TECTO-125967; No of Pages 13
Tectonophysics xxx (2013) xxx–xxx
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Tectonophysics
journal homepage: www.elsevier.com/locate/tecto
Shear heating in extensional detachments: Implications for the thermal history of the
Devonian basins of W Norway
Alban Souche ⁎, Sergei Medvedev 1, Torgeir B. Andersen 1, Marcin Dabrowski 2
Physics of Geological Processes (PGP), University of Oslo, P.O. Box 1048, Blindern, 0316 Oslo, Norway
a r t i c l e
i n f o
Article history:
Received 9 July 2012
Received in revised form 27 May 2013
Accepted 4 July 2013
Available online xxxx
Keywords:
Shear heating
Supra-detachment basins
Numerical modelling
Post-Caledonian extension
a b s t r a c t
The supra-detachment basins in western Norway were formed during Lower to Middle Devonian in the
hangingwall of the extensional system of detachments known as the Nordfjord-Sogn Detachment Zone
(NSDZ). The basins experienced elevated peak temperatures (up to 345 °C adjacent to the main detachment
fault) during the terminal stages of the extension. In this study, we show that the heat generated by the deformation of the rocks within the shear zone of the detachment, also known as shear heating, played a role in
the thermal evolution of the entire region during the extension and could have induced the elevated peak
temperatures of the supra-detachment basins. We numerically analyse the influence of the rheology and
the deformation style within the shear zone, the rate of exhumation of the footwall, and the thickness of
the sedimentary accumulation on the top of the system. The model reproduces the elevated temperatures
recorded in the supra-detachment basins where locally 100 °C and 25% of the total heat budget can be attributed to shear heating. We show that this additional heat source may increase the peak temperatures of the
hangingwall as far as 5 km away from the detachment.
© 2013 Elsevier B.V. All rights reserved.
1. Introduction
Large-scale extensional detachments, resulting from gravitational
collapse of over-thickened crust, or extension driven by plate motions,
have been recognized and documented in a number of places and provide a major mechanism to exhume large segments of the lower crust
and mantle (Dewey, 1988; Lister et al., 1986; Manatschal, 2004;
Wernicke, 1985). The main geological feature of such crustal-scale excision is the juxtaposition of lower and upper crustal rocks across relatively narrow (100 m to a few kilometre thick) damage and shear
zones (e.g. Andersen and Jamtveit, 1990; Norton, 1986). Based on geological observations, the ‘simple shear model’ or modified versions of it,
explains the excision of the lithosphere along low angle normal faults
and shear zones. However, despite the general acceptance of this
model, several important questions remain regarding the dynamics
and the overall thermal evolution of crustal-scale detachments.
A detachment consists of a footwall (“lower plate”), a shear zone,
and a hangingwall (“upper plate”).
⁎ Corresponding author at: Institute for Energy Technology (IFE), Instituttveien 18,
NO-2007 Kjeller, Norway. Tel.: +47 22 85 60 48.
E-mail address: alban.souche@ife.no (A. Souche).
1
Present address: Centre for Earth Evolution and Dynamics (CEED), University of
Oslo, P.O. Box 1048, Blindern, 0316 Oslo, Norway.
2
Present address: Computational Geology Laboratory, Polish Geological Institute –
National Research Institute, Wrołcaw 53-122, Poland.
1) The footwall records higher-grade synkinematic metamorphism
and younger metamorphic cooling ages than the hangingwall. The
footwall is often referred to as a metamorphic core complex, and
in extreme cases the entire crust may be stretched off to expose
mantle rocks (Manatschal, 2004; Wernicke, 1985).
2) The shear zone is usually characterized by the ductile deformation with normal-sense kinematic indicators in the mylonites,
which are commonly capped by brittle fault rocks. In the case of
the Nordfjord-Sogn Detachment Zone, the mylonites may be several kilometers thick and associated with displacement in the order of
100 km (Andersen and Jamtveit, 1990; Hacker et al., 2003).
3) In the upper stratigraphic level of the hangingwall, syntectonic
supra-detachment basins develop along active faults (e.g. lowangle-normal faults, listric normal faults or strike-slip transfer faults)
and are commonly found in tectonic contact with the shear zone
(Andersen, 1998; Osmundsen et al., 1998, 2000; Seguret et al., 1989).
The thermal evolution of a crustal-scale detachment has to be
considered with two distinct aspects: fast exhumation and highlylocalized deformation. Pressure-temperature-time (PT-t) paths of the
footwall metamorphic core complexes typically indicate fast decompression at near-isothermal conditions during exhumation from the
lower to the upper crust level (Labrousse et al., 2004). The time window
bracketing such isothermal exhumation is only a few million years
(Hacker et al., 2010; Johnston et al., 2007) and is significantly smaller
than the time required for thermally equilibrating the crust by diffusion
0040-1951/$ – see front matter © 2013 Elsevier B.V. All rights reserved.
http://dx.doi.org/10.1016/j.tecto.2013.07.005
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
2
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
The other characteristic feature controlling the thermal evolution
of the detachment zone results from the relative movement between
the foot- and the hangingwall. Large strain can accumulate inside the
shear zone of the detachment, and this may have a significant effect
(several tenths of Myr, e.g., Turcotte and Schubert, 2002). Therefore, the
thermal structure of the crust after fast exhumation of a high- to
ultra-high-pressure [(U)HP] terrane might be strongly perturbed by interaction between a “hot” footwall and a “cold” hangingwall.
a
5E
6E
Bremangerlandet
Hornelen
Frøya
Gjegnalundsbreen
Svelgen
Kalvåg
Hornelen basin
Ålfobreen
Sandane
Hyen
61.6 N
Hovden
Florø
Håsteinen basin
Trondheim
Askrova
Førde
Kvamshesten basin
Norway
Bygstad
Dale
Bergen
Oslo
Fure
Fjaler
Værlandet
61.2 N
Sørbøvåg
Hyllestad
Solund basin
Leirvik
Devonian basins (hangingwall)
Lavik
Sula
Major detachment faults
Hardbakke
Ytre Sula
Caledonian nappes
(hangingwall / shear zone)
10 km
b
Western Gneiss Region (footwall)
Erosion
W
E
Devonian basins
Burial depth basins
>10 km
Exhumation of the footwall?
Shear heating?
Fig. 1. a) Geological map of the Nordfjord-Sogn Detachment Zone in western Norway. b) Schematic cross section (E-W) across the detachment. Heat flux introduced at the base of
the Devonian basins can come from the exhumation of the footwall or from shear heating in the shear zone.
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
3
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
2. Geological constraints of the model
2.1. Footwall conditions and strain partitioning in the detachment
The NSDZ is one of the largest extensional detachments in the world,
with a displacement of 70 to 100 km (Andersen and Jamtveit, 1990;
Hacker et al., 2003). The extension was initiated during the latest stages
of the Caledonian continent-continent collision when the crust reached
its maximum thickness and partially crystallized to (U)HP eclogites in
the Late Silurian to Early Devonian (Hacker et al., 2012; Krogh et al.,
2011; Kylander-Clark et al., 2009). The eclogites are found in the WGR
(Fig. 1), which constitutes the exhumed footwall of the NSDZ. Petrological and thermochronological studies of the footwall give consistent
PT-paths (Fig. 2) with a common near-isothermal decompression
from peak-pressure conditions within maximum 10 Ma (Hacker et al.,
2003, 2010; Johnston et al., 2007; Kylander-Clark et al., 2009; Root et
al., 2005; Young et al., 2011).
The HP rocks exposed structurally below the basins (here we use
data from near the Solund and Hornelen basins) provide useful constraints regarding the Caledonian metamorphism (Fig. 2) and the kinematic evolution of the NSDZ (Hacker et al., 2003; Johnston et al.,
2007; Young et al., 2007). We take the maximum burial of the HP
units to represent the initial depth (~ 60 km) where the exhumation
started along the shear zone. Based on the decompression path of
the HP units, we constrain our model to account for 40 km of vertical
exhumation bringing the HP rocks to a depth of 20 km, at which the
retrograde mineral assemblages and their cooling age have been determined by 40Ar/39Ar geochronology (Andersen, 1998; Chauvet
2.5
Pressure (lithostatic) [GPa]
on the heat budget of the system in form of shear heating (Brun and
Cobbold, 1980; Burg and Gerya, 2005; Campani et al., 2010; Hartz and
Podladchikov, 2008; Leloup et al., 1999; Schmalholz et al., 2009;
Scholz, 1980). Shear heating rate depends on the non-elastic strain
rate and the deviatoric stress experienced by rocks during deformation. The amount of shear heating produced in a geological system
is difficult to measure and modelling provides a viable tool for
assessing the role of shear heating.
We base our study on the geological observations of the
Nordfjord-Sogn Detachment Zone (NSDZ) of western Norway. In the
top stratigraphic level of the detachment, several supra-detachment
basins formed (Fig. 1) as response to repetitive tectonic motions during the exhumation of the footwall (Osmundsen et al., 1998; Seguret
et al., 1989; Seranne et al., 1989; Steel et al., 1977). These supradetachments basins are commonly named the Devonian basins due
to the presence of Devonian plant fossils in the sediments (Høegh,
1945; Kolderup, 1916, 1921, 1927). The special feature of the Devonian basins is the documented low-grade metamorphism (Seranne and
Seguret, 1987; Souche et al., 2012; Svensen et al., 2001; Torsvik et al.,
1988). Souche et al. (2012) showed that the temperature conditions
at the burial depth of 9.1 km was as high as 345 °C which can be translated to a geothermal gradient of 38 °C/km. Additionally, the peak
temperatures in the Devonian basins increase with proximity to the
detachment contact and thus suggest that the lateral variation of
maximum temperatures was controlled by a dynamically evolving
system, most likely linked to the detachment. Was shear heating contributing to the heat budget in the system? Could it have an impact on
the lateral variation of the peak temperatures within the supradetachment basins?
In this study, we analyse the thermal evolution of the detachment
zone and estimate the contribution of shear heating on the heat budget of the area. We first present a compilation of available geological
data from the NSDZ to build a well-founded geological model that can
be studied numerically. We compare the results of our numerical
studies with independent observations from the supra-detachment
basins of western Norway and discuss the importance of the shear
heating in this particular geological setting.
Hacker et al., 2003 (peak)
Hacker et al., 2003 (retro)
Johnston et al., 2007 (peak)
Johnston et al., 2007 (retro)
Average of error ellipses (1σ)
HP
(WGR)
2
HP
(Caledonian nappes)
~60 km
1.5
1
0.5
0
k
ea
~40 km
s
ion
dit
n
co
“Isothermal”
decompression
p
HP
(retrograde overprints)
~20 km
400
500
600
700
800
Temperature [oC]
Fig. 2. Peak and retrograde Caledonian metamorphism from the WGR and HP Caledonian nappes exposed beneath the Solund and Hornelen basins. Thermobarometry data
from Hacker et al. (2003) and Johnston et al. (2007).
and Dallmeyer, 1992; Cuthbert, 1991; Fossen and Dallmeyer, 1998;
Young et al., 2011).
The mylonites of the NSDZ have a maximum thickness of 5.5 km
north of the Hornelen basin, but the precise original thicknesses cannot
be accurately determined because of excision by later faults. Marques
et al. (2007) estimated the shear strain partitioning in the NSDZ and
showed that the lower parts of the mylonites experienced considerable
flattening across the foliation in addition to simple shear (ratio of simple to pure shear ~1). The upper ~2 km of the mylonites are, however,
characterized mainly by simple shear, with a shear strain of approximately 20 (Andersen and Jamtveit, 1990; Hacker et al., 2003; Marques
et al., 2007), used here as the bulk shear strain along the NSDZ.
2.2. The hangingwall peak temperatures
Following fluid-inclusion and mineralogical studies of the basins
(Svensen et al., 2001), a recent study using Raman spectroscopy of carbonaceous material (RSCM) from Devonian plant fossils (e.g. Souche et
al., 2012) documents high peak temperatures from the Solund,
Hornelen, and Kvamshesten basins. The temperatures locally reached
345 °C for an estimated burial of 9.1 km in the Kvamshesten basin
(Table 1). This suggests a geothermal gradient of approximately
38 °C/km, which is clearly at variance with commonly accepted geothermal estimates for a thickened orogenic crust or even for a crust in
early stages (415-395 Ma) of post-orogenic extension (Fossen, 2010;
Gabrielsen et al., 2005; Hacker, 2007; Spengler et al., 2009). Souche et
al. (2012) concluded that the elevated basin temperatures could only
be explained by an additional heat source introduced at the base of
Table 1
Burial depth and peak temperature estimates of the Devonian basins of W Norway.
Fluid-inclusion and
mineralogical studies (*)
Raman spectroscopy of
carbonaceous material (**)
Burial
depth(***)
T (local)
T (average)
Hornelen
9.1 ± 1.6 km
250 ± 20 °C 295 ±
Kvamshesten 9.1 ± 1.6 km
250 ± 20 °C 345 ±
Solund
13.4 ± 0.6 km 315 ± 15 °C 284 ±
295 ±
301 ±
30
30
30
30
30
Distance from
the detachment
°C at 1 km
°C at 1 km
°C at 8 km
°C
°C
*Svensen et al. (2001); **Souche et al. (2012); (***) The burial depth is estimated from
lithostatic pressure assuming a constant rock density of 2600 kg/m3 (Svensen et al.,
2001).
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
4
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
the hangingwall during the extension. This conclusion also is in agreement with previous description of ductile deformation of the basin-fill
sediments adjacent to the detachment, which could suggest elevated
temperatures in this area (Braathen et al., 2004; Norton, 1987).
2.3. Model: constraints and simplifications
The different geological aspects presented in the previous sections
allow us to build a simplified geological model of the evolution of the
detachment zone used in the numerical analysis (Fig. 3). The initial
configuration of the model (Fig. 3a) corresponds to the thickened
Caledonian crust with a Moho depth set to 70 km corresponding to
approximately twice the average thickness of continental crust. We
assume that the deformation was initiated by the development of a
lithospheric-scale shear zone with a constant dip angle α = 30°
(Andersen and Jamtveit, 1990; Norton, 1986).
The active deformation within the shear zone and the exhumation of
the footwall is considered within a time window denoted t*. After that
time the system still. We assume constant rates of deformation for simplicity. During deformation, the footwall translates upward along the detachment and accounts for the exhumation of the lower crustal body
(hf = 40 km) from 60 to 20 km depth (marked by a star in Fig. 3 a and
b). At the same time, the hangingwall translates downward along the detachment and creates accommodation space in the basin formed at the
surface. We assume immediate erosion and sedimentation processes so
that the model does not produce variations in topography. At time t*,
the lower crustal body is at 20 km depth and the basin is at maximum
depth (bmax). Accordingly to the estimated burial depth of the Devonian
basins (Table 1), we consider two values for bmax, 10 km (for the
Hornelen and Kvamshesten basins) and 15 km (for the Solund basin).
An important element of the model is the style of deformation
within the shear zone. Despite observations of Marques et al.
(2007) on the significant component of pure shear within the lower
parts of the NSDZ (Section 2.1), we assume simple shear only. This
simplification reduces the number of model parameters, which are
poorly constrained and might be related to the earlier stages of exhumation (Andersen et al., 1994). Thus, simple shear (Fig. 3c) accommodates the total relative displacement between the foot- and
hangingwall. According to the observations and measurements
along NSDZ (see Section 2.1), we set a bulk shear strain (γbulk) of 20
at time t*. Given γbulk, α, and the total displacement between the
foot- and hangingwall, we estimate the thickness of the shear zone
to be D = 5-5.5 km depending on bmax, which is consistent with the
observations (Marques et al., 2007).
The assumption of uniform translation (and thus, without any deformation) of the foot- and hangingwall is a simplification of our
model. We have, however, tested the importance of the lateral
stretching of the units, and found that a moderate extension does
not cause significant changes of the peak temperatures in vicinity of
the detachment zone.
a) Initial configuration (t=0)
Normal sense reactivation of the detachment fault along the
Solund and Kvamshesten segments (Fig. 1) in the Permian, and
again in the Late Jurassic has been documented by palaeomagnetic
and geochronological studies of fault breccias along the NSDZ
(Andersen et al., 1999; Eide et al., 1997; Torsvik et al., 1992), and
may have contributed to some exhumation. In this study, we assume
that the Devonian basins reached their peak temperature conditions
during or shortly after the major post-Caledonian extension and
were not influenced by later reactivation of the detachment faults.
3. Model
3.1. Governing equations
We analyze the thermal evolution in the model by solving the
transient heat balance equation:
ρC p
!
"
∂T !
þ v " ∇T −∇ " ðk∇T Þ ¼ Hr þ H s
∂t
ð1Þ
where T is the temperature, t is the time, k is the thermal conductiv!
ity, v is the rock velocity, ρ is the rock density, Cp is the rock specific
heat capacity, and Hr is the radiogenic heat production (see characteristic values in Table 2). The shear heating term Hs is the product of the
deviatoric stress τ and the non-elastic strain rateγ̇ (Brun and Cobbold,
1980; Burg and Gerya, 2005; Turcotte and Schubert, 2002) so that:
H s ¼ τ γ̇
ð2Þ
The hangingwall and footwall translate along the detachment without internal deformation. All the deformation is accommodated within the shear zone and thus shear heating (Eq. (2)) is only used in the
shear zone.
Completing the mathematical model of thermal evolution requires
the specification of initial and boundary temperature conditions, velocities and rheological relationship. The assumption of constant
translation during exhumation (Section 2.3) gives the velocities within the foot- and the hangingwall by dividing the total displacements
by t*. The initial and boundary temperature of the system and specification of Eq. (2) for the shear zone are discussed below.
3.2. Initial geotherm
The initial thermal situation before the extension has to be considered in the geological framework of the Caledonian orogeny (peak
condition data points on Fig. 2). Thus, to obtain the initial thermal
state, we perform an auxiliary modelling of the evolution of the lithosphere during the Caledonian collision. In this experiment we assume that, prior to collision, the lithosphere was in steady state and,
during the orogeny, the lithosphere thickened uniformly by pure
b) Final configuration (t=t*)
c) Model geometry
(not to scale)
depth [km]
0
20
0
α
bmax
upper crust
basin
bmax
20
40
40
lower crust
70
hf
D
α
hf
=40 km
60
60
80
0
γ bulk=2
Moho
32
α=30°
α
Moho
mantle lithosphere
20 km
80
20 km
D = (bmax+hf)/sin(α)/γbulk
Fig. 3. a) Initial and b) final configuration of the different rock units in the model. The maximum basin depth bmax = 10-15 km. The total vertical motion of the footwall hf = 40 km.
c) Sketch of the translation motion between the foot- and the hangingwall units along the constant dipping shear zone (α = 30°). All the deformation is accommodated inside the
shear zone with a resulting bulk shear strain γbulk = 20 at time t*. The thickness of the shear zone D is estimated to be 5-5.5 km.
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
5
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
Table 2
Thermal and rock properties of the different units of the model.
Parameters
symbol
Thermal conductivity*
Specific heat capacity
Rock density
Radioactive heat production**
Unit
Value
W/(m.K)
J/(kg.K)
Kg/m3
W/m3
k
Cp
ρ
Hr
sediments
upper crust
lower crust
mantle
a = 0.64; b = 807
1000
2600
1.2 × 10−6
1000
2700
1 × 10−6
a = 1.18; b = 474
1000
2900
4 × 10−7
a = 0.73; b = 1293
1000
3300
2 × 10−8
*k=a+b/(T+350), where T is the temperature expressed in °C (after Clauser and Huenges (1995)). ** after Hasterok and Chapman (2011).
shear. We used a 1D computer code LiToastPhere (Hartz and
Podladchikov, 2008) to perform the calculations.
We consider the Caledonian continent-continent collision to last
for 25 My (Corfu et al., 2006; Johnston et al., 2007; Torsvik and
Cocks, 2005) and to result in 100% thickening of the crust (double
of the original thickness) which corresponds to an average strain
rate of 8.7 × 10-16 s-1. The model lithosphere consists of three layers,
composed of a wet quartzite upper crust, a diabase lower crust, and a
dry dunite mantle (Ranalli, 1995). The initial thickness of the upper
crust is set to 16 km (Hasterok and Chapman, 2011). The lower
crust extends to the Moho discontinuity with initial depths of 30, 35
and 40 km (corresponding to the model “Mxx”). The initial depth of
the mantle lithosphere is set to 140 km (Artemieva, 2006). The thermal properties used in the calculations are listed in Table 2.
We performed a series of numerical tests varying the initial Moho
depth and the amount of radioactive heat production and considering
the influence of shear heating. We compare the model results with
the data of the peak metamorphic conditions of the Caledonian
nappes below the Solund basin and the Hornelen basins (Fig. 2).
The best fit geotherm is obtained with the model M35-sh (Fig. 4),
which accounts for shear heating during the deformation of the lithosphere. Fig. 4 also presents model results obtained with similar
3
M4
0
5
h
-s
40
0-s
35
M
No shear heating
h
-s
h
M
2 ~ 72 km
Shear heating
M3
Pressure (lithostatic) [GPa]
2.5 ~ 88 km
M3
M3
0
100% thickening (pure shear) in 25Ma
Strain rate = 8.7e-16s-1
1.5 ~ 54 km
1
~ 37 km
Geological data (peak conditions):
Hacker et al. (2003)
Johnston et al. (2007)
Average of error ellipses (1σ)
0.5
400
500
600
700
800
Temperature [oC]
Fig. 4. 1D model predictions (solid and dash lines) versus pressure-temperature (PT)
estimates of the HP units in the footwall of the Nordfjord-Sogn Detachment Zone
(Hacker et al., 2003; Johnston et al., 2007). Different model setups are compared for
different Moho-depth indicated by “Mxx”, and with or without shear heating indicated
by “-sh”. The grey shades represent the spread of the curves “M35” and “M35-sh” by
varying the radioactive heat production by ±25%. We observe a good fit between
the representative PT data and those predicted by the model when the bulk deformation of the lithosphere is taken into account. Without shear heating during collision,
the entire system is too cold in comparison to the geological data.
setups but where shear heating was ignored (M30, M35 and M40).
The shear heating models are warmer by an average of 100 °C, illustrating the importance of this heat source during deformation. The results also illustrate greater dependence on the initial conditions and
rock properties for the models without shear heating.
We use the preferred temperature profile M35-sh (Fig. 4) to build
the initial temperature distribution. We also use this profile to constrain the bottom boundary condition during the model evolution.
As the footwall moves upward during exhumation, the material incoming into the system assumed to sustain its temperature initially
prescribed by M35-sh. This assumption ignores the thermal relaxation of the deep lithosphere after Caledonian orogeny, which is approximately valid for fast exhumation considered here.
3.3. Kinematic of the shear zone
Restoring the exact distribution of velocities and stress field within
the shear zone is beyond the scope of this study. Instead, we approximate the velocities within the shear zone using two end-member approaches, in which either shear strain rate (γ̇-cst) or shear stress (τ-cst
model) is kept constant across the shear zone. The two kinematic
models correspond respectively to an upper and lower bound on the estimate of shear heating expressed by Eq. (2). Whereas the γ̇-cst model
forces uniform deformation and thus active heat production even in
strong areas of the shear zone, the τ-cst model results in localized deformation of weaker parts of the detachment, and thus produces less heat.
It is important to keep in mind that neither the γ̇-cst nor the τ-cst
models reproduce exactly the actual deformation process within the
Nordfjord-Sogn Detachment and the main purpose in using these two
virtual kinematic models is to deliver an upper and lower bound on
the shear heating effect.
3.3.1. Rheology
The shear zone accumulates the entire deformation in the model
(Fig. 3c) and hosts the production of shear heating. This deformation
occurs in the range of 0-60 km, and thus involves the transition from
brittle to ductile behaviour of rocks.
The brittle behaviour is described by Byerlee’s relationship:
#
$ y
τBy ¼ 1−p f μg ∫ ρdy
ð3Þ
0
See Table 2 for definitions and values of parameters used. This
constitutive relationship assumes no deformation for the stress
values below critical, τBy, which increases with depth (Fig. 5). We
approximate Eq. (3) using a power law relationship (after Nanjo et
al., 2005; Turcotte and Glasscoe, 2004)
τ ¼ τ By
qffiffiffiffiffiffiffiffiffiffi
nb
γ̇b =γ̇c
or γ̇b ¼
τ
τ By
!n
b
γ˙c
ð4Þ
where γ̇c is a critical strain rate and γ̇b is the strain rate of the brittle
deformation. Here we define γ̇c as the bulk strain rate imposed by
the motion of the foot- and the hangingwall (γbulk/t *). For large
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
6
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
a
b
τ shear stress [MPa]
0
τ
0
100
200
300
400
γ = 2.1x10-13
After exhumation
(Wet Quartzite)
20 15 0
1 5
=1
nb
20
Brittle/Ductile
Transition zone
zit
uart
et Q
l (W
it t ia
n
e)
I
τBy
Depth [km]
10
it
artz
Qu
Dr y
al (
e)
e
ittl
Br
30
it i
Int
re
ilu
fa
0
γc
γ
40
Fig. 5. a) Brittle failure approximated by power law relationship between shear stress and shear strain rate. The power law expression (Eq. (3)) approaches Byerlee’s brittle failure
(where τBy(y) is function of depth) for large power law exponent (nb = 20 in this study). b) Depth distribution of the shear stress within the shear zone (γ̇-cst model and t* = 3
My). A single stress profile is drawn for dry and wet quartzite at the initial conditions. The brittle-ductile transition is smoothed (illustrated by yellow) because of the use of approximation presented on Fig. 5a and Eq. (4). Lateral variation of the temperatures at the end of the exhumation (t*) scatters the shear stresses for a given depth as illustrated
by the gray shade for the wet quartzite simulation.
power law exponent (nb = 20), the deviatoric stress is approximately equal to the critical stress τBy (yield strength) for a wide range of
strain rate values centered about the critical strain rate γ̇c .
The ductile behaviour is described by a power-law for dislocation
creep:
τ ¼ expðE=RTnd Þ
qffiffiffiffiffiffiffiffiffi
n
γ̇d =A or γ̇d ¼ τ d A expð−E=RT Þ
nd
ð5Þ
where γ̇d is the strain rate of the ductile deformation. The Arrhenius
term exp(E/RTnd) results in a strong decrease of the stress with temperature and, for standard geotherm, with depth (Fig. 5b). The shear zone
exposed today between the Devonian basins and the WGR is mainly
formed of Caledonian nappes represented by meta-sedimentary units
and some deformed slices of basement. The rheological weakness of
the meta-sediments is approximated by flow laws for wet and dry
quartzite (see Table 3).
We assume the brittle and the ductile mechanism of deformation
to act simultaneously on the total shear strain rate, and it allows us to
write:
γ̇ ¼ γ̇b þ γ̇d
ð6Þ
For a given shear stress value across the shear zone, Eqs. (4) – (6)
yield a strain rate profile. For a given strain rate value across the
shear zone, a shear stress profile is obtained by solving Eq. (6) after
Table 3
Material and rheological (Ranalli, 1995) parameters used in the numerical experiments.
Parameters
Symbol
Unit
Value
Gravity
Friction coefficient
Pore fluid factor
Power law exponent
Gas constant
g
μ
Pf
nb
R
m.s−2
J.K−1
Power law exponent
Pre-exponential multiplier
Activation energy for
dislocation creep
nd
A
E
−
MPa-nd.s-1
kJ.mol−1
9.81
0.7
0.4
20
8.314
Quartzite: Wet/Dry
2.3/2.4
3.2 × 10−6/6.7 × 10−6
154E + 3/156 × 103
substituting Eqs. (4) and (5) into it. Our approximation to the shear
stress (Fig. 5b) follows the minimum of the two stress regimes and
smoothes the transition zone.
3.3.2. Constant shear strain model (γ̇-cst model)
The constant shear strain rate model results in a linear velocity
profile across the shear zone (Fig. 6a) defined by the hangingwall
and footwall velocities. The deformation is homogeneous and distributed across the shear zone, without localization, and with the same
intensity in the “strong” or “weak” (thermally weakened) parts of
the shear zone. Therefore, we expect the shear heating term (Eq.
(2)) to be maximized using this kinematic model.
γ̇¼
γbulk
t⁎
ð7Þ
The corresponding shear stress, variable across and along the shear
zone, is determined at any points by substituting Eqs. (4) and (5)
into Eq. (6) and solving it using Eq. (7). This approach insures a
mass conservation in the model but induces large variations of
shear stress across the shear zone due to the variations of rheological
properties.
3.3.3. Constant shear stress model (τ-cst model)
The motivation for the τ-cst model is to augment the flow pattern
by avoiding large stress variation across the shear zone and allowing
for localization of the deformation. The localization has a direct impact on the distribution of heat produced by shear heating since
large deformations mostly occur in “weak” parts of the shear zone.
Thus, the τ-cst model is used as end-member kinematic model minimizing the impact of shear heating as opposed to the γ̇-cst model. A
value of the shear stress τ, uniform across and variable along the
shear zone, is obtained by integrating Eq. (6) across the shear zone
(in the y’ direction, Fig. 6) and using the bulk strain rate (γbulk/t*) as
a constraint:
D
D
&
'
˙ 0 ¼ ∫ γ̇b þ γ̇d dy0
∫γdy
0
0
"
#
!n
D
b
γ bulk
τ
n
0
D¼∫
γ̇c þ τ d Aexpð−E=RT Þ dy
τBy
t⁎
0
ð8Þ
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
7
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
0
20 km
Brittle
Ductile
Depth [km]
20
y’
x’
Line of zero velocity
40
Brittle-ductile transition
Velocities across the
shear zone
60
a)
b)
Couette
+
flow
γ-cst
d)
c)
Poiseuille
flow
=
τ-cst
Fig. 6. Typical velocity profiles of the two kinematic models. The profile (a) γ̇-cst is derived for the constant shear strain model and gives a uniform linear velocity profile across the
shear zone, constant in time. The profile (d) τ-cst is derived from the constant shear stress model (example of wet quartzite rheology at initial temperature distribution) and is the
sum of Couette (b) and Poiseuille (c) flows in the shear zone. The profile (d) is calculated at each time step with the updated temperature and may change as the model evolves.
the boundaries of the shear zone. However, significant rheological
variations along the shear zone result in velocity fields characterized
by a non-vanishing divergence and lead to a violation of mass and
momentum conservation condition. To limit these effects, we augment the kinematic profile by adding a Poiseuille flow contribution
(Fig. 6c) determined by keeping the integrated mass balance constant along the shear zone. This results in a non-zero pressure gradient in the shear zone. We note that due to adding the Poiseuille flow
component the original assumption of the constant shear stress
across the shear zone is not strictly satisfied. Yet, we choose to
refer to this model as τ-cst throughout the manuscript, as the strain
The strain rate profile is then calculated from Eqs. (4) and (5) and it
results in localized deformation in weak parts. Depending on the
dominant deformation mechanism, rock weakens at the upper part
of the shear zone in the brittle regime (where τBy is smaller,
Eq. (3)) and at lower part of the shear zone in the ductile regime
(where temperature is higher, Eq. (5)). Therefore, highly localized deformations occur along the boundary with the hangingwall in the
upper part of the shear zone, and along the boundary with the footwall in the lower part of the shear zone (Fig. 6b).
The velocity profile resulting from the constant shear stress assumption (Couette flow, Fig. 6b) respects the no-slip conditions at
a) Initial conditions
b) Exhumation of the footwall (t*=3 My)
0
0
depth [km]
20
20
60
70
HP unit
Moho
20 km
80
40
60
60
20 km
d
HP unit
Moho
20 km
80
e) Peak temperatures
ne
maxima
zo
350
0
ar
15
ay
aw
aw
ay
km
5k
m
250
co
nta
ct
s
he
20
150
782
Basin
600
HP unit
Moho
40
400
60
200
80
50
[oC]
10
Temperature in the basin [oC]
20
40
80
Basin
15
HP unit
Moho
32
40
0
Basin
15
c) Thermal relaxation (30 My)
Shearing
1
20 km
10
Thermal relaxation
3
5
7
9
Time [My]
Fig. 7. Temperature distribution in the model (a) initial; (b) at the end of exhumation (3 My); (c) after thermal relaxation (30 My). (d) Temperature evolution in the markers located at the bottom of the basin taken at the contact, 5 km, and 10 km away from the shear zone (white circles in a-c). The arrows mark the temperature maxima and illustrate the
scattered time for the peak conditions. e) Maximum temperature field recorded during the entire simulation. The model used here is set with t* = 3 My, a dry quartzite rheology
with γ̇-cst kinematic model, and bmax = 15 km.
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
8
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
Dry Quartzite
0
Depth [km]
AA
BB
A
15
20
[J.m-3]
1010
Wet Quartzite
B
109
t
τ-cs
t
s
γ -c
Range for upper crust
radioactive heat production
t
t
s
γ -c
τ-cs
108
20 km
60
107
Fig. 8. Distribution of total heat density produced by shear within the detachment zone for t* = 3 My (note the logarithmic colour-scale) and bmax = 15 km. Results are presented
for τ-cst and γ̇-cst models with dry and wet quartzite rheology. Shear heating produces 1-2 orders of magnitude more heat than the radioactivity of the upper crust during the same
time interval. A, AA, B, and BB indicate the position of cross sections presented in Fig. 10.
rates related to the Poiseuille flow are significantly smaller than the
total strain rates (Fig. 6c, d). The resulting velocity profile of the
τ-cst model is the sum of Couette and Poiseuille flows as illustrated
Fig. 6d.
3.4. FEM strategy
The model domain is 80 km deep and 120 km wide (Fig. 3). We
use a triangular mesh generated with Triangle (Shewchuk, 1996,
Dry Quartzite
depth [km]
3My
τ-cst
20
20
+143 oC
40
+110 oC
40
60
60
0
depth [km]
0
γ -cst
20 km
80
10My
Duration of the detachment
0
0
γ -cst
20 km
80
τ-cst
20
20
o
+81 C
40
[oC]
o
+67 C
40
60
60
20 km
80
100
80
20 km
80
60
Wet Quartzite
depth [km]
3My
+92 oC
40
20 km
20
0
20 km
80
0
γ -cst
20
τ-cst
20
o
+49 C
60
80
+67 oC
40
60
60
40
40
τ-cst
20
20
0
depth [km]
0
γ -cst
80
10My
Duration of the detachment
0
40
+39 oC
60
20 km
80
20 km
Fig. 9. Temperature anomaly induced by shear heating for different kinematic models across the shear zone, for dry and wet quartzite rheology, for t* = 3 and 10 My, and for
bmax = 15 km). The same colour code is used in all the plots. The numbers indicate the maximum difference in peak temperatures induced by shear heating.
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
2002) to discretise the model domain and accurately represent the
boundaries separating materials of different thermal properties. The
mesh is updated at each time step accordingly to the subsidence of
the basin on the top of the model.
Eq. (1) is implicitly solved with a finite element code using the
Streamline Petrov-Galerkin scheme. A similar discretisation has been
documented and used for geodynamical problems and we refer to the
paper of Braun (2003) for more technical details. Temperature and
velocity fields are linearly interpolated using 3-nodes elements.
The temperature-dependent thermal conductivity is calculated using
temperatures from the previous time step. We consider zero flux on
the lateral boundaries and Dirichlet boundary conditions on the top
and bottom (Section 3.2).
Shear zone related calculations are performed on an auxiliary regular grid onto which the modelled temperature is projected. The regularity of this grid allows accurate integration of Eqs. (4) – (6) and/or
(8) across the shear zone to define the stress field, strain and strain
rate. The resulting velocities and shear heating are then interpolated
back to the main grid each time step.
4. Results
We performed a systematic study by varying the time of the exhumation t* (3, 5, 10, and 15 My), the shear zone rheology (wet and dry
quartzite), the kinematic model of the shear zone γ̇ -cst and τ-cst),
and the maximum basin depth bmax (10 and 15 km). All the simulations
are repeated twice, with and without accounting for shear heating.
4.1. Temperature evolution and peak conditions
For any t*, the exhumation of the footwall and the subsidence of the
hangingwall leads to an unsteady temperature field. Thermal relaxation
is reached by 20 to 30 My after t*. The different thermal stages of the
model are illustrated in Fig. 7a-c. Note that for the similar initial and
final geometry and for the same distribution of the thermal properties
of rocks, the initial and final geotherms (Fig. 7a and c) will be equivalent. The peak temperature profile (Fig. 7e), however, depends greatly
on the parameters tested in our numerical experiments.
To facilitate the comparison of our model results with geological
observations (Section 2), we extract the peak temperature of each
point during the evolution of the model up to thermal relaxation
(Fig. 7e). The peak temperatures are reached at different time for different localities, often well after t = t* (Fig. 7d).
9
4.2. Heat produced by shear heating
The distribution of total heat produced by shear heating within
the detachment shear zone is presented in Fig. 8 (example presented
for t* = 3 My). The results for γ̇ -cst model show a maximum heat
produced in the upper 10-20 km, which corresponds to the largest
stress field at the brittle-ductile transition. The results for τ-cst
model are characterized by two areas with high heat production
(corresponding to areas of localized deformations), one in the upper
brittle domain, and one in the lower ductile domain. Irrespective of
the different spatial distribution of the heat production in the two kinematic models, the amplitude remains within the same range.
Shear heating may be locally two orders of magnitude higher than
the heat produced by radioactive decay in the upper crust. This effect
can be seen as a punctual heat source during the active deformation
in the detachment zone.
4.3. Temperature anomaly produced by shear heating
We illustrate (Fig. 9) the temperature anomaly produced by shear
heating by plotting the difference of modelled peak temperatures
obtained with and without shear heating. For dry quartzite models
and t* = 3 Ma, shear heating can introduce more than 100 °C at the
base of the basin. The impact depends on the rheology and on t*,
but remains significant (within 30-80 °C) for slower exhumation
and weaker rheology in the shear zone.
In the model configurations presented here, a value of the pore
fluid factor of pf = 0.4 in Eq. (3) is assumed along the shear zone.
The effect of pore-fluid pressure on the effective stress is still a matter
of debate (e.g. Nur and Byerlee, 1971; Sibson, 1994). However, the
impact of varying the pore fluid pressure within acceptable range is
not critical with respect to thermal effects related to shear heating.
The two kinematic models, γ̇ -cst and τ-cst, were introduced in
Section 3.3 as an upper and lower bound to estimate shear heating.
Fig. 9 shows that for the same parameters, γ̇ -cst model is hotter
than τ-cst model. However, despite the large difference in the
model velocity profiles (Fig. 6), the estimated temperature anomaly
produced by shear heating is on the same order. A difference of approximately 30 °C separates at most the temperature anomaly
obtained from both models, which shows that the style of deformation within the shear zone does not influence much the amplitude
of the induced temperature anomaly in the system.
5. Comparison with geological observations
160
A
AA
(Dry Quartzite rheology)
140
B
BB
(Wet Quartzite rheology)
Shear Strain
120
on
ati
rm )
efo ime
d
g
ed re
aliz tile
loc (Duc
on
ati
rm
efo ime)
d
ed reg
aliz tle
loc (Brit
100
80
60
40
Average shear strain = 20
(Model input)
20
0
0
1
2
3
4
5
Shear zone thickness [km]
Fig. 10. Strain profiles across the shear zone for the τ-cst models. Two shear bands representative for brittle and ductile deformation are observed along upper (left) and bottom (right) parts of the shear zone. The average strain across the shear zone is 20. The
rheology and the position at which the cross sections are taken have a minor impact on
the style of final strain profile. See locations of cross sections on Fig. 8, t* = 3 My.
5.1. Shear strain partitioning
Fig. 10 shows the strains along the cross sections A, AA, B, and BB
(Fig. 8) at the end of the simulations. Independent of the rheology and
the cross section position, the strain across the shear zone, described
by the τ-cst model, localizes mostly along the upper part, beneath the
hangingwall with maximum strain higher than 200 and along the
basal part of the shear zone next to the footwall with maximum strain
around 35-60. Strains higher than 100 are indeed speculated in the
upper part of the NSDZ (Hacker et al., 2003). However, there is no convincing field evidence for high-strain environment at the base of the
shear zone such as produced by the τ-cst model. This ductile shear localization can be seen as an artefact of our simplified shear zone geometry
(e.g., the strict rigidity of the footwall). The τ-cst model also produces a
strain gap in the middle of the shear zone, reducing the effective thickness of the shear zone. This is in contrast with the deformation pattern
observed along the NSDZ, which exhibits continuous and large strain
(20) over several kilometres (Andersen and Jamtveit, 1990).
In contrast to τ-cst, the γ̇ -cst model satisfies the observations on
continuous strain within the NSDZ, but lacks the high shear localization in the upper part of the detachment. Although the two models
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
10
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
Kvamshesten/Hornelen basins
Burial depth = 9.1 km
a
Max temperature in
the basin [oC]
340
Solund basin
Burial depth = 13.4 km
d
GEOLOGICAL DATA
Svensen et al. 2001
Kvamshesten
300
300
220
260
Dry Quartzite
8
6
2
4
b
Max temperature in
the basin [oC]
220
Solund
Dry Quartzite
8
6
4
6
4
2
e
Kvamshesten
340
300
380
340
Hornelen
260
300
220
260
180
Wet Quartzite
8
6
4
220
2
c
Solund
Wet Quartzite
8
2
f
340
Max temperature in
the basin [oC]
γ -cst τ-cst
t* = 3 My
t* = 5 My
t* = 10 My
340
Hornelen
260
180
MODEL RESULTS
380
Souche et al. 2012
Kvamshesten
300
380
340
Hornelen
260
300
220
260
180
No shear heating
220
No shear heating
8
6
4
distance from the detachment
[km]
2
contact
2
contact
8
6
4
distance from the detachment
[km]
Solund
Fig. 11. Comparison of model temperatures and geological data for the Kvamshesten and the Hornelen basins (a-c, max. burial depth 9.1 km) and for the Solund basin (d-f, max.
burial depth 13.5 km). Data from Svensen et al. (2001) present average peak temperatures for the entire basins, Souche et al. (2012) presents local peak temperatures. γ̇-cst and
τ-cst corresponds to the two kinematic models used for the shear zone. t* is the time window for the development of the detachment zone. (c) and (f) present the model results
obtained without shear heating; (b) and (e) present the model results obtained with wet quartzite rheology for the shear zone, and (a) and (d) with dry quartzite. The maximum
depth of the basin (bmax) was set to 10 km in (a-c) and 15 km in (d-f).
differ in the style of deformation and in correlation to observations,
the resulting heat produced within the shear zone based on those
models does not differ significantly. Thus, we can conclude that
exact reproduction of the flow within the shear zone has limited importance while considering the thermal budget of the NSDZ.
5.2. Implications for the supra-detachment basins thermal history
In Fig. 11, we compare the modelled peak temperatures with
the geological data (Section 2.2). Svensen et al. (2001) documented average values for peak temperatures of 260 ± 20 °C for the entire
Kvamshesten and Hornelen basins (horizontal lines on Fig. 11a-c) and
315 ± 15 °C for the Solund basin (horizontal lines on Fig. 11d-f).
Souche et al. (2012) reported the local estimates of peak temperature
of the sediments as a function of the distance to the detachment (data
points on Fig. 11).
The maximum reported burial depth of the Devonian sediments is
9.1 km for the Kvamshesten and the Hornelen basins and 13.4 km
for the Solund basin (Svensen et al., 2001). Thus, we compare the
geological data with the modelled peak temperatures taken at 9.1
(Fig. 11a-c) and 13.4 km (Fig. 11d-f). Fig. 11 presents the modelled
temperature curves for three different t* and two kinematics within
the shear zone (γ̇-cst and τ-cst) in each sub-figure.
All the results obtained without shear heating (Fig. 11c and f)
show low peak temperatures and fail to reproduce the geological
data. In addition, these modelled temperatures show a relatively
small lateral variation through the basins. The lateral temperature
variation does not exceed 40 °C in a range of 10 km away from the
detachment contact. That low thermal gradient illustrates the role
played by the hot rising footwall on the heating of the basin. This
heat source is not sufficient to induce significant variations of the
peak temperatures in the supra-detachment basins.
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
11
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
2.5
Model results:
Geological data:
Markers before exhumation
No shear heating
γ-cst model
τ-cst model
2
Johnston et al., 2007 (peak)
Johnston et al., 2007 (retro)
Corresponding samples
Average of error ellipses (1σ)
HP
(Caledonian nappes)
Pressure [GPa]
~60 km
“Isothermal”
decompression
1.5
~50 km
~40 km
HP
(retrograde overprints)
1
~30 km
0.5
~20 km
~10 km
0
200
250
300
350
400
450
500
550
600
650
700
750
Temperature [oC]
Fig. 12. Comparison of retrograde overprints (Johnston et al., 2007) with PT-t paths from numerical models of our study. All the model results are obtained for t* = 3 My, a dry
quartzite rheology, and bmax = 15 km.
The models including shear heating (Fig. 11a, b, d and e) show significantly higher peak temperatures and a stronger dependence on
the proximity to the shear zone. Wet quartzite models, however, do
not show good qualitative agreement with observations (Fig. 11b
and e). Dry quartzite models (Fig. 11a and d) result in higher temperature rise and compare well to geological observations. The temperature curves for γ̇Z -cst models (any t*) and τ-cst models (t* b 5 My)
show good correlation to the average peak temperatures of Svensen
et al. (2001) and to the local estimates of Souche et al. (2012).
The results obtained with dry quartzite (Fig. 11a and d) show
about 100 °C or more of lateral variation of peak temperatures within
10 km distance from the detachment. This clearly indicates the significance of shear heating by locally increasing the temperature in the
sediments at the vicinity of the detachment fault. These models reproduce the geological data of 295 °C in the Hornelen basin. In comparison to a maximum of 220 °C calculated without shear heating,
shear heating alone accounts for as much as 25% of the thermal budget of the sediments.
Figs. 9 and 11 show that the influence of shear heating depends
greatly of the distance to the shear zone, and is in accordance with
the geological observations from Souche et al. (2012). The numerical
results of the present study show that rocks located more than
5-10 km away from the detachment are not affected by shear heating.
5.3. PT-t path and retrograde overprint of rocks within the shear zone
The studies of PT-t paths representative for the exhumation of HP
units within the NSDZ show isothermal or temperature-increasing
decompression (Hacker et al., 2003; Johnston et al., 2007; Labrousse
et al., 2004) illustrated in Fig. 12. Can shear heating lead to temperature increase or support isothermal condition during exhumation of
the HP units? Below we demonstrate that shear heating is unlikely
an explanation for such an effect and other mechanisms should be
found to explain the phenomenon.
We compare these observations with our model results. We select
several markers with initial depth between 60 and 25 km and reconstruct
their PT-t paths. In our estimations, we use lithostatic condition to
calculate pressure, which makes our pressure estimates as upper bounds
(as the background extensional stress may induce underpressure) and
conclusions of this section more robust. Markers initially shallower than
45 km exhibit strong thermal overprint related to shear heating. The temperature increase during the exhumation of these markers, up to 100 °C,
compares with some of the geological data but is observed at much
shallower depths in the model. In contrast, markers initially deeper than
45 km, comparable to the initial decompression depth of the HP units,
monotonically cool during decompression.
The model results show the limited effect of shear heating in producing temperature increase at the early stage of the decompression
of the HP units found in the NSDZ. Either these units were originally
at shallower depth and were subjected to overpressure (Petrini and
Podladchikov, 2000), or were possibly affected by heat generation
during retrograde metamorphism reactions (Connolly and Thompson,
1989; Haack and Zimmermann, 1996).
6. Conclusions
We have developed a simple numerical model to understand the
thermal budget and the role of shear heating in crustal-scale extensional detachments. Using geological observations of the NSDZ we
have developed a conceptual model of evolution of such a system.
The model was augmented by constraining the non-steady-state
geotherm associated with the end of Caledonian orogeny. We have
built two simplified kinematic models for the deformation within
the shear zone. The model results were tested against peak temperature conditions documented within three Devonian basins of western
Norway. Our numerical study has shown:
1. Shear heating within the highly-deformed NSDZ can generate a
temperature increase up to 100 °C in the detachment shear zone
and is a necessary mechanism to explain the elevated peak temperatures within the Devonian basins of western Norway.
2. Shear heating may account for 25% of the thermal budget of the
supra-detachment basins and may affect the peak temperatures 5
to 10 km away from the detachment shear zone.
Please cite this article as: Souche, A., et al., Shear heating in extensional detachments: Implications for the thermal history of the Devonian basins
of W Norway, Tectonophysics (2013), http://dx.doi.org/10.1016/j.tecto.2013.07.005
12
A. Souche et al. / Tectonophysics xxx (2013) xxx–xxx
3. The amount of shear heating produced in the system is predominantly dependent on the exhumation rate and the rheological parameters of the rocks forming the shear zone.
4. Shear heating has a limited impact on the documented heating
during decompression of the HP units in the NSDZ.
Acknowledgements
The Norwegian Science Council provided basic funding though a “Centre of Excellence” grant to Physics of Geological Processes. Statoil and Det
norske viteskapsakademi are thanked for a Vista grant to AS. Det norske
oljeselskap is thanked for research grant to SM. The two reviewers, Prof.
Taras Gerya and Prof. Stefan Schmalholz, are thanked for their comments
and suggestions that helped improving the manuscript.
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