Download 6 2D thermo-mecanical finite element models of n Contents

6
2D thermo-mecanical finite element models of
accretion in the Central Cyprus margin∗
Contents
6.1
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . 136
6.1.1 Morphotectonic features in accretionary margins . . . 136
6.2
From the Cyprus trench to the Central Anatolian Plateau . . . 138
6.3
Modeling accretion in the Central Cyprus margin . . . . . . . 139
6.3.1 Model design and parameterization . . . . . . . . . . . 142
6.3.2 Model strategy and representation of results . . . . . . 143
6.4
Model
6.4.1
6.4.2
6.4.3
6.5
Discussion . . . . . .
6.5.1 A new view at
6.5.2 A new view at
6.5.3 A new view at
6.6
Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158
Results . . . . . . . . . . . . . . . . . . .
Standard model . . . . . . . . . . . . . .
Suite one - sedimentation rate variations
Suite two - viscosity parameters . . . . .
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
.
. 146
. 146
. 149
. 152
. . . . . . . . . . . . . . . . . . . .
uplift in South Turkey . . . . . . .
forearc highs . . . . . . . . . . . .
vertical motions in forearc regions
.
.
.
.
.
.
.
.
. 155
. 155
. 156
. 157
∗ To be modified and partially submitted to Nature Geoscience as: D. Fernández-Blanco, C.
Fuller, M. Utsav, T. Cassola, G. Bertotti, and S. Willett, Explaining the Uplift in South Turkey
133
“Essentially, all models are wrong,
but some are useful.”
George E. P. Box
135
CAP Model
Abstract
Thermo-rheological and sedimentological changes acting during subduction in the
overriding crust influence the landward formation of uplifted plateau-like terrains.
Here, we analyze the impact of these factors on the mechanics of uplift of the
southern margin of the Central Anatolian Plateau. We compiled
tectonosedimentary and geophysical studies from the Cyprus trench to Central
Turkey in an N-S crustal transect, which is used as a constraint to our 2D
thermo-mechanical numerical models. Models show that the growth by sediment
accretion of the Cyprian forearc basin system, let to thermal weakening, lower
crustal viscous deformation and subsequent surface uplift of the modern Taurus
Mountains. The viscous rheology of the overriding plate and the amount of
sediments in the forearc basin control the type and timing of the surface uplift. We
further discuss this thermo-viscous mechanism in the context of other plateau-like
areas.
136
6.1
Chapter 6
Introduction
Subduction margins over the globe can be categorised as erosional or accretionary
depending on the direction of interplate sediment transfer; net crustal loss of sediments and trenchward migration occur in erosional margins, whereas net accretion of
sediments and landward migration take place in accretionary margins [e.g. Clift and
Vannucchi, 2004]. Accretionary margins develop complex trench-parallel structures,
such as an accretionary wedge, a forearc basin system, and a volcanic arc [Gutscher
et al., 1996, 1998; Kukowski et al., 2002] (Fig. 6.1). Whereas the structural development of topography in the seaward areas of the system is related to brittle
deformation caused by the friction exerted by the subducting slab and the landward
accretion of material, the mechanisms of uplift for forearc highs, plateau-like areas
with high elevations that develop inland at some accretionary margins, are open to
debate [e.g. Fuller, 1996; McNeill et al., 2000].
The modern Central Taurus Mountains (CTM) in southern Turkey, an uplifted
area of some 350 km E-W and ∼200 km N-S, can be perceived as the southern flank
of the Central Anatolian Plateau (SCAP). Surface uplift in this region has been
accurately dated as latest Miocene or younger (post–8 Ma) by recent studies, which
attribute the upward motion to break-off of the Cyprus slab [e.g. Cipollari et al.,
2012; Schildgen et al., 2012b; Cosentino et al., 2012; Cipollari et al., 2013; Schildgen
et al., 2014a]. However, these studies fail to explain (i) why the maximum surface
uplift occurs above an area where the Cyprus slab is still presently attached, at least
partially [Bakırcı et al., 2012], and miss to address (ii) the coeval subsidence taking
place in South Turkey offshore [Walsh-Kennedy et al., 2014] or (iii) the presence of
regional contractional structures in this region (chapter 5).
In this contribution, we test the hypothesis that the SCAP formed as the forearc high of the Cyprian subduction margin, effectively fragmenting the preceding
Miocene forearc basin and differentiating the offshore Cilicia and onshore Mut basins
(chapter 5 and Fig. 6.1). We use a series of 2D thermo-mechanical finite element models (FEM) to investigate the mechanics of the surface uplift of the SCAP in relation
to the Cyprian subduction margin. Simulations are constrained by the present-day
thicknesses and geometries, geological observations, and kinematic constraints of an
N-S transect in which we integrated tectonosedimentary and geophysical studies of
the Cyprian accretionary margin (Fig. 6.2). Models include, but are not limited to,
visco-plastic and critical wedge mechanics, isostatic compensation, and the effects
of variations in sediment input, rheological behavior, and thermal conductivity. We
look for an overall agreement on the evolution of the major morphotectonic units
(between the model and nature), focusing particular attention on the uplifted CTM
and its bounding domains to the south and north. We propose applying this approach to other accretionary margins.
6.1.1
Morphotectonic features in accretionary margins
The terms assigned to the different morphotectonic features observed in arc-trench
regions are defined in the following according to how they are used in this contribution. In general, terms to describe natural examples are used after their original
definition as in Karig and Sharman [1975]; Dickinson and Seely [1979], while simpler,
more visual names are used instead for the description of features in the models [e.g.
Fuller et al., 2006a].
137
CAP Model
Trench
Trench-slope
Trench fill
s.l. basin
Trench-slope
break
Accretionary
Forearc-high
Forearc basins
Residual
?
Subductin
area of mechanical
accretion
Not to scale
Messaoria
Basin
Troodos
Ophiolite
Levantine
basin
s.l.
?
Overriding plate
Kyrenia
Range
Cilicia
Basin
N
Central Taurus
Mountains
No
volcanic arc
Mut Basin
?
Cyprus S
la
area of mechanical
accretion
Not to scale
2D Models
Proside
Oceanic
?
(CVP)
Anatolian plate
Structural
Highs
Wedge-top
basin
lithosphe
Not to scale
?
unknown mechanism
b
Proside
basin
?
unknown mechanism
g plate
Cyprus Arc S
Volcanic arc
Intramassif
re
Forearc-high
Forearc
probasin
Forearc
retrobasin
Retroside
Generic forearc
Accretionary wedge
Mechanical domain
Figure 6.1: Terminology of the features of a conceptual accretionary margin with forearc
high and their correlation to the main features seen in our study area and those in our models. Generic terminology is used in the same manner to Karig and Sharman [1975]; Dickinson and Seely [1979]. The terms of “trench-fill basin” and “accretionary forearc basin”
are referred to in the model description as “pro-side” and “wedge-top” basins, respectively,
and the terms “residual” and “intramassif forearc basins” of Dickinson and Seely [1979]
are called, during the modeling description, as “landward” and “seaward” forearc basins,
respectively.
From the subducting slab landward, accretionary systems are composed of: (i)
the trench-fill basin, which is a sedimentary infill at the front of the trench, on
top of the contact point between the subducting and the overriding plates; (ii) the
trench itself, an area grown by active deformation that includes the trench-slope,
shaped by imbricate seaward-verging thrust faults; and (iii) the trench-slope break,
in which the slope changes to landward-dipping in response to the appearance of
landward-verging thrusts. The trench region might contain accretionary forearc
basins inDickinson and Seely [1979] sense that are often carried on top of thrust
sheets as piggy-back basins. Inland of the trench-slope break, (iv) forearc basins
138
Chapter 6
might appear as residual forearc basins located between the trench and the arc
massif, as intramassif forearc basins with sediments that unconformably lie on the
arc massif, or as a constructed basin, a large single forearc basin (not shown in
Fig. 6.1) enclosing the two previously described types [Dickinson and Seely, 1979].
Before reaching the volcanic arc, which delimits the forearc from the backarc regions,
(v) a forearc high might be present, separating the residual and the intramassif
forearc basins or enclosing them within the intramassif forearc basin (Fig. 6.1).
In the models, simpler nomenclature is used. Convex-up features that develop in
relation to shear zones are termed “structural highs”. Basins commonly appear enclosed between adjacent structural highs and are named based on their location with
respect to the pro- and retro-sides of the models (sea- and landward, respectively)
and the area of active deformation. Developing at the pro-side of the model, (i) proside basins and (ii) wedge-top basins [DeCelles and Giles, 1996] are differentiated
on the basis of their distinctive wedge-top position for the latter, or lack thereof for
the former. Heading to the retro-side of the model, a well-developed (iii) structural
high separates the pro-side area of active deformation (shear zones) from a stable
area, where the (iv) forearc basin develops. If present, the (v) forearc structural
high divides the forearc basin in pro- and retro-side forearc basins at the front and
back of the forearc high, respectively.
The term “negative-alpha basin” is used for model and natural cases alike, in the
Fuller et al. [2006a] and Willett and Schlunegger [2010] sense; this is, a basin with a
surface slope reverse to that of the wedge on top of which the basin remains stable
with no internal deformation, and passively sliding above the subduction thrust, as
long as it is restricted and steadily infilled by its bounding highs.
6.2
From the Cyprus trench to the Central Anatolian Plateau
A regional transect that runs north from the East Mediterranean to the Central
Anatolian Plateau interior is shown in Fig. 6.2. This crustal-scale transect encloses
relevant tectonosedimentary and geophysical data from on- and offshore studies.
This integrative effort, unique for the area, allows the determination of four key
features; (i) the shape of the subducting slab and the geometry of its contact with the
continental crust, (ii) the overall distribution of crustal thicknesses, (iii) the relative
age and cutting relationships of the main faults, and (iv) the position, geometries
and continuity of the Miocene rocks.
The transect shows three first-order tectonic features (Fig. 6.2). To the south, the
subducting African plate is represented by a continental fragment, the Erastothenes
Seamount, and farther north by oceanic crust, starting at >15 km depths from
around 35°and northward. The Troodos Ophiolite in the center of the transect is
a sliver of oceanic crust in between plates. Toward the north of the transect, the
Anatolian overriding plate consists of thickened continental crust.
Crustal thicknesses along the transect range from a minimum of ∼25 km to
a maximum of approx. 45 km. In the southern sectors of the transect, thickness
changes are well detected by the gravimetric signal of Ergün et al. [2005] (Fig. 6.2B) and the Moho models of Koulakov and Sobolev [2006] (Fig. 6.2-C). In the African
plate, average thicknesses of ∼28 km are observed below the Erastothenes Seamount
south of Cyprus. More to the north, between 34°30’ and 37°N (Fig. 6.2-C), the geophysical models concur on a significant increase in Moho depth, from some 28 km
CAP Model
139
to more than 40 km, which we correlate with the steepening (up to 40°) of the subducting slab (Fig. 6.2-E). The thinnest oceanic crust (∼25 km) is seen below the
trench area. Northward, thickening occurs in relation to the Troodos Ophiolite, the
detachment depth of which is uncertain. This crustal thickening is probably the
result of the Troodos emplacement, with probable thrust doubling. This is also the
location underneath which the locked underthrust of the Erastothenes Seamount is
found. The extent of the continental crust underneath the Troodos Ophiolite and
the position of its transition to oceanic crust more to the north remain enigmatic. In
the southern regions of the overriding Anatolian plate, a relevant crustal thickening
takes place below the CTM. Farther north, the Anatolian crust thicknesses decrease
gently from circa 45 km to values in excess of 35 km. Here, Pn tomography was
used [Mutlu and Karabulut, 2011] instead of gravity data [Özeren and Holt, 2010]
(Fig. 6.2-C), which points to crustal thickness values up to 10 km thicker.
South-verging thrusts led to fragmentation of the area into present-day sedimentary basins and structural highs (Fig. 6.2-D). These thrusts have been linked either
directly (in the south) [e.g. Stephenson et al., 2004; Calon et al., 2005a, b], or indirectly (in the north) (Chapter 5), with the Cyprian subduction thrust and become
older when progressing northward along the section (Fig. 6.2); they are as young
as present-day at the trench [e.g. Stephenson et al., 2004], Pliocene or younger in
north Cyprus [e.g. McCay, 2010], mid-Pliocene in the Cilicia Basin (Chapter 5),
and pre-Miocene in the Mut Basin [e.g. Çiner et al., 2008]. Pliocene contractional
structures are seen in the Cilicia Basin, but no evidence of any Miocene or younger
thrust system is known for the Mut Basin or for the transition between both basins.
This transition is marked instead by a long-wavelength monocline [Çiner et al., 2008]
that accommodates ∼4 km of relative vertical displacement between initially laterally equivalent Miocene rocks (Chapter 5).
From south to north, Miocene rocks outcrop in four different localities: the
Messaoria Basin in Cyprus, the offshore Cilicia Basin, the Mut Basin on the CTM,
and the Tuz Gölü Basin, in the plateau interior. Originally deposited on top of a
peneplaned surface, these rocks allow the determination of the relative end-result
vertical motions since the Miocene. The Miocene basins seen along the transect
show dissimilar geometries. Rocks above pre-Miocene substratum thin southward in
the asymmetrical Messaoria Basin [Calon et al., 2005a, b; McCay, 2010]. Messinian
and younger deposits of the Cilicia Basin also thin southward, but the geometry
remains speculative for pre-Messinian Miocene, since its bottom is not seen in the
seismic profiles of the area [e.g. Aksu et al., 2005a]. This asymmetry is not seen for
the Mut Basin nor farther north. Basin terminations are erosive in all four Miocene
basins, with the exception of the southern margin of the Messaoria Basin.
6.3
Modeling accretion in the Central Cyprus margin
Our modeling analysis is suitable to describe the central segment of the Cyprus arc,
which is normal to the convergence direction and is where contractional structures
develop. Our models simulate a transect running roughly S−N, which is parallel
to the convergence and the main transport direction of material coming from the
Cyprus arc (Fig. 6.3). However, other relevant sources of material are the mountainous areas to the north and northeast. This and other important three dimensional
effects that might occur in the area [e.g. Wortel et al., 2010] are not taken into
account.
B
C
mGal
Turkey
0
-100
0
200
Erastosthenes
East
Mediterranean
Cyprus
CVP
Turkey
Central Anatolia
0
-100
30
20
Figure 6.2: Central Cyprus geologic and geophysical transect (I)
(a) Map view, Bouguer gravity analysis and cross-sectional representation of geophysical data available for the study area.
~110 km
Vertical exaggeration ~ 3,5 50
Crustal thickness via Pn tomography, as in Mutlu & Karabulut, 2011
Moho depth via real data inversion, as in Koulakov & Sobolev, 2006
Moho depth via seismics + gravity, as in Özeren & Holt, 2010
Moho depth CRUST 2.0, as in Özeren & Holt, 2010
0
200
10
Moho depth CRUST 2.0, after modification by Koulakov & Sobolev, 2006
Reflaction seismics, as in Mart & Ryan, 2002
Bouger gravity analysis, as in Ergun et al, 2005
Aeromagnetic survey MTA - Interpreted by Ates et al., 1999
Taurus Range
50
Cilicia Basin
40
Ergun et al., 2005
Trench
Northeast Mediterranean
40
30
20
10
Km
mGal
Km
A
140
Chapter 6
50
Volcanics
Oceanic crust
Troodos Ophiolite
Miocene
Paleogene
Plio-Q
Messinian
Mantle lithosphere
Continental crust
Figure 6.2: Crustal-scale transect running N-S (650 km at around 33°30’ E). The transect stretches from the Levantine Basin to the Tuz
Gölü Basin, passing though the Cyprus arc trench, Cyprus island, the Cilicia Basin, and the Taurus Mountains, finally reaching the Central
Anatolian interior. Panel A represents a map view of the area. Panel B shows the two major gravimetric studies in the area [Ates et al.,
1999; Ergün et al., 2005]. Panel C shows the interpretation of the offshore section c of Ergün et al. [2005], the interpretation of the seismic
study performed by Mart and Ryan [2002], and several cross-sections constructed from depth map models of the Moho [Koulakov and Sobolev,
2006; Özeren and Holt, 2010] and Pn tomography [Mutlu and Karabulut, 2011]. The circles are focal epicenters with Mw > 5 recorded in
a longitudinal area from 32°30’ E to 34°E in red and 31°30’ E to 34°30’ E in green. Panel D is the upper sector of the transect (up to
10 km depth), shown with ∼ 7 km of vertical exaggeration for the sake of visualization, and it integrates this Thesis and upper crustal studies
[Robertson, 1998b; Stephenson et al., 2004; Calon et al., 2005a, b; Çiner et al., 2008; McCay, 2010; Fernández-Blanco et al., 2013]. Panel E
is the crustal cross-section for the area, integrating all the aforementioned inputs.
(b) Continuation from the previous page. Upper crust and whole crust sections of the study area.
Eratosthenes SeaMt
Levantine Basin
40
~110km
Vertical exaggeration ~ 2
50
39º
40
38º
30
30
37º
20
20
36º
10
10
35º
0
0
34º
~110km
Vertical exaggeration ~ 7 10
10
Turkey
5
5
Km
Km
East
Mediterranean
B
0
0
Km
Km
A
CAP Model
141
142
6.3.1
Chapter 6
Model design and parameterization
Accretion mechanisms at accretionary margins are often described using the critical
wedge theory [Davis et al., 1983; Dahlen, 1984; Larroque et al., 1995; Malavieille,
2010], which defines the geometry of the orogenic wedge as a function of the mechanical properties of the accreting wedge and that of its surface of accretion. However,
the critical wedge excludes the ductile properties of these systems, such as the absence of brittle fracturing in their lower ends, which cause relevant deviations from
the predicted brittle morphologies. Recent research inclusive of the visco-plastic
attributes show the influence of thermal or rheological variations or that of sediment load and/or competence in the strain distribution and deformation patterns
within the accretionary wedge [e.g. Fuller et al., 2006a; Simpson, 2010; Willett and
Schlunegger, 2010; Fillon et al., 2013].
Our 2D kinematic-dynamic models consist of two coupled domains (Fig. 6.3).
The domain where mechanical laws apply represents the crust of a deforming subduction zone. Accretion of incoming sediments is driven by the tangential velocities
at the base of the mechanical domain. These velocities decrease toward, and become
zero at, the “S” point, which represents the point of contact of the subducting slab
and the continental Moho. The thermal domain covers the whole model, including
the mechanical domain. For a detailed description of the model formulation and
boundary conditions, the reader is referred to Chapter 2, sections 2.4.1 and 2.4.2.
The model is consistent with time frame and cross-sectional lengths of interest
and simulates 25 Ma of subduction in a transect of 550 km (Fig. ??). Parameters
such as amount of material incoming at the trench and convergence velocity are
set constant during the simulations. We adopted values of 3 km and 35 mm/y,
respectively, which are considerably lower than present for the incoming material at
the trench and higher than present for the convergence velocities. Our choice is based
on the needed extrapolation in time (25 Ma) of present-day values. The presentday sedimentary thicknesses in the East Mediterranean Sea range from 10 km to
15 km [e.g. Makris and Stobbe, 1984]; these values are surely the largest along the
time frame of interest, considering factors such as the narrower confinement of the
present-day Mediterranean or the presence of the Nile. Similarly, the underthrust
of the Erastothenes Seamount below south Cyprus presently decelerates subduction
motion to 9.3 ± 0.3 mm/y [Reilinger et al., 2006], thus a larger, more common,
convergence velocity value is given for the time considered for our simulations.
The subducting lithosphere is 50 My old at the left side of the model and its
70 km thickness remains constant during the running time. Since thicknesses in the
mechanical domain, which represents the crust, change as accretion takes place, an
initial thickness of 30 km is chosen on the basis of the similarity between thicknesses
in nature and in models after the run is completed, i.e., maximum values of 45 km
near the “S” point. The rest of the overriding lithosphere is 80 km thick.
Flexural rigidity is set at 2.4 x 1023 N · m for both plates (after Fuller [1996]).
Variations of this value of up to four orders magnitude did not produce substantial
changes [e.g. Forsyth, 1985, for a discussion on flexural rigidity values]. Densities
are commonly accepted values: 2.8 g/cc for the lithosphere (which includes the sedimentary cover), 3.3 g/cc for the mantle and 1.03 g/cc for the overlaying layer of
water. Cohesion and internal friction angles control the mechanical strengths in our
model. Cohesion, c, is set to 1000 Pa, a value higher than expected for the crust,
143
CAP Model
da = Incoming
thickness
A
Vx = Vc
2 Km
Crust - Mechanical domain
Vc = 10 km/Ma
Ds = 2,4 x 1023 N·m
Vtg
0
S
Fully Coupled Thermo-mechanical model
B
C = 1000 Pa
n = 3,0
Ao = X x 10 -19 Pa-3 . Ma-1
Q = X x 105 J . mol-1
φ = 24o
φo = 8o
Vtg = Vc
Vtg = 0
Do = 2,4 x 10 N·m
Full Model - Thermal domain
Crust
V = Vc (kinematic)
Oceanic
3
Lithosphere H = 0 mW/m o
T = 0 oC
V = Vmech H = 1,25 mW/m3 K = 2,0 W/(m oK) C = 1000 J/(Kg oK)
p
S
K = 2,0 W/(m K)
Cp = 1000 J/(Kg oK)
dT = 0
dx
Vx,y = 0
23
Asthenosphere
V=0
H = 0 mW/m3
K = 50 W/(m oK)
Cp = 1000 J/(Kg oK)
T = 1400 oC
Continental
Lithosphere
V=0
H = 0 mW/m3
K = 2,0 W/(m oK)
Cp = 1000 J/(Kg oK)
dT = 0
dx
Asthenosphere
T = 1400 oC
Figure 6.3: Model setup, with indication of thermal and mechanical parameters.
to maintain model stability. Lower values do not affect the outcomes [Fuller, 1996].
The internal friction angle of the crustal material, φ, is set to 27◦ and the friction
angle between the subducting and the overriding plates, φb , to 8◦ . Friction values
are set low to include the effect of fluid pressures not explicitly taken into account
and imply fluid pressure ratios within the range of those at accretionary wedges
[Fuller, 1996, and the references therein].
To let the thermal structure equilibrate, the thermal model runs for 20 My before
the crustal model onsets. Surface temperature is 5◦ C, an average between subaereal
and subacuatic temperatures, and a value of 1400◦ C is given for the asthenosphere
at the base of the model. We use values for thermal conductivity of 2 and 50
W/(m · ◦ K) for the lithosphere and the asthenosphere. Asthenospheric conductivity values are readably high to represent isothermal conditions. Heat production
has a value of 0.85 µW/m3 [Jaupart and Mareschal, 2005], occurring only in the
mechanical domain. Specific heat, c, is 1200 J/kg ◦ K for both model domains.
6.3.2
Model strategy and representation of results
The constraints presented above are used to study the mechanics of uplift of the
southern margin of the Central Anatolian Plateau. A standard model, which best
represents the known geologic evolution for the central Cyprus arc, is chosen attending to the overall agreement of large-scale structure evolution and vertical motions
through time, as well as final geometry. Based on our preferred model, the values
of sedimentation rate and viscosity for the parameter space studied are obtained as
outcomes. Then we estimate the sensitivity of the model to the influence of each
parameter using two suites: (i) Suite One for the sedimentation rate, and (ii) Suite
Two for the viscosity, where we study the activation energy, Q, and the power-law
viscosity coefficient and exponent, Aµ and nµ .
144
Chapter 6
Representation of the morphotectonic features
We intend to quantify the rates of the vertical motions and the geometrical variations
within and around the forearc high before and during its development. Changes in
the rate of the motions of the different components of the forearc basin system and
its resulting evolution can be traced over time following the position of the upper
limit of the basement, i.e., the surface representing the top of the mechanical domain
at the moment of model initiation. We track the position over time of the lowest
and highest points of the top of the basement for the area in and around which the
forearc high develops to track the development of basins and forearc highs and their
change through time (Fig. 6.4). In this schematic representation, a sedimentary
basin is confined between two highs and a structural high is enclosed between two
lows.
The width of the basins is the distance, measured on a straight line, between
its highest bounding points (B0 W for time step 2, or BP W and BR W for time
step 3 in fig. 6.4-AB). Similarly, the width of the structural highs is defined as the
distance on a straight line between its bounding lowest points (F W for time step
3 in fig. 6.4-AB). The depth/height of each feature is defined differently depending
on the morphotectonic features present at every time step. Before the development
of the forearc high (time step 2 in fig. 6.4-AB), the depth of the forearc basin is
the vertical distance between the deepest point of the basin and its position at the
moment of model initiation (B0 D for time step 2 in fig. 6.4). After the onset of the
forearc high growth (time step 3 in fig. 6.4-AB), and in order to better represent the
relative vertical motions of the forearc high and that of its frontal and rear basins,
a new base level change takes place at every time step. Restitution of the base
level occurs with respect to an arbitrary point, which coincides with the landward
boundary of the forearc basin system at every time step. This point is defined with
a deliberate increase of initial model topography in the retro-side of the model, and
continuously subsides at a similar pace to that of the area where the forearc high
develops. The height of the forearc high (F H), and the depth of the pro- and retrobasins (BP D and BR D, respectively), are measured with respect to the new base
level at every time step. The base level change removes (at least part of) the primary
subsidence signal and allows the quantification of the relative motions of the forearc
high and its bounding basins. It also allows the distinction between the growth of
(i) a subsiding high, a feature that subsides less than its surroundings but remains
at deeper absolute levels than the reference point, and that of (ii) an uplifting high,
an uplifted area that overcomes the primary subsidence and is then treated as the
forearc high. Only the case of the uplifting high (forearc high) is represented in the
plots.
We anticipate that the first-order variations observed in our models would will
occur in any accretionary margin (Fig. 6.4). Therefore, the measured widths and
depths/heights of each morphotectonic element have been converted into percentages
in relation to the maximum basin depth and width of the standard model, and are
depicted as the base and altitude of the triangles, respectively. The maximum basin
depth accounts for the depth of the initial forearc basin and the depth of the forearc
pro-basin, whereas the maximum basin width corresponds with the width of the
initial forearc basin (see Fig. 6.4). This representation allows for easy comparison
between different time steps of the same model and among different models.
Time-step cross-sections
of top of the basement
Forearc high
heigth (FH)
Time-step
2
Base level before onset
forearc high formation
B PW
B0W
BRD
Forearc
retrobasin
width (BRW)
Forearc high
width (FW)
Forearc
initial basin
width (B0W)
Forearc
probasin
width (BPW)
BRD
FW
BRW
l ch
a
Base level after onset
forearc high formation
Time-step
3
ve
FH
10 km
20% max. width
3 2
Time-steps
Base le
D
C
Maximum basin depth of standard model
5 km
0 Ma
20%
5 Ma
10 Ma
15 Ma
20%
15 Ma
10 Ma
40%
5 Ma
0 Ma
60%
80%
20 Ma
Height/depth of the highest and
deepest points of the basement at 5 Ma,
with respect to the reference point
(after forearch high formation)
Depth of the deepest point of the
basement at 20 Ma
(before forearch high formation)
Position of basement at 25 Ma
(considered zero)
Reference point
(changed to new position evey time -step after forearc high formation)
Maximum basin width of standard model
50 km
20 Ma
Reference surface
(used every time-step before forearc high formation)
25 Ma
Standard Model
Figure 6.4: Example of a simple representation of the forearc basin system evolution (left) and simple representation of the evolution of the
standard model (right). The upper panels represent the evolution of the basement top surface for a hypothetical case (left) and our standard
model (right). The lower panels show the proposed representation of the upper cases. Basins are shown as normal triangles, isosceles are used
for the initial forearc basin, a right triangle looking left is used for the forearc pro-basin, and a right triangle facing right is used for the forearc
retro-basin. The structural highs are represented as inverted isosceles triangles only if they grow pass the chosen base level. For the right
panels, the different time steps are shown as different colors; 20 Ma in orange, 15 Ma in green, 10 Ma in blue, 5 Ma in brown, and present in
red.
Simple representation of the width and depth evolution
of the morphotectonic elements of interest
B0D
B
2 3
Forearc
initial basin
depth (B0D)
Reference Example
20% 20%
40%
Time-steps
Forearc
probasin
depth (BPD)
e
ng
Forearc
retrobasin
depth (BRD)
20% max. depth
60%
km
80%
A
CAP Model
145
146
6.4
Chapter 6
Model Results
In the following, we describe our standard model and analyze the overall geometrical
and evolutionary changes between the standard and the rest of the models.
6.4.1
Standard model
A sedimentation rate of 0.5 mm/yr and viscous parameters values of Q = 1.7 J/mol,
Aµ = 2.05 Pa−n /Ma, and nµ = 2.65 reproduce our preferred model. The complete
list of parameter values used in the standard model is shown in Table 6.1. Corresponding to the crust, the mechanical domain of the standard model is shown
(mostly in green) in fig. 6.5.
At the moment of model initiation (25 Ma) we can observe the geometry of the
mechanical domain. A roughly flat topography dominates the area, with the only
exception of the topography defined at the retro-side of the model, at the right hand
of figure 6.5. At the left-hand of this figure, the curved area below the mechanical
domain represents the down-going slab.
As seen at 15 Ma, topography has started to develop in the left-hand side of the
model (pro-side). This topography is seen in relation with two shear zones, marked
in warmer colors than their surroundings. A small asymmetric wedge-top basin
developed between the shear zones, which are seen as white horizontal isochrones.
This basin is translated between the two shear zones as they migrate retroward
(toward the right). At the same time, a large forearc (negative-alpha) basin develops
in the center and retro-sides of the model. The subducting slab becomes steeper as
the base of the crust becomes more horizontal. Both factors influence the extent of
strain around the “S” point at the base of the crust.
By 10 Ma, the wedge-top basin migrates retroward and the shear zones become
wider and more pronounced. Further deepening and thickening in the forearc basin
leads to increasing temperatures and strain accumulation at the base of the crust.
As the sediments accumulate in this forearc basin, the initial formation of a frontal
branch of a shear zone occurs at the “S” point and basement-basin contact begins
to dip toward the left (proward).
At 5 Ma, two small basins have developed in the pro-side. Toward the centralright area of the model, two shear zones branch out from the “S” point and produce
an uplifted broad bulge. These shear zones clearly delimit the forearc high uplifted
area, which dips more markedly toward the left-hand side of the model. The basement further tilts proward (left), and the formation of a monocline is clearly seen,
with horizontal deposits present on top of the bulge and dipping away from it to
its sides. Onlap relationships take place on both sides of the bulge that are more
angular in the pro-side. This bulge has divided the forearc basin into a pro-side
negative-alpha basin to the left and a retro-side landward basin to the right; it continues leading the development of onlap relationships to its sides. The deflection of
the slab proceeds toward higher angles (∼40°) at the “S” point and the base of the
crust evolves to smaller ones.
Since 5 Ma, the system evolves in a self-similar manner. The basement of the
pro-forearc basin deepens while further uplift takes place in the forearc high, in
which an area with very low values of strain is now seen in the uppermost ∼5 km.
147
CAP Model
10 km
20 km
10 km
20 km
10 km
20 km
10 km
20 km
10 km
20 km
Turkey
East Mediterranean
Cyprus Cilicia
Basin
Taurus
Mountains
CAP
Interior
Figure 6.5: Mechanical model evolution in time steps, with zoom-ins into the area where
the forearc high develops. Times represented are, from top to bottom:25 Ma, 15 Ma, 10 Ma,
5 Ma, and present. The individual lines on top of the basement represent isochrones and
should reflect the overall geometric relationships expected for strata.
148
Table 6.1: Principal parameter values of the standard model
Parameter
Description
Value
Mechanical model
̺c
̺m
̺w
φ
φb
c
D
Sα
νc
h
Sedr
Lithospheric density
Mantel density
Water density
Internal friction angle of crustal material
Friction angle between subducting and overriding plate
Cohesion
Flexural rigidity (both plates)
Slab angle at S point
Convergence velocity
Sediment thickness in the subducting plate
Sedimentation rate
2800 kg/m3
3300 kg/m3
1030 kg/m3
27◦
8◦
1000 Pa
2.4x1023 N · m
40◦
35 mm/yr
3 km
0.5 mm/yr
Thermal model
Thickness of the subducting lithosphere
Thickness of the overriding lithosphere
Age of the subducting lithosphere
Thermal model runup
Thermal conductivity (lithosphere, asthenosphere)
Surface temperature
Asthenosphere temperature
Heat production
Specific Heat
Power-law viscosity cohefficient
Power-law viscosity exponent
Activation energy
70 km
80 km
50 My
20 My
2.0, 50.0 W/(m ·◦ K)
5◦ C
1400◦ C
0.85 µW/m3
1200 J/kg ◦ K
2.05 Pa−n /Ma
2.65
1.7 J/mol
Chapter 6
Hs l
Ho l
M al ith
tr unup
Kl , a
Ts
Ta
A
c
Aµ
nµ
Q
149
CAP Model
6.4.2
Suite one - sedimentation rate variations
Recent thermo-mechanical modeling studies have focused on the impacts of sedimentation rate on the structural evolution of orogenic wedges and have successfully
shown that sedimentation rate strongly affects the stability of the wedge and the
distribution of strain within these systems [Fillon et al., 2013; Cassola, 2013]. Building upon this concept, we analyze how sedimentation rate influences the thermal
distribution underneath and within orogenic wedges, and thus affects their brittle
or ductile behavior, resulting in differential strain mode and strain localization.
In our models, sedimentation is “fill-to-spill” in that we assume that sufficient
sediment is available and that areas capable of accommodating sediments are filled to
capacity, i.e., areas between structural highs are completely filled. In this section, we
study the effects that variations in sedimentation rate produce on the development
of the different morphotectonic elements of the margin.
Table 6.2: Sedimentation rate parameterization
Model
Sedimentation rate (mm/yr)
Standard
Sr1-to-Sr14
0.5
0 to 1.4 (∆0.1)
Variation in Sedimentation Rate
0.1 mm/y
15 Ma
0.9 mm/y
10 km
5 Ma
10 km
10 km
0 Ma
10 km
10 km
10 km
Figure 6.6: Forearc high development with changes in sedimentation rates. The figure
shows 15 Ma, 5 Ma, and 0 Ma snapshots for the models with 0.1 mm/y (left column) and
0.9 mm/y (right column)
150
Chapter 6
As sedimentation rate increases, strain becomes greater in and around the “S”
point and is distributed on both sides away from it. An increase in sedimentation
rate also promotes the development of a second bulge in the retroward end of the
model. The subducting lithosphere becomes steeper retroward (to the right) and
the deflection of the base of the crust increases toward the “S” point. That is,
the “S” point increasingly subsides with increasing sedimentation, thus the surface
uplift per se (with global reference frame) decreases. However, with respect to
the “S” point only, or any fixed surface, such as the sea level, a more unexpected
behavior is observed. Starting from the model with no sedimentation, we notice that
larger sedimentation rates produce larger surface uplift. If we continue to increase
accumulation rates, surface uplift eventually decreases until a threshold is reached,
after which surface uplift becomes larger again (Fig. 6.7).
With low sediment accumulation rates, the initial forearc basin is thin, which contributes to a surface uplift that occurs earlier with respect to model onset (i.e., older),
and both upward and downward vertical motions take place over shorter periods.
In general, with lower accumulation rates the margin is more susceptible to the
influence of the other parameters; for example, the incoming sediment at the trench
highly dominates the structural evolution of the margin, developing sharper shear
zones that migrate retroward faster and structural highs and lows that are larger
in amplitude. As sedimentation rates increase, the overall relative vertical displacement between subsided and uplifted areas is consistently larger and occurs in narrower (horizontally shorter) distances, i.e., there are larger amplitudes in the main
subsidence-uplift pair (forearc high–pro-basin). Another effect from increasing the
accumulation rates is the development of thicker forearc basins that stabilize the
margin; these force deformation outside the basin into its margins and farther away.
The surface uplift becomes smaller (Figs. 6.6 and 6.8) and eventually stops happening. The uplifted and subsided areas become broader, and the points of depocenter
and maximum uplift move landward. Similarly, as the sedimentation rate increases,
sedimentary angular discordances develop more commonly instead of onlap relationships, i.e., the vertical displacements occur more gradually.
Variations in sedimentation rates influence virtually every aspect of the growth
of the margin. When sediment input is small, the development of structural highs
along the margin takes place at high rates and at early stages of the simulations. A
broad forearc high develops with very small flanking basins (Fig. 6.8, left-hand). As
sedimentation increases, a long wavelength subsidence takes place at model initial
stages and the onset of the forearc high formation is delayed (Fig. 6.8, right hand).
Eventually, with higher sedimentation values, the forearc high formation stops and
only the long wavelength subsidence is observed. We can therefore conclude that
the forearc high is a common feature developing in accretionary margins, but its
development is gradually halted by the long wavelength subsidence imposed by the
sediments.
151
CAP Model
Top of the basement
Sr14 = 1.3 mm/y
Sr11 = 1.0 mm/y
Sr8 = 0.7 mm/y
Sr6 = 0.5 mm/y
Sr4 = 0.3 mm/y
Sr1 = 0.0 mm/y
Bottom of the crust
Figure 6.7: Line trace of the bottom of the crust and top of the basement in models with
changes in sedimentation rates after a complete run. All models share a fixed “S” point for
easy comparison among them.
Sed. rate = 0.1 mm/y
25 Ma
10 Ma
Sed. rate = 0.9 mm/y
Maximum basin width of standard model
60%
80%
10 Ma
5 Ma
20%
40%
60%
80%
40%
15 Ma
60%
20% 20%
40%
60%
20 Ma
20% 20%
20%
Maximum basin depth of standard model
40%
80%
20%
Maximum basin depth of standard model
20% 20%
40%
60%
80%
Standard Model
Maximum basin width of standard model
20%
20%
Maximum basin depth of standard model
15 Ma
0 Ma
80%
Maximum basin width of standard model
20 Ma
5 Ma
20%
40%
60%
0 Ma
Figure 6.8: Forearc evolution with changes in sedimentation rate
80%
152
6.4.3
Chapter 6
Suite two - viscosity parameters
The evolution of topography in geologic systems is clearly influenced by the viscosity,
and hence the ductile behavior, of the crust. In accretionary margins this might lead
to the development of backward thrusting sequences, viscous flow and the formation
of forearc ridges [e.g. Pavlis and Bruhn, 1983; Smit et al., 2003; Fuller, 1996]. Here,
we want to expand upon these studies to explore the effects that variations of specific
viscous parameters (Aµ , nµ , and Q) have in the development of accretionary margins.
The relationship between stress and the rate of deformation is modeled by the
power-law viscosity equation:
−Q
nµ
,
(6.1)
Dij = Aµ σij exp
RT
where Aµ and nµ are constants dependent on the material, Q is the activation energy,
R is the molar gas constant, and T is the temperature. Table 6.3 shows the explored
range of values for the aforementioned parameters.
Table 6.3: Viscosity parameterization: Aµ , nµ , and Q. For every model set (left column),
the parameter being modified is marked with italic bold characters.
Model
Standard
A1-to-A7
n1-to-n7
Q1-to-Q7
Power-law viscosity
Aµ (Pa−n /M a)
nµ
2.05
0.05-4.15 (∆0.7)
2.05
2.05
2.65
2.65
2.35-2.95 (∆0.1)
2.65
Activation Energy
Q (J/mol)
1.7
1.7
1.7
1.4-2.0 (∆0.1)
The pre-exponential viscosity parameter, Aµ , affects the amount of relative
differential motion between the subsided and the uplifted areas, and the age of uplift.
Within the parameter range explored, lower values produce a deeper initial forearc
basin and slow down the formation of a subsiding high that develops at younger
ages than the standard. This occurs only weakly and is unable to compensate for
the primary subsidence signal (Fig. 6.9). Subsidence in the initial forearc basin decreases with time until formation of the structural high, which leads to accelerating
subsidence in the pro-basin and decelerating in the retro-basin.
Higher values of Aµ led to a more pronounced forearc high uplift, which develops
at earlier stages than the standard and more widespread stresses, i.e., broader shear
zones. The formation of the forearc high outpaces the primary subsidence signal, and
surface uplift in this area is eventually larger in absolute value than the subsidence
undergone by the initial forearc basin (Fig. 6.9). The rate of the motions after forearc
high development accelerates with time for both the forearc high and the pro-basin,
leading to differential motions that become more relevant with time; whereas the
rate of subsidence in the retro-basin remains roughly constant (Fig. 6.9).
Similar to Aµ , nµ exerts control on the relative differential motions and their age
and on the diffusion of deformation. Lower values of nµ within the parameter spectrum studied, led to the formation and protracted development of the initial forearc
basin, which thickens while decelerating its subsidence rate very slowly (Fig. 6.9).
153
CAP Model
The continuous development of the initial forearc basin makes the formation of the
forearc high difficult since it cannot outpace the overall subsidence.
25 Ma
10 Ma
Viscosity - A = 0.75 Pa-n/Ma
Viscosity - A = 3.45 Pa-n/Ma
Maximum basin width of standard model
Maximum basin width of standard model
20%
20%
80%
5 Ma
20%
40%
60%
80%
20% 20%
40%
60%
20% 20%
40%
60%
10 Ma
80%
60%
15 Ma
Maximum basin depth of standard model
40%
20 Ma
80%
20%
Maximum basin depth of standard model
40%
20% 20%
20%
60%
80%
15 Ma
0 Ma
Standard Model
Maximum basin width of standard model
Maximum basin depth of standard model
20 Ma
5 Ma
20%
40%
60%
80%
0 Ma
(a) Variations in viscosity parameter Aµ
Figure 6.9: Forearc evolution with changes in viscosity parameter Aµ .
Viscosity - n = 2.85
25 Ma
10 Ma
Maximum basin width of standard model
Maximum basin width of standard model
60%
80%
15 Ma
10 Ma
5 Ma
20%
40%
60%
80%
20% 20%
40%
20 Ma
60%
60%
40%
20% 20%
20%
Maximum basin depth of standard model
40%
80%
20%
Maximum basin depth of standard model
60%
40%
20% 20%
20%
80%
20%
80%
15 Ma
0 Ma
Standard Model
Viscosity - n = 2.45
Maximum basin width of standard model
Maximum basin depth of standard model
20 Ma
5 Ma
20%
40%
60%
80%
0 Ma
(b) Variations in viscosity parameter nµ
Figure 6.9: Forearc evolution with changes in viscosity parameter nµ .
154
Chapter 6
Higher values produce an early uplift that is older than that of the standard.
Soon after initiation of the model the forearc high forms, fragmenting the initial
forearc basin, which had barely started to form. Uplift rates accelerate with time
in the forearc high and its width remains roughly constant. The subsidence rates in
the pro-side basin accelerate first but slow down at more advance stages (younger
times); whereas the opposite trend in subsidence rates is recorded in the retro-basin,
with older deceleration and younger acceleration (Fig. 6.9).
Within the parameter spectrum checked (Table 6.3), the activation energy, Q,
controls the occurrence and shape of uplift. Its effects can broadly be considered
reverse to those seen for nµ . Relatively low values produce an early-staged, older,
and larger (than standard) uplift of a convex up area, blocking the development
of the large initial forearc basin, while inducing a second uplift in the retro-side.
Surface uplift in the forearc high outpaces the primary subsidence and uplift rates
accelerate with time. Subsidence rates also accelerate with time for both the proand retro-basins. The width of the pro-basin decreases with time while that of the
retro-basin increases until stabilization (Fig. 6.9).
High values accelerate the subsidence rates in the initial forearc basin, as the
basin becomes slightly narrower (Fig. 6.9). High values lessen the strain distribution
and thus increase the stability of the margin, which counteracts the upward motion
of a box-shaped area that takes place in later stages of the simulations.
Viscosity - Q = 1.6 J/mol
25 Ma
10 Ma
20 Ma
5 Ma
15 Ma
0 Ma
Standard Model
80%
10 Ma
5 Ma
20%
40%
60%
80%
20% 20%
40%
60%
20% 20%
40%
60%
15 Ma
80%
60%
20 Ma
Maximum basin depth of standard model
40%
80%
20%
Maximum basin depth of standard model
20% 20%
40%
60%
80%
20%
20%
20%
Maximum basin depth of standard model
Viscosity - Q = 1.8 J/mol
Maximum basin width of standard model
Maximum basin width of standard model
Maximum basin width of standard model
20%
40%
60%
80%
0 Ma
(c) Variations in viscosity parameter Q
Figure 6.9: Forearc evolution with changes in viscosity parameter Q.
CAP Model
6.5
155
Discussion
Accretionary margins have subduction wedges markedly compartmentalized, along
the convergence direction as a result of strain distribution (Fig. 6.1). The overriding plate in subduction wedges undergo brittle deformation in their seaward regions,
i.e., the trench, the trench-slope, and the trench-slope break areas, and remain mostly
undeformed in their landward forearc basin region. Forearc highs, uplifted flat terrains developed farther inland, are classically thought to form in margins where large
amounts of accreted material encounter a “backstop” barrier during their inland migration [e.g. Byrne et al., 1993]. Here we focus on the mechanics of uplift of one of
these areas, the southern margin of the Central Anatolian Plateau, and on the relative impact of the possible parameters influencing its development. Our simulations
demonstrate that the accretion and sedimentary thickening of the Cyprian forearc
basin system promote thermal weakening and viscous deformation in the base of the
crust that drive the subsequent surface uplift of the modern Central Taurus Mountains. Models clearly manifest that the viscous rheology of the overriding plate and
the amount of sediments in the forearc basin control the type and timing of uplift
in the forearc high without the need for a backstop.
6.5.1
A new view at uplift in South Turkey
Post-8 Ma surface uplift of the southern margin of the Central Anatolia Plateau in
south Turkey is attributed to slab break-off [e.g. Cipollari et al., 2012; Cosentino
et al., 2012; Schildgen et al., 2014a], although tomography shows the intact Cyprus
slab underlying the area with more pronounced uplift [Bakırcı et al., 2012], an area
where regional-scale contractional structures developed during the time-frame of interest (chapter 5). The analysis of the structures observed along the regional 2D
geologic transect of the central Cyprus arc demonstrate clear signs of contraction.
The type, distribution, relative age, and geometry of these contractional structures
are compatible with their development in relation to a wide accretionary subduction system [e.g. Karig and Sharman, 1975; Dickinson and Seely, 1979] (Fig. 6.1).
Thus, in this contribution, we attribute the surface uplift of the southern margin of
the Central Anatolia Plateau to shortening and the accretion of material from the
Cyprus subduction system, a mechanism that is validated by our models (Fig. 6.10).
Simulations can reproduce the overall geometry and the main contact relationships
as well as the main vertical tectonic events of the south Anatolian upper plate,
including the early continuity and later disruption of the Miocene basin(s) in the
presence of regional-scale accommodating structures in the seaward side, and in their
absence in the landward side of the system.
A conceptual representation of this surface motion mechanism in the context of
the tectonic evolution of the central Cyprus arc is shown in Fig. 6.10. Mechanical
accretion is responsible for the structural highs and lows in the seaward side of the
margin, where the development of young contractional structures, leading to ridges
and related piggy-back basins, is clearly linked with the subduction system [Aksu
et al., 2005a, b; Calon et al., 2005a, b; Hall et al., 2005a, b]. The trench in south
Cyprus, the Troodos Ophiolite, and the Messaoria Basin belong to this trench of
slope domain, a wedge-like region undergoing active accretion in which seaward imbricate thrusts systems develop. The Kyrenia Range is the trench-slope break, a
156
Chapter 6
100km
V.e.=2
Miocene basin formation
Middle Miocene
Horizontral length ~ 650km
Mechanical accretion
Miocene basin deepening
Upper Miocene
Differential vertical motions
Present-day
Cyprus Cilicia
Thermo-viscous accretion
Taurus
basin
CAVP
Central Anatolia
~45km
Lower Miocene
Domains fragmentation
S
Basement
Miocene
Accretion
~40km
~35km
~30km
Maximum crustal thickness
structural high shaped by sea- and landward verging thrusts, dividing the active
seaward areas from the landward “stable” forearc basin. The direct effects of accretion, and the vertical motions attached to it, are less readable in the forearc basin
area. Protracted accretion of material is nevertheless responsible for the thermal
expansion and ductile deformation at the base of the crust, which results in surface
uplift in the absence of ground-reaching faults (Fig. 6.10).
Viscous deformation
N
Figure 6.10: Conceptual evolution of the Central Cyprus Arc
This process led not only to the development of the uplift of the forearc high and
growth of the modern Taurus Mountains, but also to the subsidence occurring in
the intervening seaward area. The south Turkey surface uplift coupled with offshore
subsidence divided the forearc basin into a landward uplifted forearc basin, the Mut
Basin, and a seaward subsided forearc basin, the Cilicia Basin, while developing a
Miocene monocline (chapter 5) at ∼8 Ma and later times [e.g. Cosentino et al., 2012].
6.5.2
A new view at forearc highs
Although accretionary subduction margins are very diverse (plate convergence rate,
age of subducting oceanic crust, incoming sediment at the trench, dip of the subduction plate, etc.) even within the same system [e.g. Vannucchi et al., 2013], the
upper-plate morphology all over these margins is remarkably similar. Forearc highs
are often, but not always, seen in the internal parts of these margins.
A common point of view is that areas of relatively larger strength in the interior of
the forearc regions, backstops, induce a strength contrast that controls the structure
and limits the growth of the wedge. Under this view, forearc high formation is driven
CAP Model
157
by forced mechanical accretion against these larger-strength areas, which results in
upward terrain growth. The presence of a strength contrast in the inland areas seems
plausible given the variations in deformation history or metamorphic grade between
the pre-existing rocks and those formed during the development of the forearc basin
system. However, the apparent absence of backstop in some margins and the fact
that the uplift of the forearc high takes place within and on the stronger region are
indications that other mechanisms may also influence the development of forearc
highs and that of the subduction wedge [Fuller, 1996].
Our models prove that the forearc high develops above the area where the friction
exerted by the subducting plate does not affect the overriding plate. This singularity
might indeed be induced by stronger material in the retro-side of the model, but
also, and more simply, could be induced by the geometrical characteristics of the
wedge itself, i.e., the shape of the subducting slab, which controls the point where
the slab loses contact with the overriding plate. This was already postulated by
Fuller et al. [2006a] as “trench-slope breaks and forearc basins can form as a natural
consequence of the landward increase in slab dip in a margin with no lithologic
strength contrasts.” Besides, in our simulations, the uplift of the forearc high is
seen for common values of thermal conductivity, viscosity, and sedimentation rate,
and becomes more probable as the system ages. Variation in the values of these
parameters favors or complicates forearc high development without the need of a
material-strength backstop. Our models show that even in the less favorable lowviscous rheologies forearc highs eventually arise. Therefore, forearc highs should no
be considered an isolated feature conditioned to backstop presence, but rather an
integral part that develops in accretionary wedges at earlier stages of evolution in the
forearc basin systems if the conditions are favorable, and at advanced stages if they
are not. Also, we understand that forearc highs grow by protracted accretion and
do so in combination with feedback between the amount of sediments in the upperplate crust and the thermal blanketing of these forearc basin sediments. These
factors lead to a thermally activated viscosity drop at the base of the crust and
subsequent surface uplift.
6.5.3
A new view at vertical motions in forearc regions
Our models of accretionary margins show that if the amount of incoming sediments
at the trench remains constant, variations in the sedimentation rate control the age
and occurrence of surface uplift in the forearc high regions. Assuming an overfilled
forearc basin in which the sediments available from the surrounding source areas
outpace the accommodation space available, the amount of sediments that could
potentially enter the system depends solely on subsidence. Therefore, regional subsidence in the forearc basin could cause an uplift in the forearc high.
We observe that two counteracting effects compete during the growth of the forearc basin system: the thermal restrictions caused by the sediments in the forearc
basin system and their sedimentary load. Sediments have a lower thermal conductivity than basement rocks, and therefore lower heat transfer rates, which ultimately lead to changes in the thermal gradient underneath the forearc basin. As the
sedimentary infill thickens, the temperature gradient becomes more abrupt, with
increasingly larger temperatures at the base of the crust. Upward circulation is re-
158
Chapter 6
strained by the thermal resistivity or “blanketing” effect, caused by the forearc basin
sediments. Increasing temperatures at the base of the crust lead to lower viscosities,
and thus lower resistance to deformation, which in turn favors surface uplift. On
the contrary, larger sedimentary loads lead to larger isostatic subsidence. Since both
effects act simultaneously, the surface uplift induced by rheological changes at the
crustal base, and the subsequent sharper differences in the thermal temperature column, might be partially or completely masked due to regional subsidence. These
opposite controlling factors cause the region of the forearc basin to undergo a sequenced downward-upward motion in which surface uplift of the forearc high takes
place while subsidence in the surrounding areas persist, at least in a relative frame.
6.6
Conclusions
The models successfully reproduce upper plate deformation structures and modern
sedimentary architectures of the central Cyprus arc and prove that thermally activated viscous deformation at base of the Anatolian crust, yielded by the accretion of
incoming sediments of the Cyprus arc, is a plausible mechanism driving the surface
uplift that raised the modern Taurus Mountains. Frictional-mechanical deformation
occurring in the retro-side of the model controls the time sequence of structures
and topographies/bathymetries seen in the trench, trench-of-slope and seaward areas of the forearc system. Viscous-thermal deformation dominates in the landward
areas of the forearc system and controls the development of the forearc high and its
shape. Continued accretion thickens the forearc basin, which results in increasing
temperature at the base of the crust and subsequent weakened viscous flow of the
deforming material. The time delay between the onset of the accretionary system
and the moment of uplift change in relation to the velocity of accretion and the
sediment input, as well as the rheological and thermal state of the wedge. Higher
viscosity produces an older, more pronounced uplift. Decreasing values modify the
uplift from an older rounded-shaped uplift, to a box-shaped younger uplift, to the
absence of uplift. The margin stability increases when thicker basins develop in relation to either larger incoming sediment input at the trench or larger accumulation
rates. This thickening, which shifts deformation toward the forearc basin margins,
leads to younger uplift times and can force the avoidance of its occurrence, i.e., the
subduction system needs longer times to develop the forearc high. We conclude
that the topography and vertical tectonic motions seen in the Cyprian subduction
margin developed as a consequence of mechanical accretion in Cyprus and related
deep-seated deformation in south Turkey.