the satellite imagery.
The ASTER imagery was also processed for stereoscopic viewing.
The stereographic bands (3N and 3B) were used to produce a digital
elevation model (DEM). Generation of the DEM was completed in order to
produce a higher resolution topographic image than the existing Shuttle
Radar Topography Mission (SRTM) dataset, which has a resolution of 90
meters. The DEM produced from the ASTER imagery has a 15 meter spatial
resolution revealing a much more detailed topography.
Multi-spectral image processing and interpretation
Lithologic mapping of the study region was completed by constructing
false-color images from the multi-spectral ASTER bands (e.g. bands 1-9 for
this study). Multiple processing approaches were utilized, including the
combination of different bands, creating band ratios, and principal
components analysis, in order to optimize the usefulness of the imagery.
After trying multiple combinations, two variations of the ASTER image were
27
created for the identification of the lithologic units. The first is a combination
of bands 7-3-1 as a red-green-blue (RGB) image (Figure 2.2). The second is
a combination of principal components 4, 5, and 6 (PC4-5-6) as an RGB
image (Figure 2.4).
The 7-3-1 image combines shortwave infared, near infared, and visible
bands (Figure 2.2). Similar images have been used by Pena (2004) for oil
and gas exploration. This combination has the advantage of preserving
morphological features as well as displaying lithological units in different
colors (Pena, 2004). The SWIR bands are useful here since many
sedimentary rocks have distinctive spectral characteristics in this portion of
the electromagnetic spectrum (Evans, 1988). For instance, limestone has
higher reflectance than shale throughout the SWIR (Figure 2.3) (Pena, 2004).
This image was very useful for extending unit contacts wherever bare rock
was visible (Figure 2.2). This image, however, does not clearly decipher units
when large amounts of vegetation are present. This is due to the high
reflectance of vegetation in the near infrared. Therefore, an additional image
that is not affected by vegetation is needed for the areas within the study
region covered in vegetation.
The PC4-5-6 image displays principal components 4, 5, and 6 as an
RGB image (Figure 2.4). The principal components analysis combines all
nine bands into a nine dimensional analysis. The analysis begins by shifting
the origin of a nine dimensional plot to the point representing the mean of all
28
Figure 2.2. ASTER band combination image with bands 7-3-1 (RGB). Study
area outlined on Figure 2.1. Yellow lines are lithologic contacts.
29
Figure 2.3. Shortwave infrared spectral reflectance curves for limestone and
shale. Location of the ASTER bands are shown in gray. Figure adapted from
Pena (2004).
30
Figure 2.4. ASTER image composed of principal components 4, 5, and 6
(RGB). Study area outlined on Figure 2.1. Yellow lines are lithologic contacts.
31
the bands. The next step creates axes along the line of greatest spread to
create PC1 (the principal axis of an ellipse in 2 dimensional analyses). The
analysis continues by creating principal components for all nine dimensions
(Drury, 2001). The principal components 4, 5, and 6 were used here to
eliminate the axes of highest correlation (e.g. topography and vegetation).
The effects can be observed in Figure 2.4 as topography can not be
distinguished and the regions with large amounts of vegetation (e.g. the
stream valleys) are black. This image was particularly useful in displaying
Triassic deposits, which are clearly shown in purple in the NW portion of
Figure 2.4.
The photo geologic mapping of the study region was completed
utilizing both stereo photogrammetry and processed multispectral ASTER
images. By utilizing imagery, the contacts between the rock units established
in the field were able to be extended. In addition, 3 point attitudes provided
additional topographic constraints (e.g. Itzer, Morocco topographic map) on
the mapped contacts and the extracted DEM, created from bands 3N and 3B,
was useful for displaying the topography of the study region.
32
Geologic Map of the Folded Middle Atlas
By combining the field data, stereographic photo interpretation, and the
multispectral ASTER imagery, the geologic map of a portion of the Folded
Middle Atlas was created (Figure 2.5). The map region is approximately 13
kilometers by 5 kilometers and is oriented NW-SE parallel to the cross section
line and the maximum Cenozoic shortening direction in the Middle Atlas.
Structural trends throughout the mapped region are generally oriented
NE-SW with secondary E-W fold trends. Major structures include the Ait
Oufella frontal thrust, a syncline, an anticline, a NW dipping fault, and a SE
dipping back thrust. The major syncline and anticline are open with shallow
limbs and extend throughout the narrow portion of the Folded Middle Atlas
toward the NE. Within the study area, the hingeline of the syncline rotates
from a NE-SW trend to an E-W trend. This arcuate hingeline is likely due to
the bend in the AOF just to the SW of the mapped region since its trace
nearly parallels the fault trace. The back thrust is also a prominent feature in
the region mapped in other studies (e.g. Martin, 1981; Gomez et al., 1998).
However, the NW dipping fault seems to be a smaller splay that terminates
into the back thrust just outside the mapped region. A minor anticline and
syncline occur near the frontal thrust and terminate within the study region
SW of the cross-section line (Figure 2.5). A small oblique dextral thrust fault
is located in the hinge of the anticline. These structures may also be in
response to the bend in the frontal thrust. In addition, smaller undulations in
33
Figure 2.5. Geologic map of the Middle Atlas study area. Scarp localities are
shown as stars. Balanced cross section line extends from A-A’.
34
stratigraphy occur throughout the mapped region; however, these are not
significant to the larger structure.
Structural Analyses
Field data, including bedding attitudes and fault kinematic data provide
basis and justification for the balanced cross-section. Two analyses were
performed including stereonet analysis of bedding measurements in the study
region and a fault kinematic study of structural measurements along the Ait
Oufella frontal thrust. FaultKin © and StereoWin © were used to analyze and
plot the data.
Stereonet analysis of bedding orientations throughout the study region
was completed in order to determine general fold axis orientations and the
finite strain axis. The poles to bedding were plotted on an equal area
stereonet. The best fit plane and its pole were determined and also plotted
on the same stereonet (Figure 2.6). The trend of the bedding poles (319º)
corresponds to the finite strain axis and reaffirms the principal NW-SE
transport direction through the southwestern portion of the Folded Middle
Atlas. The orientation of the pole to the best fit plane is an estimation of the
fold hingelines in the study region. Data generally fit well for a cylindrical fold.
This is an essential assumption used in the construction of the cross-section
that follows since cylindrical folding is required to eliminate mass transport
into and out of the plane of section.
35
36
Figure 2.6. Equal Area stereonet plot of poles to bedding. Strike of the cylindrical best fit plane
depicted by the red line is parallel to the finite strain axis. Line denoted by point 3 in red is a
general estimate of hingeline orientation.
Fault kinematic data were collected in a road cut along the Ait Oufella
frontal thrust (Figure 2.7). The data are plotted as faults and striations with
their associated Bingham axes and as a summed fault plane solution (Figure
2.8). The fault plane solution represents the unweighted geometric moment
tensor summation (Marrett and Allmendinger, 1990). The northwest dipping
plane (211º, 30º) corresponds to the Ait Oufella Fault. This orientation is
congruent with earlier fault kinematic analyses (e.g. Gomez et al., 1996;
Gomez et al., 1998) and attests to a NW to SE transport direction. The lack
of significant oblique movement along the fault plane supports the assumption
of two-dimensional movement within the balanced cross section plane. In
addition, the infinitesimal strain axes are co-axial with the finite strain axes
(cylindrical best fit; Figure 2.6) indicating pure shear through the study region.
This is an important observation since the co-axial strain, in addition to
cylindrical folding, are assumed in the construction of the balanced cross
section that follows.
37
Figure 2.7. Exposure of the Ait Oufella Fault in a road cut. Triassic Argillite is
being thrust over the top of Neogene conglomerate.
38
39
from the road cut in Figure 2.7.
Figure 2.8. Equal area stereonet plots of fault kinematic data. Kinematic data were collected
Balanced Cross-Section of the Folded Middle Atlas
Based upon the structural analyses, the overall structural style across
the Folded Middle Atlas appears to be two dimensional with cynlindrical
folding and co-axial strain. Hence, the construction of a balanced cross
section across the southwestern portion of the Folded Middle Atlas can place
a minimum constraint on the shortening across the fold belt. The section line
extends 13 kilometers from the High Moulouya platform across the AOF to a
back thrust and is oriented N45ºW-S45ºE (e.g. parallel to the axis of principal
shortening) (Figure 2.6).
At the frontal thrust, Triassic beds are oriented parallel to the fault
surface inferring layer parallel shear (Figure 2.9). In addition, deformation in
the study region is brittle in nature and fold geometries appear to be related to
bends in the fault surface (i.e. fault-bend folding). Within the cross section,
folding was generally assumed to be concentrated along “kink axes”, which
follow the fault-bend folding style outlined by Suppe (1983) (Figure 2.9).
Stratigraphic thicknesses throughout the cross section were constrained by
rock unit contacts within the study region and earlier studies of footwall
thicknesses (Martin, 1981; Lorenz, 1976).
The structural interpretations of the cross section have been balanced
using FaultFold ©. This program assumes parallel folding through conserving
mass by maintaining equal bed lengths and thicknesses throughout
deformation. The program was utilized to forward model the deformation by
40
41
Mountains.
Figure 2.9. Balanced cross section of the southwestern portion of the Folded Middle Atlas
adjusting initial parameters (e.g. fault geometries and total slip) to fit the
observed deformation. By restoring the beds in the study region to an
undeformed state, an estimate of 3.9 kilometers of total Cenozoic shortening
across the Ait Oufella Fault was determined (Figure 2.9).
Discussion
The “thin-skinned” structural style depicted in the cross section (Figure
2.9) is suggested by field observations. First, Triassic beds in the hanging
wall lie parallel with the AOF indicating a detachment surface. Second, in a
thick-skinned model, where the fault extends downward at a steeper angle
into the basement, the amplitude of the fault bends would increase.
According to the basic kinematics of fault bend folding, larger amplitude fault
bends will be compensated by steeper fold limbs (Suppe, 1983). Hence, a
low-angle fault geometry seems more appropriate for the study area.
For the Middle Atlas, the typical “layer cake” geometry associated with
thin skinned thrust belts must be slightly adjusted since this deformation is
superimposed on a zone of earlier extension. Within the hanging wall, lower
Jurassic (Liassic) strata are prevalent; however, they are not present in the
footwall of the AOF (i.e. the High Moulouya). Instead, Triassic strata are
unconformably overlain by Cretaceous strata (Martin, 1981). This implies a
paleogeographic boundary between the two regions, such as a Mesozoic
normal fault (Figure 2.9). Based on stratigraphic discordance between the
42
hanging wall and footwall of the AOF, this normal fault bound a major
Jurassic rift basin. Though, it should be said that this steep basement fault is
not reactivated during Cenozoic convergence due to geometric constraints
provided by the fold geometries previously described. However, basement
structure may play a role in features such as the lobe-like front of the Middle
Atlas and the shape of the syncline at the front.
Straitigraphic thicknesses throughout the study area are generally
uniform; however, the Middle Liassic unit does pinch out toward the southeast
into angular unconformity with the Cretaceous (Figure 2.9). Generally, this
contact is disconformable or in slight angular unconformity, although, more
pronounced angular unconformity has been documented (Ensslin, 1992;
Gomez et al., 1998). The existence of the angular unconformity in the
restored section of the study area suggests an initial slope of approximately
3º for the Triassic and Jurassic strata in the hanging wall. This initial dip
could be a result of tilting coincident with Mesozoic extension in the basin.
A more regional objective for this study was to determine if any along
strike variation in structural style and shortening occur along the Ait Oufella
Fault. In a study by Gomez et al. (1998), located about 25 kilometers
northeast of the study area in the Folded Middle Atlas, similar thin-skinned
deformation was observed with a total of 4.7 ± 0.7 kilometers of shortening.
The determined shortening from this study (i.e. 3.9 kilometers) is less;
however, the cross section from Gomez et al. (1998) covers a larger distance
and, therefore, contains more shortening. Regardless of the shortening
43
discrepancy, the similar structural styles provide collaborative data to support
a thin-skinned deformation model for the southwestern portion of the Middle
Atlas Mountains.
Within the study region, the Triassic, composed of relatively “weak”
rocks (e.g. argillite, evaporates, and weathered basalts), serves as the
detachment layer for the thin-skinned deformation. However, progressing
towards the northwest, the detachment merges with left lateral faults (e.g. the
North Middle Atlas Fault and the Tizi-n-Tretton fault) (Figure 2.1). The
geometry at depth at the junction of these faults is not well constrained and
can not be inferred from this study.
Regardless of the structure at depth, it is important to link these
structural observations and results to the active faulting occurring in along the
front of the Middle Atlas. In the section that follows, the active setting of the
Middle Atlas Mountains is evaluated in order to determine a shortening rate
across the Middle Atlas.
44
Chapter 3 Active Tectonism in the Folded Middle Atlas
Mountains
In order to understand the recent deformation and quantify the
shortening rate across in the southwestern portion of the Folded Middle Atlas,
a neotectonic investigation of the study region is necessary. Evidence for
recent faulting along the Ait Oufella frontal thrust is evident due to faulted
stream terraces, terrace and stream profile inflections, and recent seismicity.
An mb = 4.0 earthquake was recorded in 1979 along the Ait Oufella frontal
thrust (Medina et al., 1991). However, the earthquakes throughout the Middle
Atlas are typically quite small and consequently have poorly constrained
epicenter locations. Therefore, investigating the neotectonic geomorphology
(e.g. faulted terraces and stream and terrace profiles) along the frontal thrust
provides valuable information for interpretation of the active deformation in
the study region. In addition, this investigation provides data to quantify the
preliminary uplift rate and, in conjunction with the cross-section, the
shortening rate across the southwestern Folded Middle Atlas.
The primary objectives of the neotectonic investigation were to 1)
reconstruct longitudinal stream and terrace profiles to help constrain a long
term deformation history and 2) survey the fault scarps to provide basis for
geomorphic analysis and degradation modeling. In addition, radiocarbon
dates provide age constraints on the faulted surfaces.
Collection of field data was completed using real time kinematic (RTK)
GPS. This land surveying technique utilizes the carrier frequency in the GPS
45
signal to provide real-time corrections to centimeter accuracy. A base station
receiver broadcasts the phase of the carrier that it measured from the satellite
to a mobile unit, and the mobile unit compares that with the phase that it
received. This allows the relative position of the units to be determined with
centimeter precision; although, the absolute accuracy is a function of the
accuracy of the base station location. RTK GPS points were collected on
numerous stream terrace surfaces and on the faulted surfaces.
Longitudinal Stream and Stream Terrace Profiles
The longitudinal stream and stream terrace profiles were created in
order to constrain a deformation history across the Ait Oufella frontal thrust.
Profiles spanning nearly 2 kilometers across the AOF were produced to
observe tectonic related inflections (Figure 3.1). The present day stream
profile was created using the Itzer, Morocco 1:50,000 scale topographic map
and RTK GPS points, while terrace profiles were produced using RTK GPS
points only. Measured terrace surfaces are represented on Figure 3.1 as
solid lines with interpretive dashed lines projected between the measured
surfaces. Measured surfaces through the profiles are limited due to stream
erosion.
The inflection of the present stream profile and the terrace level 1
profile is evident and attests to the recent activity along the frontal thrust
(Figure 3.1). The Ait Oufella Fault cuts directly though terrace level 1 forming
46
a composite fault scarp that is discussed in much more detail in the section
that follows. The present-day stream profile contains a large inflection nearly
a half kilometer upstream from the fault zone. This is possibly due to the
active erosion occurring in the stream. As uplift continues at the front, the
upstream gradient will increase along with topography, and, therefore, the
erosive power of the stream will increase causing the inflection to propagate
upstream. The remnant terrace level 1 is not actively eroding, and therefore,
the inflection at the fault is preserved (i.e. the fault scarp). Another possible
cause for this inflection is a change in lithology. Near the inflection is the
Triassic/Jurassic boundary where Triassic weathered basalts and argillites
are overlain by massive cliff-forming carbonate. The differential weathering
between these layers could cause an inflection in the stream profile. Another
factor that may affect the shape of the present stream profile is the contour
interval of the topographic map used in its construction. The majority of the
profile is based off of a topographic map with a resolution of 20 meters.
Therefore, the exact location of the inflection could shift slightly toward the
fault zone.
Terrace surfaces 2 through 4 were not continuous across the fault
zone; therefore, inflections in their profiles were not observed. However, by
projecting the existing data, an increase in the slope of these surfaces across
the fault zone is required. Typical, undeformed stream terrace profiles would
also expect to steepen progressively upstream; therefore, the extent of the
tectonic influence on these surfaces is not well constrained.
47
48
Figure 3.1. Longitudinal stream and stream terrace profiles. Solid lines denotes
measured surfaces, while dashed lines are inferred. Location and dip of AOF are
accurate.
Quaternary Fault Scarps
Along the Ait Oufella Fault, 1 meter and 4 meter fault scarps of
Quaternary alluvium provide basis for geomorphic analysis and degradation
modeling (Figure 3.2 and Figure 3.3). The 1 meter scarp is inferred to be a
result of a single faulting event, while the 4 meter scarp, located in terrace
level 1, is inferred to be a composite scarp. Fault scarps were surveyed using
RTK GPS. High resolution digital elevation models of the scarp forms provide
topographic perspective and justification for the placement of perpendicular
profiles for degradation modeling. Radiocarbon data provide age constraints
on the faulted surfaces and required age estimates for the modeling. By
combining radiocarbon age data with degradation modeling, an uplift rate
and, therefore, a shortening rate can be determined.
Microtopographic Mapping
The topography of the fault scarps was investigated using high
resolution digital elevation models in order to provide an overhead
perspective of the scarp forms and to identify appropriate modeling profiles.
The DEMs were created from thousands of RTK GPS points collected over
the entire area of the scarps. ArcInfo Workstation was used to create the
DEM by interpolating between the data points. A 1 meter resolution DEM
was created for the composite scarp and the profile was chosen along the
most representative transect (Figure 3.4). A 0.5 meter resolution DEM was
49
created for the single event fault scarp and four perpendicular profiles were
chosen for degradation modeling (Figure 3.5). ArcScene was utilized to
create three-dimensional perspectives of the fault scarps based on their
DEMs (Figure 3.4 and Figure 3.5).
50
Figure 3.2. Field photograph of the composite scarp. The truck is on the top
surface of the scarp, while the people are on the bottom. The photograph
was taken facing west.
51
Figure 3.3. Field photograph of the singe event scarp. Relative displacement
and scale can be assessed from the backpack and camera case. Photograph
taken facing northwest.
52
53
Figure 3.4. 1 meter digital elevation model of the composite scarp. Scarp
profile used for degradation modeling is indicated with solid black line.
54
55
Figure 3.5. 0.5 meter digital elevation model of the single event fault scarp.
Scarp profiles used for degradation modeling indicated with solid black lines.
56
Radiocarbon Age Dating
Radiocarbon dating of the faulted surfaces was completed to provide
age constraints for the determination of an uplift rate. Two types of samples
were collected from within the faulted surfaces for age dating: individual
charcoal fragments and bulk sediment. The charcoal fragments were sent to
the University of Arizona, while the bulk sediment sample was sent to the
University of Georgia.
The ages received after analyses were reported as radiocarbon ages
before present (B.P.) and can not be utilized without calibration. In order to
determine a calibrated calendar age, the “raw” ages must be calibrated
against an atmospheric carbon curve since the amount of carbon in the
atmosphere has varied with time. The ages were calibrated to calendar years
using Oxcal calibration software (Bronk-Ramsey, 1998). The calibrated ages
in years B.P. are displayed in Table 3.1 as confidence intervals to one and
two standard deviations.
Scarp
Sample Type
Radiocarbon Age
δ13C
Cal Age (2 σ)
Single Event Alluvial Charcoal
3860 ± 41 BP
22.6
2470 BC - 2200 BC
Single Event Alluvial Charcoal
3831 ± 55 BP
21.1
2470 BC - 2130 BC
Composite
4480 ± 30 BP
22.8
3340BC - 3020BC
Bulk Colluvium
Table 3.1. Radiocarbon age data
57
The age dates for the single event scarp provide a maximum age for
the last faulting event along the AOF. The age date from the composite
scarp, however, is too young to represent the age of the faulted terrace level
1. It is more likely that this is the age of colluvial wedge material deposited in
the front of the scarp after a seismic event. This age is useful since it does
reaffirm recent faulting occurring along the frontal thrust. However, the uplift
rate associated with the single event scarp (e.g. 1 meter per 4300 years) is
not constrained and only gives an estimate based on the last surface
rupturing event in this area. Therefore, another method must be used to
estimate the age of the composite scarp in order to determine a more
representative and longer term uplift rate for the Ait Oufella fault.
Fault Scarp Degradation Modeling
The fault scarp degradation modeling was completed in order to
estimate the age of the composite scarp and determine the uplift rate along
the Ait Oufella Fault. Modeling of both scarps was needed in order to obtain
these estimates and determine a shortening rate across the AOF.
The degradation modeling completed for this study utilizes the diffusion
equation analysis outlined by Hanks et al. (1984, 2000). This analysis
inherently contains two physical assumptions. The first is that mass is
conserved at a local scale, and the second is that erosional diffusivity is not
much different than depositional diffusivity, which allows for the assumption of
58
a spatially constant mass diffusivity, k (Hanks, 2000). An additional
assumption is that mass moves downhill at a rate proportional to du/dx and
that k is not time dependant. Based on these assumptions, a onedimensional, homogeneous diffusion equation with constant coefficients can
be represented as:
du
d 2u
−k 2 =0
dt
dx
(1)
The relative elevation, u, is a function of profile distance, x, and time, t. In
other words, the rate of change in elevation is proportional to its curvature
(Hanks, 2000). Therefore, a region along a profile that is concave up or has
positive curvature will have u increase with time and vise versa. For
purposes of this study, k is represented in meters squared per kiloannum (m2
ka-1), while time is represented in ka.
Single Event Fault Scarp Degradation Model
The degradation modeling of the single event fault scarp was completed in
order to determine a representative mass diffusivity constant, k, for the study
region. This value is required to determine of the age of the composite scarp
and the uplift rate along the AOF. For a better estimate of k, four profiles
were chosen perpendicular to the trend of the scarp (Figure 3.5). The points
for each profile were extracted from the RTK GPS points along a 1 meter
wide swath.
59
To model the single event scarp degradation, an analytical solution to
equation 1 can be implemented for the case of a single vertical step in
topography of 2a imposed at t=0 upon a surface of pre-existing slope b
(Hanks, 1984). This corresponds to a single vertical dip slip faulting event of
a surface with an initial slope of b. According to Hanks (1984), the solution to
equation 1 with these initial conditions is:
⎛ x ⎞
u ( x, t ) = a erf ⎜⎜
⎟⎟ + bx
⎝ 2 kt ⎠
(2)
The profile distance x is referenced from an origin at the scarp mid-height.
This produces an uplift/down drop of ± a at x=0 for time, t, instantaneously
greater than zero.
For unconsolidated materials, such as the material that comprises the
faulted surfaces, an initial vertical offset is somewhat unrealistic. In reality,
the scarp would degrade almost immediately to the angle of repose for that
specific material. However, due to the large amount of curvature induced by
the initial offset, the scarp will degrade at a much faster rate initially. For the
time scales involved in this degradation study, the inaccuracy of the initial
condition will be muted.
Utilizing equation 2, the single event scarp profiles were modeled
(Figure 3.6 through Figure 3.9). Model variables (e.g. k, a, and b) were
adjusted to best fit the points for each profile. Variable values of a and b are
reasonable due to along strike variations in initial slope and the termination of
the scarp to the NE of profile 4 (Figure 3.5 and Figure 3.9). An age of 4300
60
B.P. (t = 4.3) based on the radiocarbon dates was used for all four profiles.
Results are shown on Figures 3.6 through 3.9.
The mass diffusivity, k, was determined from the four modeled profiles
to range from 0.9 – 1.5 m2/ka. This is a reasonable range of values for a
semi-arid region such as Morocco. Similar values were determined by Hanks
et al. (1984) for regions within the Basin and Range Province (i.e. 1.1 m2/ka).
61
62
Figure 3.6. Single event fault scarp degradation model along profile 1
(Figure 3.5)
63
Figure 3.7. Single event fault scarp degradation model along profile 2
(Figure 3.5)
64
Figure 3.8. Single event fault scarp degradation model along profile 3
(Figure 3.5)
65
Figure 3.9. Single event fault scarp degradation model along profile 4
(Figure 3.5)
Composite Fault Scarp Degradation Model
In order to determine an uplift rate along the AOF, degradation
modeling was completed for the composite fault scarp. Profile points were
extracted from the RTK GPS points along a 5 meter swath. Since the
solution used for the single event scarp is not applicable for a multiple event
scarp, a separate solution must be derived. Hanks et al. (1984)
parameterizes faults that have experienced multiple episodes of faulting in
terms of a slip rate, 2A. The result is an adjustment to Equation 1 in the form
of:
du
d 2u
−k 2 = A
dt
dx
(3)
The solution to this equation shown below is for an initial scarp offset of ± a at
x >and< 0 cut on a preexisting surface of slope b. For x > 0, A corresponds to
uplift, and for x < 0, A corresponds to downdrop so that u = 0 and x = 0 for all
t:
⎤ Ax kt −4xkt
⎛ x ⎞ Ax 2 ⎡ ⎛ x ⎞
u ( x, t ) = (a + At ) erf ⎜⎜
⎟⎟ +
⎟⎟ − sgn ( x )⎥ +
e + bx
⎢erf ⎜⎜
⎝ 2 kt ⎠ 2k ⎣ ⎝ 2 kt ⎠
⎦ k π
2
where
sgn ( x ) = +1
sgn ( x ) = −1
x>0
x<0
and a and A are now positive for all x (Hanks et al., 1984).
66
(4)
Further manipulation to this solution can be done since the existing
composite scarp height is known from the RTK GPS points. The updated
solution is of the form:
⎛ x ⎞ Qx 2
u ( x, t ) = (a + H ) erf ⎜
⎟+
2
⎝2 D ⎠
⎡ ⎛ x ⎞
⎤ Q
⎟ − sgn ( x )⎥ +
⎢erf ⎜
⎣ ⎝2 D ⎠
⎦ k
D
π
e
− x2
4D
+ bx
(5)
where
H = At
D = kt
A
Q=
k
Since one half the slip rate, A, multiplied by the age, t, renders one half the
height of the scarp, H is a known constant. By assigning three new variables
to the equation, the uplift rate and the age of the composite scarp can be
solved for simultaneously by holding H constant and statistically determining
the best combination of D and Q for the given scarp profile.
A root mean square analysis was performed to calculate a best fit H for
the profile data points. This value was then entered into the model and a root
mean square analysis was completed to determine the best fit values for D
and Q (Figure 3. 10). Based on the statistical analysis, a best fit D value of
200 and Q value of 0.0095 were determined with the initial conditions of H =
1.9 meters and b = 0.020 (Figure 3.11). By entering the range k values
determined from the single event scarp analysis, the age of the composite
scarp is estimated to range from 130 – 220 ka. In addition, based on these
values, the uplift rate along the Ait Oufella Fault is estimated to range from
0.02 – 0.03 m ka-1.
67
Figure 3.10. Root mean square analysis for D variable in composite fault
scarp degration model. Best fit value D = 200.
68
Figure 3.11. Composite fault scarp degradation model along profile line on
Figure 3.4.
69
By utilizing the dip of the fault (i.e. 30º), uplift rate can be correlated to
a rate of horizontal shortening. The shortening rate across the Ait Oufella
frontal thrust, determined from the uplift rate, is estimated to range from 0.03
– 0.05 m ka-1 or 0.3 – 0.05 mm/yr.
Discussion
GPS measurements of instantaneous convergence rates across the
Middle Atlas Mountains provide a maximum estimate of horizontal shortening
of approximately 1 mm/yr (Fadil et al., 2006). However, the 95% confidence
limits on the GPS velocities are similar to the velocities themselves, so the
GPS measurements do not constrain well the shortening across the Middle
Atlas. The current study determined a much lower estimate of 0.03 – 0.05
mm/yr for late Quaternary shortening across the frontal thrust, which could be
a result of how the strain is being accommodated. Even though the Ait
Oufella Fault is the primary frontal thrust of Cenozoic shortening, more
recently, strain may be accommodated by other structures in the region (e.g.
strike slip faults, blind thrusts, folds). This would account for the discrepancy
between the two data values.
In addition to strain accommodation, the modeled degradation age
contains many possible sources of error including some of the fundamental
assumptions of mass conservation at a local scale. The scarp surface has
undergone minor cultivation, and, depending on the extent of this cultivation,
70
this could have an effect on the age estimate. However, more likely, the
diffusivity value, estimated from the single event scarp, may be
underestimated. This value has a large effect on the calculated age. By
increasing the diffusivity value to 5.0 m2 ka-1 the age of the scarp is reduced
to 40 ka with a corresponding increase in the shortening rate to 0.16 mm/yr.
However, regardless of the estimate of k, accommodation of strain by other
structures is warranted since increasing k to an unrealistic value of 15 m2 ka-1
only results in a shortening rate of 0.49 mm/yr.
In addition to the shortening rate, the modeled composite scarp age
has implications on earthquake hazard. Based on the single event scarp
height (i.e. about 1 meter), it can be deduced that the composite scarp is the
result of four earthquake events. By combining this information with the age
of the scarp, the recurrence interval is estimated to range from 33 – 55 ka.
With a maximum age constraint of approximately 4300 years on the last
faulting event, the earthquake hazard for this region is moderate to low.
However, these results are preliminary and further age constraint on the
composite scarp is needed for a more certain estimate of recurrence interval.
Additionally, from the single event scarp height, an estimate of the
magnitude for the seismic event associated with the formation of the single
event scarp can be determined. The magnitude is estimated using the
empirical relationship between coseismic displacement and magnitude
established by Wells and Coppersmith (1994). Based on the height of the
scarp (i.e. 1 meter) and the dip of the AOF (i.e. 30 º), the displacement
71
associated with this seismic event is 2 meters. Placing this value into the
empirical relationship for the maximum displacement, an estimate of M = 6.7
provides a minimum constraint on the magnitude. Alternatively, the empirical
relationship for the average displacement gives a similar estimate of M = 6.7.
Therefore, there is the potential for a large destructive earthquake along the
Ait Oufella Fault despite the long recurrence interval.
Conclusion
The NE-SW trending Middle Atlas Mountains are an active
intracontinental mountain belt within the diffuse African – Eurasian plate
boundary. The Middle Atlas study area in the southwestern portion of the
Folded Middle Atlas contains fault scarps of Quaternary alluvium that provide
basis for the determination of the horizontal shortening rate from degradation
modeling. The late Quaternary shortening rate across the Ait Oufella frontal
thrust of 0.03 – 0.05 mm/yr determined in this study in much lower than the
maximum estimate of 1 mm/yr provided by GPS measurements. Most likely,
this discrepancy is due to the compressive strain within the region being
accommodated by structures other than the AOF.
To place these active tectonic observations in a larger context, the fault
and fold geometry was assessed by completing a 12.5 km balanced cross
section across the frontal thrust. From the cross section, a minimum estimate
of total Cenozoic shortening of 3.9 kilometers was determined. This structural
72
investigation provided basis for an interpretation of thin-skinned deformation
occurring along the front of the southwestern Middle Atlas. This correlates
well with other, along strike, studies (i.e. Gomez et al., 1998).
This study integrated structural, neotectonic, and geomorphic analyses
to investigate Middle Atlas foreland deformation. The knowledge gained from
this investigation provides the first geologic estimate of shortening across the
Middle Atlas Mountains of Morocco. In addition, the evaluation of the
structural geometry for the southwestern Middle Atlas front offers new data to
assist in understanding the style of deformation occurring within the diffuse
African-Eurasian plate boundary.
73
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