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Content uploaded by Matteo Basilici
Author content
All content in this area was uploaded by Matteo Basilici on Apr 13, 2021
Content may be subject to copyright.
Università degli Studi di Urbino Carlo Bo
Department of Pure and Applied Sciences
Ph.D. COURSE IN: Basic Sciences and Applications
CURRICULUM: Earth Sciences
SERIES: XXXIII
Thermal Structure and Active Tectonics of the Frontal Zone of the Zagros Fold and
Thrust Belt in Western Lurestan, Iran: New Insights from 3-D Geothermal Analytical
Modelling and 2-D Structural Finite Element Modelling
Academic discipline: GEO/10 – Solid Earth Geophysics
SUPERVISORS:
Prof. Stefano Santini
Prof. Stefano Mazzoli
CO-SUPERVISOR:
Ph.D. Antonella Megna
Ph.D. student: Matteo Basilici
ACADEMIC YEAR 2019-2020
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Table of contents
1 – Introduction………………...………………………………………… 4
2 – Geological and Seismological Background……...…………………… 7
2.1 – Geodynamic Evolution………………………………….…………………7
2.2 – Tectonic Setting and Crustal Structures………………….………..……...11
2.3 – Stratigraphy………………………...……………………………………..15
2.4 – Seismicity…………………………….…………………………………...17
3 – 2-D and 3-D Geological Models…………………………..……….....20
3.1 – Dataset……………………..……………………………………………...20
3.2 – Methods and Results………..…..…………………………………………25
4 – 2-D and 3-D Geothermal Models…………………….……………....28
4.1 – Previous Studies on Thermal Structure…………………………………..28
4.2 – Constraints and Assumptions……………………………...……………...29
4.3 – Analytical Procedure…..…………..……………………………………..33
4.4 – Geothermal Model Results………...……………………………………...40
5 – 2-D Structural Finite Element Models………………………..………48
5.1 – FEM Methodology and Previous Studies ………………………………..48
5.2 – Co-Seismic Simulation……………………………………………………51
5.3 – Inter-Seismic Simulation……..……………...……………………….…...54
5.4 – Structural FEM model Results……..………...……………………….…..56
6 – Discussion………………………………….…………………………59
6.1 – Geothermal Model……………………………………..……………..…...59
6.2 – Structural FEM Model ……………..…………………………...………...64
7 – Conclusions……………….……………….……….…………………68
8 – References……………………………………………….……………70
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1 – Introduction
In the last century several authors tried to understand the geological and
geophysical setting of the Zagros mountain chain. The result of these studies pointed
out a mesozoic continental-continental collision between Eurasia and Arabia plates due
to the closure of the Neo-Tethys ocean (Alavi, 1994, 2007; Agard et al., 2005;
Berberian and King, 1981; Dercourt et al., 1986; Mouthereau et al., 2012; Stampfli and
Borel, 2002; Vergés et al., 2011a). At the present, the orogenic process is still active,
as confirmed by recent geodetic calculation provided by Vernant et al. (2004). It
indicates a motion of the Arabian plate relative to fixed Eurasia of ~ 20 mm/yr, this is
partitioned between right-lateral motion along NE-SW-Striking faults and NW-SE
thrusts front (Blanc et al., 2003; Vernant et al., 2004; Talebian and Jackson, 2002). The
working area of this thesis is localized in the northern part of the Zagros orogen i.e.
Lurestan arc (Figure 1). In this sector of the Zagros orogen a master blind fault known
as Mountain Front Fault and named here Main Frontal Thrust nucleated a large
earthquake (Mw = 7.3) in November 12, 2017 highlighting the deep structure of this
sector of the Zagros Mountain chain (Basilici et al. 2020b; Gombert et al., 2019; Nissen
et al., 2019; Tavani et al., 2018a; Wang et al., 2018).
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Figure 1 – Tectonic sketch map of the Zagros mountain belt (modified from Basilici et al. 2020a).
The aim of this thesis is to explain how the Main Frontal Thrust interacts with the
entire Lurestan arc in terms of thermal structure and active deformation. To reach the
prefixed scope we used a mathematical approach studying the major thrust fault by
analytical and numerical procedures in order to realize exhaustive 2-D and 3-D models.
We present here the 3-D model of the Lurestan arc thermal structure calculated
using an analytical procedure interpolated with polynomial equations. The model took
into account topography, sedimentary cover, basement, thrusts and Moho depth
constrained by the physical properties of the rocks as thermal conductivity and heat
production rate (Basilici et al., 2019; 2020a).
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To study the active deformation, instead, we realized a 2-D model of co-seismic
and inter-seismic deformation of the Main Frontal Thrust applying the finite element
modelling methodology performed here using Marc Software (MSC Software
Corporation). In this case we set a pre-built model that was divided into several
domains, to which average values of Young’s modulus, Poisson’s ratio and density was
assigned. To resolve the system, the entire model was discretized into an equivalent
assemblage of small structures (elementary component, named mesh). As a result, for
each unit, a solution was formulated and combined to obtain the solution for the entire
system (Basilici et al., 2020b). The boundary condition was set up based on available
GPS dataset (Vernant et al., 2004).
Both 3-D thermal structure model and 2-D active deformation FEM model were
calculated on a geological model describing the main Lurestan arc crustal structures.
The geometry of the tectonic features was built on the base of previous published
geological cross section (Tavani et al., 2018a; 2020; Vergés et al., 2011a). At the end,
we compared the obtained thermal structure with previous studies and the active
deformation model with an investigation of a large-scale features of topography.
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2 – Geological and Seismological Background
2.1 – Geodynamic Evolution
The first recognizable tectonic event in the study area occurred at the end of the
Carboniferous period, with the onset of the Variscan orogeny. Prior to this time, the
entire region was a stable continental platform. Paleomagnetic poles indicate that Iran
remained close to Arabia during at least the early part of the Permian, to detach itself
subsequently, at the end of the same period (Berberian and King, 1981).
The entire Mesozoic represented the era in which the High-Zagros Alpine Ocean
in the southern region of Iran closed, during this time the Arabian foreland was
subjected to progressive subsidence and normal faults parallel to the old continental
margin was formed, these extensional movement was linked to the opening (occurred
at that time) of the Neo-Tethys ocean (Jassim and Goff, 2006; Berberian and King,
1981).
A phase of rifting occurred from the Early to the Late Jurassic time (Pratt and
Smewing, 1990). The extensional deformation is testified by syn-sedimentary faults,
block tilting, unconformities and by facies and thickness change (Tavani et al. 2018a;
2018b) that produced a strong control on the structure of the present mountain belt
(Bahroudi and Talbot, 2006; Jackson and Fitch, 1981; Tavani et al., 2020). Schettino
and Scotese (2002) suggested that the mechanism of this rifting phase was due to a
tensional force related to the north easterly subduction of the old Neo-Tethys oceanic
lithosphere beneath the active margin of Eurasia extending from the northern Turkey
to the Sanandaj-Sirjan Zone of Iran.
Another phase of oceanic floor spreading around the northern margin of the
Arabian Plate is marked by an important regional unconformity related to Mid
Tithonian which marks the boundary between Megasequences AP7 and AP8 (Sharland
et al., 2001). The oceanic floor spreading continued during Early Cretaceous until Late
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Cretaceous in which the south-eastern oceanic crust was abducted onto the Arabian
Plate, it formed the ophiolite succession actually representing the Neo-Tethys ocean
remainder (Figure 2), In Early Campanian a pulse of compression formed a foreland
basin around the northern margin of the Arabian Plate (Warburton et al., 1990) and the
oceanic lithosphere of the southern Neo-Tethys are interpreted to have been subducted
into the southern subduction zone (Searle et al. 2004). Also during this time, pre-
abducted ophiolites were metamorphosed and finally involved into the Zagros suture
zone during late Miocene (Jassim and Goff, 2006; Leterrier, 1985).
9
Figure 2 – Sketch of the geodynamic evolution of the Lurestan arc from Late Cretaceous to Present
(modified from Agard et al., 2011). Abbreviations: SSZ – Sanandaj-Sirjan Zone; ZFTB – Zagros Fold
and Thrust belt; SFB – Simply Folded Belt; ZIZ – Zagros Imbricated Zone; MRF – Main recent Fault;
HZF – High Zagros Fault; MFT – Main Frontal Thrust.
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Evidence of subduction and the final closure of the Neo-Tethys ocean are
attributable to Paleocene-Eocene (Ghasemi and Talbot, 2006; Mouthereau et al. 2012).
In late Eocene (~35 Ma) further compressional event occurred driven by the distal
continental margin negative buoyancy, resulting in the effective continent-continent
collision between Eurasia and Arabia during Oligocene (~25 Ma) (Agard et al. 2005;
Gavillot et al., 2010; Mouthereau et al. 2012).
The temporal evolution of the Zagros topography and adjacent Iranian plateau
uplift scarcely well understood but it is generally believed that the subsequent growth
of the large Zagros anticlines occurred during Miocene (~15 - 12 Ma; Mouthereau et
al. 2012) to continue during Plio-Pleistocene in the Zagros Fold and Thrust belt
(Hessami et al. 2001). The presence of dextral strike-slip dominated and reverse
dominated domains within the Kurdistan area indicates that an oblique convergence
guides the entire history of the Zagros orogeny since the Late Cretaceous obduction of
the Neo-Tethys oceanic realm (Allen et al., 2011; Sadeghi and Yassaghi, 2016).
Currently it is believed that convergence slowed down by ~30% from ~5 Ma to
present day (Austermann and Iaffaldano, 2013). On the contrary, from ~5 Ma the
Zagros Fold and Thrust belt in the Kurdistan region experienced major basement-
involved thrusting, (Allen et al., 2004; Austermann and Iaffaldano, 2013; Mouthereau,
2012; Koshnaw et al., 2017). During this recent period the shortening in the northern
part of the Zagros Fold and Thrust belt was ~49 km (Blanc et al., 2003) with an active
deformation rate of ~10 mm/yr (Allen et al., 2004), but a qualitative analysis of striated
fault planes suggests a variable shortening direction along the strike of the mountain
belt. However, shear sense indicators from thrust surfaces in various parts of the
mountain belt reveal a dominant SW direction of transport (Alavi, 2007; Csontos et al.,
2012; Hessami et al., 2006).
The last geodetic calculation was performed in 1999 September and 2001
October, it confirmed a northward motion of the Arabian plate of ~20 mm/yr in the
fixed-Eurasia reference frame (Vernant et al., 2004). The Iran block is actually
considered as a microplate located in between the Arabian plate and the Eurasian plate
11
in sensu stricto (Alavi, 1994; 2007; Berberian and Berberian, 1981; Colman-sadd,
1978; Stocklin, 1968).
2.2 – Tectonic Setting and Crustal Structures
The Lurestan arc is part of the so-called Zagros Fold and Thrust belt (ZFTB). It
comprises the Simply Folded Belt (SFB) and the Zagros Imbricated Zone (ZIZ; Figure
3). The SFB constitutes the external and less-strained part of the active Zagros orogen
wedge, bounded from the ZIZ to the north-east by the High Zagros Fault (HZF). The
suture zone between the ZIZ of the Arabian Plate and the Sanandaj-Sirjan Zone (SSZ)
of the Iran block is represented by the Main Recent Fault (MRF; Alavi, 1994;
Berberian, 1995). An important feature of the Arabian plate margin is the coexistence
of thin-skinned and thick-skinned settings (Le Garzic et al., 2019; Nissen et al., 2011;
Tavani et al., 2018a; 2018b).
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Figure 3 – Sketch map of the Lurestan arc tectonic setting. The yellow line represents the political
boundary between Iraq and Iran. Red lines are the traces of Main Recent Fault (MRF) and High
Zagros Fault (HZF). The red dashed line is the trace of Main Frontal Thrust (MFT).
The MRF is the most studied fault of the entire mountain chain. It is the major
seismically active right-lateral strike-slip fault and it is structurally distinct along its
entire length (Berberian, 1995). Based on the drainage pattern restoration, the offset of
the MRF may be as much as 70 km geometrically linked to a shortening of ~50 km
which is a substantial fraction of the 85 – 140 km total Arabia – Eurasia convergence
over the last 3 – 5 Ma. The calculated horizontal slip rate on the MRF would then be
13
in the range of 10 – 17 mm/yr, with a vertical component of 0.1 – 0.2 mm/yr
(Authemayou et al. 2006; Talebian and Jackson, 2002). Using a FEM model, Vernant
and Chéry (2006) calculated a strike-slip fault accommodation of ~25% of the total
lateral shear strain, the remaining ~75% could occur in the orogen itself by a more
distributed deformation.
The High Zagros Fault (HZF) is a reverse fault with a vertical displacement more
than 6 km. Uplift is ongoing since the Early Miocene. In the south-eastern Zagros this
fault reaches the Mountain Front Flexure (MFF) and it was associated with several
earthquakes with Mw ≥ 5 (Berberian, 1995).
The Mountain Front Flexure (MFF; Falcon, 1961) named here Main Frontal
Thrust (MFT, Basilici et al. 2020a; 2020b) represents the major thrust structure and,
therefore, the main topic of this thesis. It is a large basement blind fault showing the
typical ramp-flat geometry of more than 1350 km long and a max value of 9.7 km of
cumulative displacement. It accommodated most of the shortening (Hessami et al.,
2006) and it has important structural, topographic, geomorphic and seismotectonic
characteristics (Berberian, 1995; Emami et al., 2010; Koshnaw et al., 2017; Sherkati et
al., 2006; Tavani et al., 2018a). Topographical and geophysical evidences suggest that,
in the Lurestan arc, the MFT developed at 8.1 – 7.2 Ma (Homke et al., 2004; Koshnaw
et al., 2017). Tavani et al. (2020) suggest that it nucleates in the inner portion of the
necking domain of the Jurassic rift, where the mid-crustal ductile level is sufficiently
thick to promote the development of a large and interconnected decollement, from
which the MFT emanates.
The most north-eastward topographic expression of the MZT is represented by
the Anaran anticline which is also the most recent (1.5 Ma) and external fold of the
Zagros orogen (Emami et al., 2010). It was interpreted by Vergés et al. (2009; 2011a)
as forming part of the multi-detachment fold system of the Lurestan arc governed by
the varying mechanical stratigraphy. The sinusoidal shape of the MFT and its different
position in the Kirkuk embayment and Lurestan arc was promoted by lateral
segmentation of the Jurassic rift accommodated by transfer faults (Tavani et al., 2020)
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and it is probably that in this boundary area (Lurestan arc – Kirkuk embayment) the
deep faulting is decoupled by the shallow folding (Nissen et al., 2019). The main
seismic events of the ZFTB are largely nucleated by this thrust fault (Talebian and
Jackson, 2004). Recently, earthquake magnitude reached high values as testified by
events of November 12, 2017 (Mw = 7.3; Tavani et al., 2018a) and November 25, 2018
(Mw = 6.3).
Several authors investigated the thickness of the crust of the Iran sector using
different techniques. On this thesis we took into account only those work where the
Lurestan sector is comprised.
Paul et al. (2006; 2010) calculated a Moho depth comprises in the range of 40 -
55 km for the Lurestan arc using two temporary passive seismic experiment across the
central and northern Zagros. Jiménez-Munt et al. (2012) produced the first 3-D model
of the crustal geometry calculated from the geoid height and elevation data combined
with thermal analysis, confirming previous studies (Paul et al., 2006; 2010) and
tightening the Lurestan arc Moho depth in a range of 44 – 52 km. All results obtained
by previous authors were combined and used in Tunini et al. (2015) to produce a
numerical 2-D model of the thermal, compositional, density and seismological
structure of the crust and upper mantle along two transect across the entire Arabia-
Eurasia collision zone. A recent 2-D model of the sedimentary cover-basement
interface and Moho depth was performed by Teknik et al. (2019) coincident with a
CIGSIP (China-Iran Geological and Geophysical Survey in the Iranian plateau) seismic
profile. Therefore, the results are constrained mainly by a receiver function seismic
section, but also by published highest resolution data about crustal Iranian geophysics
available at present. Today new Moho depth investigation techniques are improving,
as the Moho estimation from GOCE Gravity Data (Heydarizadeh-Shali et al., 2020)
and phase velocity ambient noise tomography (Movaghari and Javan-Doloei, 2019)
confirming, in Iran, the results of previous authors.
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2.3 – Stratigraphy
The characteristic collision zone of the Zagros mountain chain is constituted by
two main domains belonging to Arabian plate and Iran block respectively. ZFTB
formed by the typical sedimentary succession of the Arabian margin (Figure 4), resting
on top of a Pre-Cambrian crystalline basement which does not take part in the typical
shallower folding, the basement-cover interface is therefore a decollement level (Alavi,
2007; Blanc et al., 2003; Casciello et al., 2009; Colman-sadd, 1978; Emami et al., 2010;
Hessami et al., 2001; Massaro et al., 2019; Sepehr and Cosgrove, 2004; Vergés et al.,
2011a). Probably the Lurestan arc does not present the Hormuz salt, a sequence of
evaporites that were deposited during Early Cambrian and located instead, on the Fars
arc (Talbot and Alavi 1996; Vergés et al., 2011a; 2011b). Teknik and Ghods (2017)
and Teknik et al. (2019) calculated the total thickness of the Arabian Cover using a
fractal spectral method to the aeromagnetic map of Iran and a gravity magnetic
modelling respectively. The result is an Arabian Cover thickness of the Lurestan arc
comprised between 6 and 16 km. Therefore the basement is not exposed in the ZFTB
and the nearest exposures are mapped in central Iran (Saghand Area; Ramezani, 1997;
Ramezani and Tucker, 2003).
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Figure 4 – Mesozoic-Cenozoic stratigraphy of the Lurestan arc sedimentary cover (modified from
Beydoun et al., 1992; Jassim and Goff, 2006; Motiei, 1993; Sepher and Cosgrove, 2004).
17
The ZIZ is a more stratigraphic complex zone, it is the core of the orogen where
the intensity of deformation has reached its maximum. It presents the remainder of the
Neo-Tethys ocean obducted in an ophiolitic succession (Figure 2), a Radiolarite Zone
and a series of magmatic units (Alavi, 2007; Agard et al., 2005; Berberian and King,
1981; Dercourt et al., 1986; Ghasemi and Talbot, 2006; Leterrier 1985; Sadeghi and
Yassaghi, 2016; Vergés et al., 2011a).
The Sanandaj-Sirjan Zone (Iran Block) is formed mainly by methamorphic rocks,
by Jurassic to Early Eocene calk-alkaline magmatic rocks, and by the products of
Middle Eocene gabbroic plutonism (Alavi, 1994; Baharifard et al., 2004; Berberian
and Berberian, 1981; Leterrier, 1985).
2.4 – Seismicity
The present seismicity of the Zagros orogen is studied from late ‘60s (e.g. Falcon,
1969; Dewey and Grantz; 1973) and it covers the entire width of the Lurestan arc.
Berberian (1976) produced the first seismotectonic map of Iran. Today, earthquake
data from 1967 are available on Global CTM and USGS catalogues (Figure 5).
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Figure 5 – Earthquake with Mw ≥ 5.5 recorded from 1967 to the present day in Lurestan arc. The red
dot is the main seismic event of November 12, 2017 (Mw = 7.3). The data are from the USGS
catalogue (https://earthquake.usgs.gov/earthquakes/search/; last access: 04 May 2020).
There are two distinct types of seismic active fault plane solution in the Lurestan
arc, a dextral strike-slip NW-SE striking faults and reverse faults respectively (Allen
et al., 2011; Authemayou et al., 2006; Basilici et al. 2020a; Talbot and Alavi, 1996;
Talebian and Jackson 2002). The strike-slip faulting nucleates seismic events with
increasing magnitude going from SW into the SFB, where the reverse faulting is
19
dominating, to NE in proximity of the MRF. Here the magnitude of strike-slip fault
nucleated earthquakes reaches its maximum (Mw ~ 5; Talebian and Jackson, 2002).
The seismicity confirmed the study of Vernant and Chery (2006) in which they
calculated an accommodation of ~25% of the total lateral shear strain carried by the
MRF.
The reverse faulting comprises minor shallow reverse faults and deep thrusts
localized into the basement (Hessami et al., 2006; Jackson, 1980; Jackson and Fitch,
1981), the master blind thrust, named here MFT, can generate large earthquakes
(Berberian, 1995) as confirmed by November 12, 2017 (Mw = 7.3) seismic event
nucleated at depth of ~19 km (Basilici et al. 2020a; Gombert et al., 2019; Nissen et al.,
2019; Tavani et al., 2018a; Wang et al., 2018). Therefore, the totality of the seismicity
(seismogenetic level) occurs into the crust or upper lithospheric mantle (depth range of
5 – 35 km). There are not recorded earthquakes into the deep mantle below the ZFTB
(Engdahl et al., 2006). This particularity of the Arabia-Eurasia collision zone made the
authors understand that the subduction of the Arabian plate below the Eurasian plate is
inactive at present and the slab put in place a break off (Austermann and Iaffaldano,
2013; Monthereau et al., 2012).
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3 – 2-D and 3-D Geological Models
3.1 – Dataset
To improve analytical and numerical calculations in the Lurestan arc, a geological
model of the area is requested as a base for mathematical procedures. For this purpose,
we used the published balanced cross sections located on the study area. In Figure 6
we show geological cross sections located on the Lurestan arc and used to build 2-D
and 3-D geological models. In each geological cross sections (shown in Figures 7, 8
and 9) the entire sedimentary succession was unified in a unique geological unit i.e.
Arabian sedimentary cover, resting on top of the crystalline basement (Basilici et al.,
2019; 2020a; 2020b).
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Figure 6 – Tectonic sketch map of the Zagros mountain belt. The yellow line represents the political
boundary between Iraq and Iran. Red lines are the traces of the Main Recent Fault (MRF) and the
High Zagros Fault (HZF). The red dashed line is the trace of the Main Frontal Thrust (MFT). The
black segments AB, CD and EF are the cross sections by Tavani et al. (2018a; 2020) and Vergés et
al. (2011a) respectively and shown in Figures 7, 8 and 9. The violet line GH is the 2-D geological
model shown in Figure 10. The blue rectangle is the area of the 3-D geological model shown in Figure
12. Red star is the epicentral position of the main seismic event of 12 November 2017 (Mw = 7.3;
Tavani et al., 2018a).
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In Figure 7 is shown the balanced geological cross section (AB segment of Figure
6) by Tavani et al. (2018a). It based on geological observation, interpretation of seismic
reflection profiles and well data, and crosses the epicentral area of 12 November 2017
earthquake (Mw = 7.3) so providing a comprehensive picture of the MFT geometry.
Figure 7 – Published geological cross section by Tavani et al. (2018a), AB segment in Figure 6. The
entire sedimentary cover was unified in a unique unit (Basilici et al., 2019; 2020a; 2020b) The focal
mechanism is projected in section and it shows the hypocentre of the 12 November 2017 earthquake
(Mw = 7.3).
23
In Figure 8 is shown the balanced geological cross section (CD segment of Figure
6) by Tavani et al. (2020). It based on geological observation, geomorphic analysis,
interpretation of seismic reflection profiles and earthquake data. It provides to highlight
the thick-skinned structural setting of the Lurestan arc central part.
Figure 8 – Published geological cross section by Tavani et al. (2020), CD segment in Figure 6. The
entire sedimentary cover was unified in a unique unit (Basilici et al., 2019; 2020a; 2020b). Focal
mechanisms are projected in section and show the hypocentres of Mw > 5 earthquakes associated
with the aftershock sequence of the 12 November 2017 earthquake (Mw = 7.3; Figure 7).
24
In Figure 9 the balanced geological cross section is shown, (EF segment of Figure
6) as provided by Vergés et al. (2011a). Paul et al. (2010) constrained the Moho depth
using seismic receiver functions and seismic tomography sections. It provides to
highlight the deepest part of the Lurestan arc.
Figure 9 – Published geological cross section by Vergés et al. (2011a), EF segment in Figure 6. The
entire sedimentary cover was unified in a unique unit (Basilici et al., 2019; 2020a; 2020b). Moho
depth and geometry are constrained by Paul et al. (2010).
25
3.2 – Methods and Results
In Figure 10 is shown the 2-D geological model (GH segment in Figure 6). It was
implemented on the basis of the geological section by Tavani et al. (2018a; Figure 7)
because its location crosses the epicentre of 12 November 2017 earthquake (Mw = 7.3).
In order to include the entire MFT in the section, this was integrated with the two
further section by Tavani et al. (2020; Figure 8) and Vergés et al. (2011a; Figure 9).
Moho depth and geometry are included in the 2-D model by using a projection of the
Moho calculated by Paul et al. (2010).
Figure 10 – 2-D Model (GH segment in Figure 6). This composite section was built by integrating
sections by Tavani et al. (2018a; Figure 7) for the northern sector, by Tavani et al. (2020; Figure 8)
for the central section and by Vergés et al. (2011a; Figure 9) for the southern sector. The Moho was
projected using profile by Paul et al. (2010). The crust of the model was subdivided into three different
homogeneous zone as shown in the legend (modified by Basilici et al., 2019; 2020a; 2020b).
Putting the 2-D model (Figure 10) and the published balanced geological cross
section (Figures 7, 8 and 9) in prospective into a 3-D space (Figure 11), it was possible
to build a 3-D geological model of the Lurestan arc (blue rectangle in Figure 6). The
used 3-D space covered an area of 320 km × 240 km, approximately as the area of the
entire Lurestan arc. The chosen model’s depth was 80 km in order to provide a better
view of the entire crust and part of the lithospheric mantle. The 3-D model was built
using Blender, a free and opensource 3-D computer graphics software downloadable
26
in its official website (https://www.blender.org; last access: 26 December 2020).
Blender’s features include 3-D modelling, fluid simulation and particle simulation and
it was recently used to show complex 3-D geometries in many fields of study including
geology (Basilici et al., 2020a; Florinsky and Filippov, 2019).
Figure 11 – Prospective of 2-D model (Figure 10) and published balanced geological cross section
by Tavani et al. (2018a; Figure 7), Tavani et al. (2020; Figure 8) and Vergés et al. (2011a; Figure 9)
into a 3-D space. The Violet line and the blue rectangle represent the 2-D model and the 3-D model
area respectively (Figure 6).
The 3-D model (Figure 12) results in a volume of 6,144,000 (320 × 240 × 80) km3
of rock. The topography was integrated using a 30 m resolution ASTER GDEM. The
model took into account the MRF and MFT, while minor faults were not included to
simplify the modelling procedure. In this 3-D case the crustal thickness was included
using the 3-D Moho depth and geometry calculated by Jiménez-Munt et al. (2012).
27
Figure 12 – 3-D geological model (modified from Basilici et al., 2020a) (a) Complete 3-D model of
6,144,000 (320 × 240 × 80) km3 of rock volume, based on geological section by Tavani et al. (2018a;
Figure 7), Tavani et al. (2020; Figure 8) and Vergés et al. (2011a; Figure 9). Topography is from a
30 m resolution ASTER GDEM. (b) MFT footwall topography (red line) including deep sector
beneath the branch line with the MRF (violet line). (c) Moho geometry based on Jiménez-Munt et al.
(2012), with Moho depth in km (white numbers). In order to provide a better view of MRF and MFT
geometry, the lithospheric mantle is not shown as a different block beneath the Moho discontinuity.
28
4 – 2-D and 3-D Geothermal Models
To obtain a better knowledge of the Zagros thrust belt and the impact of the MFT
on the thermal structure of the Lurestan arc, we improved an analytical methodology
which took into account the temperature variation due to the re-equilibrated conductive
state after thrusting. Frictional heating, heat flow density data, and a series of
geologically derived constraints are used in the model. The mathematical procedure is
based on available information about Moho (Francois et al., 2014; Jiménez-Munt et al.,
2012; Paul et al., 2010) and thermophysical characteristics of rocks (Byerlee, 1967,
1978; Tunini et al., 2015). The analytical methodology was applied on the 2-D
geological model presented in the previous chapter to obtain geotherms and isotherms
(Basilici et al., 2019). The resulting 2-D geothermal model was compared with the
numerical results obtained by previous studies (Kargaran and Neubauer, 2015;
Shekarifard et al., 2012; Tunini et al., 2015). Subsequently, the same procedure was
applied on the 3-D geological model presented on the previous chapter to elaborate the
first 3-D model reproducing the thermal structure of the Lurestan arc (Basilici et al.,
2020a).
4.1 – Previous Studies on Thermal Structure
Several authors studied the thermal structure of the Zagros complex and adjacent
zones, resulting in a series of models in which they calculated surface heat flow,
isotherms, geotherms and thermomechanical scenarios for the Zagros collision zone
(Főrster et al., 2010; Francois et al., 2014; Kargaran and Neubauer, 2015; Schutz et al.,
2014; Shekarifard et al., 2012; Tunini et al., 2015; Vernant and Chéry, 2006). Vernant
and Chèry (2006) adopted a surface heat flow of = 40 mW/m2, assuming radiogenic
sources in the crust, and a mantle heat flow of = 10 mW/m2 to calculate the first
thermal field of the Zagros thrust belt used to produce a 2-D mechanical finite element
29
model. Főrster et al. (2010) calculated geotherms of the Arabian Shield in Jordan based
on density, thermal conductivity and radiogenic heat production of a unique set of
samples from uppermost crust down to the lithospheric mantle. Shekarifard et al.
(2012) produced a 1-D thermal model of Alborz (Northern Iran) based on
physical/chemical equations and geological assumptions to reconstruct burial and
thermal history, taking into account different tectonic scenarios. Schűtz et al. (2014)
calculated heat flow in Mesozoic sediments of Central and Southern Israel using the
classical approach of heat-flow determination by implementing in the analysis high-
precision continuous temperature logs obtained in air- and/or water filled boreholes.
Francois et al. (2014) produced a series of 2-D numerical models of the Zagros/Central
Iran to investigate the continental subduction and to explain the current topographic
uplift, incorporating a free upper surface erosion, rheological stratification, brittle-
elastic-ductile rheologies, metamorphic phase changes and account for the specific
crustal and thermal structure of the Arabian and Iranian continental lithospheres.
Kargaran and Neubauer (2015) investigated the lithospheric thinning and subsequent
formation of the Iranian plateau using a 2-D numerical modelling, taking into account
heat flow and mineral thermobarometry behaviour from west to east. Tunini et al.
(2015) applied a combined geophysical-petrological methodology in order to study the
thermal, compositional, density and seismological structure of the crust and upper
mantle along two transect across the Arabia-Eurasia collision region. They calculated
a 2-D numerical model of thermal structure showing isotherms from surface to
asthenosphere.
4.2 – Constraints and Assumptions
Molnar et al. (1983) considered three sources of heat that affect the temperature
in a region of thrust faulting: heat supplied from the mantle, radiogenic heating within
the crust and frictional heating along the fault. The effect of these three source of
heating can be isolated from one another, treated separately, and then combined.
30
The analytical procedure was applied on a series of pseudo-well (Figure 13) traced
on geological models both 2-D and 3-D. Subsequently, we interpolated the model data
by a second-order power-law polynomial equation, obtaining the best-fit analytical
curves approximating the geothermal values. In the case of the 3-D model, we
constructed isothermal surfaces using the minimum curvature spline interpolation
technique (Basilici et al., 2019, 2020a).
Figure 13 – Pseudo-well representation (modified from Basilici et al., 2019). is the heat flux at
Moho; and are density of the sedimentary cover and basement, respectively; and are
thermal conductivity for the sedimentary cover and basement, respectively; and are heat
production rates for sedimentary cover and basement, respectively; is the basement thickening
value.
To perform the 2-D geothermal model we considered five pseudo-wells in total
(2, 3 and 4 are shown in Figure 14) located along the 2-D geological model presented
31
in the previous chapter (Figure 10, GH segment in Figure 6). For each pseudo-well we
considered the following parameters and assumption:
1. Altitude. The surface was considered as flat at 0 m.a.s.l.
2. Thickness of the Arabian cover . We considered a constant thickness of the
sedimentary cover of 7.5 km. The density of the Arabian cover was fixed to
= 2.55 × 103 (Teknik et al., 2019).
3. Thickness of the Arabian basement . We considered a variable thickness of
the basement based on the depth of the Moho discontinuity calculated by Paul et
al. (2010). The density of the basement was fixed to = 2.80 × 103
(Teknik et al., 2019).
4. Thrust depth and angle. The pseudo-wells cross the MFT at different depths with
a certain angle.
5. Amount of basement offset by the MTF . Taking into account the
reconstructed geometry of the Arabian margin by Le Garzic et al. (2019), the
timing of activity of the faults in our 2-D geological model (Figure 10) is variable
and it is divided into three steps: (i) Campanian/Maastrichtian, 84 – 66 Ma for
the MZT and HZF, (ii) Miocene, 20 – 10 Ma for the MRF and the Marakhil,
Seykh Saleh, and Miringeh faults, and (iii) 10 Ma-Present for the MFT. We
considered a total shortening of 20 km in the range of 20 – 10 Ma. The resulting
slip rate and the basement thickening were in the range of 1 – 2 mm yr-1 and
3.5 – 5.8 km respectively (Basilici et al., 2019; 2020a). The friction coefficient
was fixed to 0.6 according to Byerlee (1967; 1978).
6. Constant heat production rate for the sedimentary cover = 1.0 µW/m3
(Tunini et al., 2015).
7. Constant heat production rate for the basement = 0.4 µW/m3 (Tunini et al.,
2015).
8. Thermal conductivity for the sedimentary cover = 2.0 Wm-1K-1 (Tunini et
al., 2015).
32
9. Thermal conductivity for the basement = 2.2 Wm-1K-1 (Tunini et al., 2015).
10. Heat flux at the Moho ) = 20 mW/m2 (Francois et al., 2014). However,
taking into account that higher values have been proposed for the Red Sea zone
of the Arabian plate (Főrster et al., 2010; Schütz et al., 2014), a value of ) =
25 mW/m2 was also considered in the calculation of the geotherms along our
profile.
Figure 14 – Conceptual model for geothermal calculation along the 2-D geological model (Figure 10,
GH segment in Figure 6, modified from Basilici et al., 2019). The graph describes the parameters
used in the analytical calculation: Crustal thickening values ) are identified by blue numbers,
thrust angles by red numbers and Moho depth values by green numbers.
To perform the 3-D geothermal model we considered a series of twenty pseudo-
wells (Figure 15) located into the 3-D geological model presented in the previous
chapter (the blue rectangle in Figure 6). For each pseudo-well it was considered the
same parameters and assumptions used to perform the 2-D geothermal model but here
it was improved the analytical procedure including altitude (taking into account the
topography by a 30 m resolution ASTER GDEM) and the depth of the Moho
discontinuity calculated by Jiménez-Munt et al. (2012) in which the 3-D geometries of
33
the crust-mantle interface is visible. The thickness of the Arabian cover was considered
as variable based on the 3-D geological model (Figure 12).
Figure 15 – The 3-D geological model (Figure 12, blue rectangle in Figure 6, modified from Basilici
et al., 2020a) with the location of pseudo-wells. The grey pseudo-wells are numbered from 1 to 20.
The section traces (blue lines, HG, PQ, NO, LI) were used to perform the analytical procedure.
4.3 – Analytical Procedure
In the ideal case with no thrusts, we obtain the following equations of heat flow
(1)
(2)
In the hypothesis of a constant heat production rate for the cover ( and the
basement (, the equations of temperature for the two layers are:
34
(3)
(4)
where:
(5)
Assuming a single slip basement thrusting, besides generating a new source of
heat due to friction along the fault, produces a perturbation in the propagation of the
mantle heat and an increase in radiogenic heat within the basement (Basilici et al.,
2019; 2020a; Candela et al., 2015). In the calculation of the heat source represented by
frictional heating associated with thrust faulting, the shear stress () was considered,
which was obtained by Sibson’s (1974) formulation for favourable oriented thrust
faults at the chosen depths under hydrostatic pore-fluid conditions with a fixed friction
coefficient.
The heat flow and the temperatures for the two units of the model are obtained by
the following equations:
(6)
(7)
(8)
where is the average slip rate along the thrust, and is the shear
stress at the thrust depth .
In the case of thrusting, the procedure to compute the temperature follows the
analytical method developed by Molnar et al. (1983):
35
(9)
where is the new equilibrium state, is the final temperature of the new
equilibrium status , and is the time dependent term. Furthermore,
is calculated by the followed equations:
(10)
where
(11)
with
(12)
and is the perturbed initial state, which depend on the heat source type, the
thrust depth , the crustal thickness , the thermal capacity , and diffusivity
. Therefore, the parameter of the perturbed initial is used to compute the
coefficient (Equations (3)), then time dependent term (Equations (2)), and
lastly the time dependent temperature adding the final temperature (Equations (1)). In
particular, the terms and of the equation (1) depend on the thermal
source type and on thrusting involving the sedimentary cover (Basilici et al., 2019;
2020a; Megna et al., 2014) or the basement (Basilici et al., 2019; 2020a; Candela et al.,
2015).
To compute the geotherms the calculation is separated into two parts, one for the
sedimentary cover (z ≤ ) and another for the basement ( ). In the
36
basement the new perturbed conductive status associated with two principal heat
sources – i.e. mantle heat flow (MH) and basement radiogenic heat () – and the
additional heat source due to friction on the fault () are considered. For the
sedimentary cover, the basement radiogenic heat increment at the bottom of this layer,
as a consequence of basement overthrusting in addition to the previous equilibrium
status () and the crustal radiogenic heat , is estimated. Moreover, for this
layer the temperature increases due to friction source (). This approach can be
mathematically summarized as follows:
(13)
(14)
where and are the respective temperatures calculated at the top of the
basement ( ). In particular, , the only not-time-dependent term, is the
temperature due to the radiogenic heat in the cover layer. Considering the
overthrusting, is the temperature associated with the perturbed radiogenic
heat in the basement, whereas and are the temperatures associated
with the new perturbed state of the mantle and of the new heat source due to
overthrusting. Therefore, based on analytical method developed by Molnar et al.
(1983), each time-dependent term has the following generic expression:
(15)
with
(16)
37
In this study, instead, Molnar et al. (1983) equations are modified considering a
constant radiogenic heat for the two layers (Tunini et al., 2015; Basilici et al., 2019;
2020a; table 1).
Without the contribution of the frictional heat, the final temperature for the
sedimentary cover is given by:
, (17)
and the perturbed initial state associated uniquely with the radiogenic heat increase is
given by
(18)
whereas the perturbed initial state associated with the mantel heat is given by
(19)
Considering only the contribute of the basement radiogenic source, the final
temperature for the basement is given by:
(20)
and the perturbed initial state associated uniquely with increased basement thickness
is given by:
(21)
38
(22)
For both units of the model, the perturbation of the mantle-derived heat flow
causes small temperature changes (Basilici et al., 2019; Candela et al., 2015). The
temperature increment due to frictional heating is added as follows:
(23)
(24)
(25)
and in this latter case the final temperature coincides with the perturbed initial state.
The table 1 is the list of the terms , , and for each source type, dividing
the compute of the temperatures for the cover layer ( ), and for hanginwall
( ) and footwall ( ) of the thrust in the basement.
39
Table 1 – Terms of the equations used for the analytical procedure (Basilici et al., 2020a).
40
4.4 – Geothermal Model Results
We calculated surface heat flow , isotherms and geotherms along
the 2-D geological model (Figure 10). In Figure 16a we show the surface heat flow
for a ) = 20 mW/m2 obtained using an interpolation of a second-order power-law
polynomial equation which takes into account the value calculated using five pseudo-
wells. The heat flow increases its value from the point G to the point H due to the heat
flow produced by MFT friction , basement thickening and deepening of
the Moho. The resulting surface heat flow ranges from ~41 to ~ 65 mW/m2 (Basilici et
al., 2019). Into the undeformed Arabian plate the temperature increases constantly,
moving to point H the temperature shows an increment due to thrusting and deepening
of the Moho. The temperature increase from ~400°C at 30 km depth along pseudo-well
1, to ~500°C at the same depth along pseudo-well 5, where the MFT offset occurred at
greater depths (Basilici et al., 2019).
Figure 16 – (a) Surface heat flow for a ) = 20 mW/m2 calculated using five pseudo-wells.
(b) 2-D thermal structure of the Zagros thrust belt for a ) = 20 mW/m2 (modified from Basilici et
al., 2019).
41
In Figure 17 we show geotherms for ) = 20 mW/m2 and ) = 25 mW/m2. With
the same , the absence of thrust is visible in the geotherm relative to pseudo-well 1
that shows a constant increment of temperature with depth. The geotherm relative to
pseudo-well 2 shows a smooth change with respect to that pseudo-well 1 because of
the basement thickening as a result of thrusting. On the other hand, geotherms relative
to pseudo-wells 3, 4 and 5 have higher gradients due to the presence of thrust at greater
depths. Moreover, variations between geotherms for ) = 20 mW/m2 and ) = 25
were computed for pseudo-wells 2,3 and 4, in order to examinate temperature changes.
These variations increase with depth, resulting in a maximum values of ~100 °C at
depth of ~45 km (Figure 17; Basilici et al., 2019).
Figure 17 – Calculated geotherms corresponding to the five pseudo-wells shown in Figure 16 for a
) = 20 mW/m2 (solid lines) and ) = 25 mW/m2 (dashed lines, for the pseudo-well 1 it was not
calculated). The maximum change of temperature between the two geotherms with ) = 20
mW/m2 and ) = 25 mW/m2 is indicated by the dotted grey line (modified from Basilici et al.,
2019).
42
Megna et al. (2014) pointed out that the largest temperature changes are obtained
by varying the heat flow, while temperature changes are in the range of 2 – 5% when
thickness, thrusting depth, thermal conductivity, timing of activity, and slip rate values
vary by 10% of their initial chosen value (Basilici et al., 2019).
We calculated the geotherm for each pseudo-well of the 3-D geological model
(Figure 12). The result is a series of twenty geotherms shown in Figure 18. For a better
visualization of the results we produced four section (HG, PQ, NO, LI on Figure 19)
on the 3-D geological model (Basilici et al., 2020a).
43
Figure 18 – Calculated geotherms corresponding for the twenty pseudo-wells shown on Figure 19.
Each graph contains geotherms relative to the respective cross-sections HG, PQ, NO and LI (modified
from Basilici et al., 2020a).
44
As observed on the 2-D model, the absence of thrust is visible by a constant
increment of temperature with depth (geotherms relative to pseudo-wells 5, 10, 15 and
20 represented by yellow lines in Figure 18). A weaker temperature increase occurred
for the geotherms of pseudo-wells 4, 9, 14 and 19 (green lines in Figure 18), where the
thrust was shallower and implied a smaller amount (3.5 km) of basement offset. On the
other hand, the other geotherms show an increase of temperature variation in
correspondence to the MFT. In particular, the different trend of geotherms depended
on both the basement thickness modified by thrusting and on the thrust depth
(Basilici et al., 2020a). Megna et al. (2014) checked the sensitivity of this type of
analytical procedure, a variation of about 10% of thrust depth or slip rate produces a
change of frictional heat that modified the final temperature by values never exceeding
2%.
We interpolated the temperature data from 50°C to 500°C using a 50°C of step in
order to provide a better overview of the temperature trend along each cross section
(HG, PQ, NO and LI showed in Figure 19). The second-order power-law polynomial
equation used to interpolate the data consisted of best-fit analytical curves with a rms
(root mean square) error comprised between 0.6 and 3.75 (for the 100°C and 500°C
isotherms, respectively; Basilici et al., 2020a).
45
Figure 19 – Isotherms for the 100 – 500 °C temperature range (100°C steps) calculated using a
second-order power-law polynomial interpolation for cross-sections HG, PQ, NO and LI (modified
from Basilici et al., 2020a).
46
The isotherms shown in Figure 19 define a general trend of increasing temperature
moving from SW to NE. Only the isotherm pattern along cross-section HG defined an
upward curvature, the isotherm pattern along the other cross-sections show a gentle
downward curvature, while more straight, SW-ward dipping isotherms characterize
cross-section LI. The deepening of the isotherms involved depth changes in the range
of ~ 20 km for the 100°C and 500°C respectively (Basilici et al., 2020b).
The thermal structure shows a clear correlation with the geometry of the MFT
which is the main thrust fault of the region. According to Tavani et al. (2020), the MFT
developed in the necking domain of the Jurassic rift system ahead of an array of
inverted Jurassic extensional faults, the related structural architecture resembling that
of a crustal-scale footwall shortcut. As the main crustal ramp of the MFT underlies at
depth the MFF, the sinusoidal shape of the latter in the Lurestan region would derive
from the re-use of the originally segmented, inverted Jurassic rift system (Tavani et al.,
2020). Despite this long-lived paleo-tectonic control and the fundamental role
underlying thrust (MFT) controlling it, are recent processes. These are dated to ca. 3
Ma based on stratigraphic information (Homke et al., 2004), and to ca. 5 Ma based on
low-temperature thermochronology data (Koshnaw et al., 2017). According to the
latter studies, the major morphotectonic feature constituting the Mountain Front
Flexure developed during the Pliocene by basement-involved thrusting as deformation
migrated downward at deeper crustal levels. Therefore, late-stage, crustal ramp-
dominated thrusting (Butler and Mazzoli, 2006) not only led to the development of a
prominent geomorphological boundary between the high Zagros Mountains and the
low foothills to the SW (Berberian, 1995; Falcon, 1961; Emami et al., 2010) but also
appears to have exerted a major control on the thermal structure of the Lurestan salient.
Using the minimum curvature spline interpolation technique we constructed a series of
isothermal surfaces, for the 100°C to 500°C isotherms (Figure 20). This technique
shows a good fit with the second-order power-law polynomial interpolation (Basilici
et al., 2020a). The Figure 20 shows how the first-order, general pattern of foreland-
ward deepening of the isotherms was rendered more articulated by along-strike
47
variations, mainly consisting of a SE-ward climbing of the isotherms (Basilici et al.,
2020a).
Figure 20 – Isotherm depth contour maps obtained by spline interpolation for the 100 – 500 °C
temperature range (modified from Basilici et al., 2020a).
48
5 – 2-D Structural Finite Element Models
In order to unravel how and where co-seismic and inter-seismic deformation of
the MFT impacts spatial and temporal patterns of rock uplift of a mountain range, a 2-
D elastic finite element model along the 2-D geological model presented in chapter 3
(Figure 10) was performed (Basilici et al., 2020b). The aim of the finite element model
(FEM) is to simulate inter-seismic stress and strain accumulation, and to obtain
information about the vertical surface displacement associated with both inter-seismic
and co-seismic stages in this region characterized by thrust-related seismicity.
5.1 – FEM Methodology and Previous Studies
The FEM methodology is actually one of the best methods to study complex
systems in various study fields. In geophysics it has been successfully applied and
validated by recent studies in different regions of the world (e.g. Candela et al., 2015;
Carminati and Vadacca, 2010; Liu et al., 2015; Megna et al., 2005; 2008; Vigny et al.,
2009; Zhu and Zhang, 2013). This method consists of the geometric construction of a
model containing the fault setting of interest. To resolve the system, the model was
divided into an equivalent assemblage of small finite elements (elementary components
or mesh). As a result, for each element, a solution was formulated and combined to
obtain the approximated solution for the entire system. The smaller the elementary
components are, the closer the system is to the continuum case, the grater complexity
of the solution becomes. Generally, it needs to research a good compromise between
accuracy, numerical cost and complexity of the studied problem. At the end, a
sensitivity analysis was carried out in order to quantify the solution error.
Several authors applied FEM methodology to understand the structural
implications of the Zagros mountain (Austermann and Iaffaldano, 2013; Francois et
49
al., 2014; Tunini et al., 2015; Vernant and Chèry, 2006). Vernant and Chèry (2006)
used a 2-D FEM to investigate the influence on the Zagros deformation of the obliquity
of convergence, the rheological layering of the lithosphere and a possible weakness of
the MRF. Calculating, as a result, the percentage of accommodation of the MRF.
Austermann and Iaffaldano (2013) performed a global dynamic 2-D FEM of the
mantle/lithosphere system to test the most recent uplift across the Arabia-Eurasia
collision zone, including Zagros orogeny. Francois et al. (2014) produced a series of
2-D numerical models of the Zagros/Central Iran that incorporate free upper surface
erosion, rheological stratification, brittle-elastic-ductile rheologies, and metamorphic
phase change and account for the specific crustal structure of the Arabian and Iranian
continental lithospheres. They tested the impact of the transition from oceanic to
continental subduction and the topographic consequences of the progressive slowdown
of the convergence rate during continental subduction. Tunini et al. (2015) applied a 2-
D FEM model combining geophysical-petrological methodology in order to study the
thermal, compositional, density and seismological structure of the crust and upper
mantle along two transect across the Arabia-Eurasia collision region.
In this thesis we used the 2-D FEM methodology in order to investigate both inter-
seismic and co-seismic deformation related to the MFT along the 2-D geological model
presented in chapter 3 (Figure 10). The chosen software was Marc software (MSC
Software Corporation) of which the first commercial version is available since 1972.
Here it is possible to set up the contact type between two different surfaces (faults on
our instance) and boundary conditions.
The pre-build 2-D geological model (Figure 10) was divided into three
homogeneous domains, to which values of Young’s modulus, Poisson’s ratio and
density were assigned considering an elastic rheology (Table 2). The friction
coefficient was set as µ = 0.6 (Byerlee, 1967; 1978) for the contact (MFT on our
instance). The model was divided into 14,314 quadrilateral element and 15,298 nodes.
Near faults, homogeneous zone boundaries and top surface, the sides of each single
50
quadrilateral element are about 1 km long, increasing in size away from contacts
(Basilici et al., 2020b).
Domain
Formation or Group
Lithology
Young’s
modulus
(Pa)
Poisson’s
ratio
Density
(Kg/m3)
Arabian
cover
Foredeep infill
Asmari, Shabazan, Ilam,
Sarvak, Garau, S-N-B,
Sehkaniyan, Sarki, Baluti,
Kurre Chine, Geli Khana,
Beduh, Mirga Mir – Chia
Zairi
sandstone
limestone
and shale
3.62E+10
0.23
2.55E+03
Arabian
basement
basement
granite
6.00E+10
0.22
2.80E+03
Sanandaj-
Sirjan
Zone
volcanic arc
volcanic
rock
5.16E+10
0.25
2.80E+03
Table 2 – Parameter values used in the FEM model (Basilici et al., 2020b). Density values are from
Teknik et al. (2019; also used in Basilici et al., 2019; 2020a). Values of Young’s modulus and Poisson
ratio were defined according to previous FEM models by Zhao et al. (2004) and Zhu and Zhang
(2013). Formations and groups belonging to the Arabian sedimentary cover are from Tavani et al.
(2018a).
Numerical modelling included two independent FEM simulation carried out using
the same mesh: a first procedure was used to investigate surface vertical motions
associated with the MFT co-seismic stage by modelling the characteristic earthquake,
while a second procedure was used to analyse inter-seismic stress and strain
accumulation (Basilici et al., 2020b). A sensitivity analysis was carried out in order to
quantify the impact of variations of density values and Young’s modulus. Varying the
density from 2.7E+03 to 2.9E+03 kg/m3 (Basilici et al., 2019; 2020a; Teknik et al.,
2019), and Young’s modulus from 4.5E+10 to 5.5E+10 Pa in the Sanandaj-Sirjan Zone
and from 5.5E+10 to 6.5E+10 Pa in the basement (Zhao et al., 2004; Zhu and Zhang,
2012) we obtained uncertainties always smaller than 4%. In particular, the uncertainty
51
varies nonlinearly along the cross-section M-M’ for Young’s modulus variations
within the previously specified ranges, with a mean value of 2% for the central zone of
the transect (175 – 225 km in Figure 10). On the other hand, the uncertainty is almost
constant in the case of the density, with a maximum value of 3.7% (Basilici et al.,
2020b).
5.2 – Co-Seismic Simulation
To produce a realistic simulation of a co-seismic phase of a fault, we need to
select the characteristic earthquake that identifies a seismic cycle. In Figure 21 are
shown earthquakes with Mw ≥ 5.5 recorded from 1967 to the present day in western
Lurestan.
52
Figure 21 (modified from Basilici et al., 2020b) – Earthquakes with Mw ≥ 5.5 recorded from 1967 to
the present day in western Lurestan. Epicentral location, event location, magnitude and fault plane
solutions are from USGS catalogue (https://earthquake.usgs.gov/earthquakes/search/ ; last access: 06
October 2020).
The November 12, 2017, Mw = 7.3 earthquake was selected as the characteristic
earthquake (i.e. a seismic event rupturing the entire fault) to investigate the behaviour
of the MFT. Data on the recurrence interval of similar large-magnitude events
nucleated along the MFT are not available, as catalogue data from the study area start
from 1967 (Basilici et al., 2020b).
53
Modelling of the co-seismic stage was divided into two different steps (Figure
22). The first step consisted in setting the boundary conditions of the model, so as to
observe the effects of gravity alone, it allows the achievement of “equilibrium
conditions” between gravity, hydrostatic pressure of the Earth’s mantle, and
compaction of rocks and contacts. The boundary conditions of the model were the
following:
1. The surface was free to move in all direction.
2. The SW and NE boundary were locked in the horizontal direction and free to
move in the vertical directions.
3. The base was treated as a Winkler’s foundation (Williams and Richardson,
1991). This model base was used to simulate hydrostatic pressure of the Earth’s
mantle: free horizontal movement was allowed and vertical motion was
controlled by an elastic spring with stiffness coefficient equal to
,
where is the base length, is the thickness of the model and is the average
of Young’s modulus of the rocks included in the model.
The second step, implemented in succession, consisted in the sudden movement
of a part of the fault plane, in order to simulate stick-slip behaviour. It considers free
to move (by stick-slip) a portion of the fault located at a depth between 14 and 20 km,
with a total slip of 4 m (as inferred for the November 12, 2017 earthquake; Vajedian et
al., 2018). This procedure allowed us to compute the surface motion in the vertical
direction (Basilici et al., 2020b).
54
Figure 22 (modified from Basilici et al., 2020b) – Geometry and boundary conditions for the co-
seismic simulation. (a) Step 1: the surface is free to move in all directions under the influence of
gravity (applied as volumetric force on each elementary component), while the lateral boundaries are
locked in the horizontal direction and the base of the model is treated as a Winkler’s foundation
(William and Richardson, 1991). (b) Step 2: the boundary conditions are the same as those for step
1, but the fault portion coloured in white is let free to move instantaneously, while the rest of the fault
(coloured in red) is locked.
5.3 – Inter-Seismic Simulation
The inter-seismic simulation was used to analyse inter-seismic stress and strain
accumulations. It shares the first step with the co-seismic simulation procedure,
however, in this case, the second step consisted of the horizontal movement of the SW
boundary towards the NE boundary, in order to simulate observed movements
constrained by the GPS stations located at the surface (Figure 23, Basilici et al., 2020b).
The unique GPS data available in our study area are those recorded by the Ilam
station (Vernant et al., 2004) for the period 1999-2001. The related velocity projected
along our model section results in a value of 2.3 ± 1 mm yr-1 towards the NE,
55
considering a fixed Central Iranian Block (Sanandaj-Sirjan Zone) reference frame.
Therefore, the motion of the SW boundary during the second step was set up to move
Ne-ward with a horizontal velocity of 2.3 mm yr-1 in correspondence with the projected
position of the Ilam GPS station in the model. The second step involves keeping the
seismogenic MFT patch locked (this is the fault that is inferred to have slipped during
the November 12, 2017 earthquake; Gombert et al., 2019; Nissen et al., 2019), together
with the upper crustal fault splays branching out from the upper portion of the main
MFT. On the other hand, the deeper (NE) portion of the MFT was let free to move (by
stable sliding), thus simulating a creeping detachment in the middle to lower crust
(Basilici et al., 2020b).
Figure 23 (modified from Basilici et al., 2020b) – Geometry and boundary conditions for the inter-
seismic simulation. (a) Step 1: the surface is free to move in all directions under the influence of
gravity (applied as volumetric force on each elementary component), while the lateral boundaries are
locked in the horizontal direction and the base of the model is treated as a Winkler’s foundation
(William and Richardson, 1991). (b) Step 2: a horizontal NE-ward velocity of 2.3 mm yr-1 is
introduced in correspondence of the projected position of the Ilam GPS station. The fault portion
coloured in white is that let free to move by stable sliding, while the rest of the fault (coloured in red)
is locked.
56
This procedure allowed us to compute the amount of equivalent Von Mises stress
:
, with
=
(27)
where is the stress component and is the Kronecker delta, of equivalent (Von
Mises) strain
, with
=
(28)
where is the strain component (Zhuang et al., 2019), and of surface motion in the
vertical direction, accumulated in a given time interval during the inter-seismic stage.
The calculation procedure was set up to cover a period of 1000 years in order to reach
appreciable values of accumulated stress, strain and surface displacement.
5.4 – Structural FEM Model Results
The values of resulting vertical surface displacement for co-seismic scenario is
shown in Figure 24. The modelled hypothesis of a characteristic earthquake of Mw =
7.3 yields curves of vertical surface displacement (Figure 24) is characterized by two
peaks. The first peak is positive and shows an increment from 0 to 1.2 m at a distance
around 162 km, while the second peak is negative and shows a decrement to -0.4 m at
ca. 190 km. These values are in very good agreement with the results of the analysis of
geodetic data obtained by Feng et al. (2018) relative to the November 12, 2017 Iran-
Iraq earthquake, which pointed out an uplift of ca. 1 m in the same area and a
57
subsidence of the zone immediately to the NE of 0.4 m (Basilici et al., 2020b). The
values of resulting vertical surface displacement, total strain and equivalent of Von
Mises stress for inter-seismic scenario is shown in Figure 26.
Figure 24 (modified from Basilici et al., 2020b) – Output of the FEM showing vertical surface
displacement for the co-seismic stage (characteristic earthquake of Mw = 7.3).
Figure 25 (modified from Basilici et al., 2020b) – Output of the FEM showing: (a) vertical surface
displacement for the inter-seismic stage (1000 years), (b) equivalent strain and (c) equivalent Von
Mises stress for the inter-seismic stage (zoom of the central part of the modelled section). Black fault
segments are locked, whereas white fault segments are unlocked (i.e. free to slip by stable sliding).
58
The model output of vertical displacement for the inter-seismic stage (integrated
over a 1000 years time span) is characterized by a plateau in the SW part of the model,
increasing from 150 km NE-ward to define a wide bulge between 190 and 240 km and
then gently decreasing to the NE (Figure 25a). A zone of maximum resulting total
strain occurs in the hanging wall of the MFT (Figure 25b), above the deeper portion of
the seismogenic detachment segment. This region of maximum deformation is
comprised between the Sheykh Saleh Fault to the NE and the Miringeh Fault to the
SW, being located just NE of the November 12, 2017 earthquake hypocentre. This
latter falls anyway in a zone of relatively high strain (in the range of 9.0E-05 to 9.5E-
05). It is worth noting that the zone of marked strain accumulation reaches the surface,
maintaining similar values to those characterizing the locked detachment at depth. The
accumulated strain decreases gradually both NE-ward and SW-ward, defining a ca. 140
km wide perturbed area. The Von Mises stress (Figure 25c), besides displaying an
expected peak (exceeding a value of 6.2E+06Pa) at the junction between creeping and
locked detachment segments, is characterized by roughly elliptical, concentric regions
of high stress elongated in a sub-horizontal direction. Stress accumulation at the surface
is not as marked as that occurring at depth (particularly in the 10-20 km range), as
stress diffusion appears to follow a horizontal preferential direction. As a matter of
fact, the perturbed (stressed) region exceeds 175 km in the horizontal direction (Basilici
et al., 2020b).
59
6 – Discussion
6.1 – Geothermal Model
During last years, new modelling methods and software have been developed
allowing us to investigate a geologically complex structure as the “Main Frontal
Thrust” of the Lurestan region, Zagros thrust belt (Iran). GPS measurements show that
the northward relative motion of the Arabian Plate is still active today, with oblique
convergence occurring at a rate of ca. 2 cm/yr with respect to fixed Eurasia. Our
purpose was to reconstruct a 2-D and 3-D geothermal models of the Lurestan arc and
to provide thermal constraints.
To reach our aim, we compared the resulting 2-D model (Figure 16) with the
numerical results obtained by previous studies (Kargaran and Neubauer, 2015;
Shekarifard et al., 2012; Tunini et al., 2015; Vernant and Chéry, 2006) to improve the
description of the thermal structure of Lurestan sector (Figure 26). Figure 26a shows
how an increment of (from 20 to 25 mW/m2) produced a shift of the curve
without changing its shape. Both curves remain within the range of 40 – 80 mW/m2,
which is consistent with recent studies by Shekarifard et al. (2012) and by Kargaran
and Neubauer (2015), but values obtained in this study display a marked variation
with respect to published results by Kargaran and Neubauer (2015), it passes from
lower values in the foreland and thrust front region to higher values in the thrust belt
interior. The curve obtained for a = 20 mW/m2 provides a better fit as
compared with results by Kargaran and Neubauer (2015), in particular, a value
congruency of about 61 mW/m2 is met for pseudo-well 4, while a minor difference
occurs for pseudo-well 5. The curve obtained for a = 25 mW/m2 shows a
congruency in correspondence of pseudo-well 3, but displays a larger difference
towards thrust belt interior.
60
Figure 26 – (a) Surface heat flow for a ) = 20 mW/m2 and ) = 25 mW/m2 compared
with Kargaran and Neubauer (2015) curve. (b) 2-D thermal structure of the Zagros thrust belt for a
) = 20 mW/m2 compared with the isotherms by Tunini et al. (2015; modified from Basilici et al.,
2019).
Figure 26b shows how our temperature isolines are generally consistent with those
obtained by Tunini et al. (2015), a part for some difference that occurs in the thrust belt
interior. The trend of our results is in good agreement with the trend obtained by
Vernant and Chéry (2006; Basilici et al. 2019). Isotherms contour maps obtained by
spline interpolation (Figure 20) allowed us to produce a 3-D model of the thermal
structure of the Lurestan arc subsurface (Figure 27).
61
Figure 27 – Three different views of the 3-D thermal model. The model shows isotherms obtained
using spline interpolation for the 100 – 500 °C temperature range (100°C steps), together with the
MFT and the Moho, this last was calculated by Jiménez-Munt et al., (2012; modified from Basilici et
al., 2020b).
62
The 3-D model of the thermal structure of the Lurestan arc (Figure 27) shows a
good fit with the thermal structure obtained by second-order power-law polynomial
interpolation (Figure 28), thus confirming the consistency and robustness of the
thermal structure 3-D model (Basilici et al., 2020b).
Figure 28 – Thermal structure of the Lurestan arc as defined by isotherms calculated using a second-
order power-law polynomial interpolation along cross-sections HG, PQ, NO and LI (located in Figure
19; black lines are 100, 200, 300, 400 and 500°C isotherms). Surface topography is from the 30 m
resolution ASTER GDEM (modified from Basilici et al., 2020b).
The 3-D model of the thermal structure (Figure 27) shows how isotherms are
characterized by a general SW-ward deepening, which clearly outline a higher
temperature axial zone of the mountain belt. On the other hand a “colder” zone is
present in the outer zone of the SFB and the adjacent foreland, there, the widely spaced
63
isotherm pattern gives way to a pattern characterized by progressively more densely
spaced isotherms, which become closer to each other moving to the SE portion of the
Lurestan arc. The NE-ward temperature increment was due to crustal thickening
associated with the thrust, as the NE dipping MFT produces a progressively increasing
amount of basement offset by the fault . Also, this along-strike, SE-ward
temperature increase was controlled by the 3-D geometry of the MFT, which is
characterized by a low-angle crustal ramp in its SE portion, changing to a steep ramp
rapidly reaching lower crustal depths to the NW. The resulting thermal structure tended
to mimic the general arc-shape geometry of the mountain belt, as readily observable
by comparing the isotherm depth contours with the trace of the MFT in Figure 20. The
junction between the Lurestan arc and the Kirkuk embayment also was clearly reflected
by the crustal thermal structure. Therefore, the geometry of the MFT shows a clear
correlation with the thermal structure of the area (Basilici et al., 2020b).
Tavani et al. (2020) concluded that the sinusoidal shape of the MFT and its
different position in the Kirkuk embayment and Lurestan arc was promoted by lateral
segmentation of the Jurassic rift accommodated by transfer faults. Despite this long-
lived paleo-tectonic control and the fundamental role of structural inheritance, the
development of the MFT and its activity controlling it. Therefore, late-stage crustal
ramp-dominated thrusting (Butler and Mazzoli, 2006) not only led to the development
of a prominent geomorphological boundary between the high Zagros mountains and
the low foothills to the SW but also appears to have exerted a strong control on the
thermal structure of the entire Lurestan arc (Basilici et al., 2020b).
64
6.2 – Structural Model
The results of finite element modelling (FEM), including predicted uplift pattern
and rate along the analysed crustal section, are compared with a tectonic
geomorphology analysis in order to unravel the most reliable tectonic scenario for the
active tectonic setting of the study area.
To produce a cumulative vertical surface displacement curve, we integrated the
calculated vertical surface displacement of both co-seismic and inter-seismic stages
over a time span of 1000 years (Figure 29). As a data on the recurrence interval of the
characteristic earthquake are not available, the reasonable hypothesis of a characteristic
earthquake occurring every 1000 ± 500 years on average has been adopted (Basilici et
al., 2020a).
Figure 29 – Cumulative vertical surface displacement along the trace of cross-section GH integrated
over a time span of 1000 years (modified from Basilici et al., 2020a). It resulting from the sum of co-
seismic and inter-seismic stages considering a characteristic earthquake recurrence interval of 1000
years (blue line); 500 years (green dashed line); and 1500 years (red dashed line).
To improve the description of the rock uplift, we compared the resulting
cumulative vertical surface displacement (Figure 29) with an investigation of the large-
scale features of topography provided by Basilici et al. (2020a) on the same trace GH
of the 2-D geological model (Figure 10). The geomorphology analysis was inferred
from the qualitative and quantitative analyses of the features of the landscape and
drainage network carried out using a Digital Elevation Model (the 30 m resolution
ASTER GDEM), satellite images (Google Earth, 2019) and orthophotos, and a GIS-
based analysis of the DEM. The qualitative analysis of the topography was aimed at
collecting information that could allow identifying the contribution exerted by variable
65
erodibility of outcropping rocks vs. surface motions in the formation of the landscape.
This included an analysis of the active and relic drainage network, carried out through,
e.g. identification and mapping of abandoned fluvial paths and point of capture, and of
erosional and depositional landforms suggestive of a multi-stage development of the
relief. Quantitative constraints to the features of the relief were obtained from the
analyses of the spatial distribution of elevation and local relief and from river
longitudinal profiles. The elevation parameter depends on both resistance to
weathering of outcropping rocks and uplift. However, the maximum elevation is
considered as primarily influenced by resistance to erosion of outcropping rocks, while
the mean elevation is considered as more closely representing a response to surface
uplift (England and Molnar, 1990). On the other hand, the local relief is considered a
robust indicator of uplift, particularly in regions underlain by rocks with rather
homogeneous resistance to erosion (DiBiase et al., 2010). The elevation and local relief
were analysed both along a profile and in map view. A swath profile was constructed
following the methodology of Perez-Peña et al. (2017) along a 320 km long and 40 km
wide transect centered on the trace of the 2-D geological model (Basilici et al., 2020a).
66
In Figure 30 is shown the spatial distribution of topography features compared
with the cumulative vertical surface displacement of Figure 29.
Figure 30 – (a) Spatial distribution of topography features. The red line is the local relief, the black
line is the elevation along the transect and the grey zone represents the spatial distribution of elevation
into the 40 km wide transect. (b) Cumulative vertical surface displacement along the trace of cross-
section GH integrated over a time span of 1000 years (modified from Basilici et al., 2020a). It
resulting from the sum of co-seismic and inter-seismic stages considering a characteristic earthquake
recurrence interval of 1000 years (blue line); 500 years (green dashed line); and 1500 years (red
dashed line).
It is worth noting that the uncertainty involved in the extrapolation of co-seismic
displacements over a 1000 years period is not a major issue in this context, as our focus
is on the definition of the position of areas of major surface deformation rather than on
the absolute value of vertical displacement in Figure 29). Within this framework, a
likely contribution of post-seismic afterslip taking place during years or even decades
following the main shock (Copley, 2014; Copley and Reynolds, 2014) would also be
taken into account by our modelling of surface displacement, assuming slip occurred
along the same thrust originating the earthquake. A different issue is that the reference
timescale of the FEM (102 – 103 years) differs by one – two orders of magnitude from
that involved in the construction of modern topography (106 years; Whipple, 2001).
Nevertheless, the modelled seismic behaviour of the MFT provide useful insights into
the source of the late part of the surface displacement recorded by topographic features
67
and river network, also allowing one to infer whether one or more structures are
contributing to such a displacement. Our results show that the local high relief in
correspondence of the Mt. Bamo (Figure 30) also occurs in the region of maximum
accumulated strain, which reaches shallow depths according to FEM of inter-seismic
deformation (Figure 25b). At ca. 215 km, a further maximum of vertical displacement
in the model occurs, the uplift decreasing then slightly to the NE (Figure 29). Overall,
this entire area of the model defines an uplifted crustal block in the hanging wall of the
mid crustal, gently dipping segment of MFT. The vertical surface motion pattern of
Figure 29 results from rather distinct contributions provided by continuous regional
uplift characterising inter-seismic stages and focused deformation in the area of the Mt.
Bamo associated with co-seismic displacement.
Therefore, our FEM suggests that the Quaternary development of the relief
defining the MFT is mainly related to the co-seismic deformation, while generalized
uplift of the orogen segment located more to the NE is associated with stable sliding
along the deeper portions of the MFT. Consistent with results of numerical modelling
of topography growing over seismically active blind thrusts (Ellis and Densmore,
2006), the two uplifted areas are separated by a belt characterized by co-seismic
subsidence associated with the large earthquake (Mw > 7) occurring in the region
(Basilici et al., 2020a).
68
7 – Conclusions
This thesis provides, for the first time, a comprehensive picture of the thermal
structure of the Zagros fold and thrust belt calculated using an analytical procedure. It
resulted in 2-D and 3-D models that emphasize the major role played by a crustal thrust
fault (MFT) in controlling the thermal structure of the region, a very good correlation
between high temperature gradients in the crust and the amount of basement offset by
the MFT testified it. The study area presents a zone of relatively high geothermal
gradient due to the marked effect of recent (Pliocene) thick-skinned inversion and
associated basement thrusting, this could have interesting implications in terms of
geothermal energy potential (Basilici et al., 2019, 2020b).
The structural finite element modelling presented in this thesis shows how the
vertical surface displacement is the result of a combination of slip accumulated during
large (Mw > 7) seismic events and continuous displacement along a gently dipping, mid
crustal thrust detachment. Co-seismic and inter-seismic stages models allowed us to
gain new insights into the relative contribution of each process in the development of
the relief: while inter-seismic deformation produces a generalized uplift of the whole
crustal block in the hanging wall of the mid crustal segment of the major thrust
detachment, co-seismic displacement controls localized uplift of a distinct topographic
feature located above the main upper crustal ramp of the same major thrust fault
(Basilici et al., 2020a).
The resulted information provided by this thesis is of pivotal importance for any
quantitative modelling of the tectonic behaviour of this actively deforming Iran sector,
geodynamic studies, as well as in geothermal/hydrocarbon exploration. It may also
have important implications for active tectonic studies and seismotectonic modelling
in an area recently affected by a Mw > 7 earthquake.
69
We explained, in this thesis, how the MFT interacts with the entire Lurestan arc
in terms of thermal structure and active deformation. The used mathematical approach,
analytical for the thermal modelling and numerical for the structural modelling, is able
to realize exhaustive 2-D and 3-D models. A similar approach can be used to study
thrust zones in other locations of the world. Santini et al. (2020) applied a similar
analytical procedure to the outer Albanides and adjacent Adriatic crustal sector
(Albania) to calculate surface heat flow, geotherms and isotherms in an area affected
by a strong earthquake (Mw = 6.2) during November 2019. The authors observed a
thermal structure strongly influenced by frictional heating produced by faults, similarly
to Zagros thrust belt.
The FEM method is growing during this years, new powerful softwares are now
in development phase. Thanks to modern technologies, numerical procedures will be
applied, next years, to the 3-D geological models and will be capable to give a better
explanation of the fault activity-active deformation relationship.
To improve this work, studies will be realized in future. We will work to develop
a 3-D FEM modelling using the 3-D geological model (Figure 12) as a base. The 3-D
numerical procedure will be capable of showing the activity of the MFT for its entire
fault’s surface and to highlight the active deformation in a 3-D view, observing the
surface vertical displacement in the map, we will explain how inter-seismic and co-
seismic displacements control uplift of distinct topographic feature in the entire
Lurestan arc. A 3-D geological model will be realized in Albania, to improve,
subsequently, a structural 3-D FEM to understand outer Albanides active deformation
and to observe how co-seismic and inter-seismic displacements control the
development of the mountain chain relief.
70
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Acknowledgments
I would like to offer my sincere thanks to supervisors Prof. Stefano Santini and
Prof. Stefano Mazzoli. I am particularly pleased to them for giving me this study and
work opportunity.
I thank the co-supervisor Ph.D. Antonella Megna for following my work with
patience, great expertise and meticulous attention to detail.
I thank all collaborators from University of Naples “Federico II”: Prof. Stefano
Tavani, Prof.ssa Alessandra Ascione and Ph.D. Ettore Valente; from University of
Camerino: Prof. Pietro Paolo Pierantoni and Ph.D. student Simone Teloni; and from
Total Upstream Denmark: Ph.D. Vincenzo Spina.
I thank the reviewers Prof.ssa Chiara Invernizzi and Prof. Domenico Liotta for
their useful comments and suggestions.
At the end, I would like to express my gratitude to my family and my girlfriend
for their support during the entire Ph.D. course.
