Tectonophysics 406 (2005) 197 – 212
www.elsevier.com/locate/tecto
Western Alpine back-thrusting as subsidence mechanism in the
Tertiary Piedmont Basin (Western Po Plain, NW Italy)
B. Carrapa a,*, D. Garcia-Castellanos b
a
Universität Potsdam, Institut für Geowissenschaften, University of Potsdam, Germany
Institute of Earth Sciences Jaume Almera (CSIC), Barcelona (previously at Vrije Univ. Amsterdam)
b
Received 29 September 2004; received in revised form 11 May 2005; accepted 24 May 2005
Available online 8 August 2005
Abstract
Basin formation dynamics of the Tertiary Piedmont Basin (TPB) are here investigated by means of cross-section numerical
modelling. Previous works hypothesised that basin subsidence occurred due first to extension (Oligocene) and then to
subsequent loading due to back-thrusting (Miocene). However, structural evidence shows that the TPB was mainly under
contraction from Oligocene until post Pliocene time while extension played a minor role. Furthermore, thermal indicators
strongly call for a cold (flexure-induced) mechanism but are strictly inconsistent with a hot (thermally induced) mechanism. Our
new modelling shows that the TPB stratigraphic features can be reproduced by flexure of a visco-elastic plate loaded by backthrusts active in the Western Alps in Oligo-Miocene times. Far-field compression contributed to the TPB subsidence and
controlled the basin infill geometry by enhancing basin tilting, forebulge uplift and erosion of the southern margin of the basin.
These results suggest that the TPB subsidence is the result of a combination of mechanisms including thrust loading and farfield compressional stresses.
D 2005 Elsevier B.V. All rights reserved.
Keywords: Western Alps; Subsidence; Sedimentary basins; Numerical modelling
1. Introduction
Investigation of the subsidence and structural evolution of syn-orogenic sedimentary basins allows the
basin formation kinematics to be unravelled. The
Tertiary Piedmont Basin (TPB), the southwestern extension of the Po Plain (Fig. 1), is located on the
* Corresponding author. Tel.: +49 331 977 5078.
E-mail address: carrapa@geo.uni-potsdam.de (B. Carrapa).
0040-1951/$ - see front matter D 2005 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2005.05.021
suture between the Alps and the Apennine belts and
was generated by post-collisional subsidence next to
the Alpine/Apennine orogen.
Subsidence and sedimentation of the TPB started at
the beginning of the Oligocene during a period of
important tectonic movement within the western Mediterranean area, including the opening of the Ligurian
Sea to the south (e.g. Gueguen et al., 1998) and the
formation, mainly in post Tortonian time, of the Apennine thrust belt to the east (e.g. Castellarin, 2001).
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B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
Despite these regional tectonic events and despite
being located on top of the Alpine/Apennine junction
on highly deformed basement (e.g. Miletto and
Polino, 1992), the TPB clastic infill is relatively undeformed and there are no major tectonic disturbances
separating the basin from its source areas (Carrapa et
al., 2003a). A few normal faults identified through
fieldwork investigation, mainly in Oligocene–early
Miocene sediments, have previously been considered
responsible for the early TPB evolution (Fig. 1).
These faults have been linked to the extensional
phase eventually responsible for the opening of the
Ligurian Sea (Mutti et al., 1995). These authors suggested an inversion in the stress field from extensional
to compression sometime in the late Oligocene–early
Miocene. However, the main Miocene subsidence of
the TPB has been associated with compressional tectonics, possibly related to the thrust activity developed
in the south-western Alps (Gelati et al., 1993; Roure et
al., 1990). Evidence of thrust tectonics in the Western
Alps together with Oligocene–Miocene syn-sedimentary compressional structures in the TPB (Carrapa et
al., 2003a; Hoogerduijn Strating et al., 1991; Schmid
and Kissling, 2000) suggest that the basin was mainly
undergoing NE–SW to NW–SE shortening since Oligocene time while extension played a minor role in
the evolution of the basin.
Despite previous qualitative geological observations on the tectonic evolution of the TPB, so far,
no attempt has been made to test and quantify basin
formation mechanism(s). Resolution of this problem
has important implications for the general evolution of
the Western Alps and associated sedimentary basins.
Cross-sectional numerical models are used as a
first approach to test the foreland basin-forming hypotheses and simultaneously advance the understanding of basin formation and kinematics in this
tectonically complex area. Two different flexural
199
models (Model I and Model II; after Zoetemeijer
(1993) and Garcia-Castellanos et al. (1997) are used
to test if the compressional scenario related to a
foreland basin setting can explain the subsidence history of the TPB.
The ultimate goal of this study is to check if a
flexural mechanism can explain the entire TPB subsidence, thus leading to a better understanding of the
tectonic and dynamic relationships between the TPB
and its surrounding areas and in general of the driving
mechanisms responsible for sedimentary basin formation in collisional tectonic settings.
2. Geological setting of the TPB and surroundings
areas
The TPB is located within the Internal Western
Alpine Arc, coincident with structures related to the
junction of the Alpine and the Apennine thrust belts
(Figs. 1 and 2). Geographically and kinematically the
TPB can be considered as the westernmost extension
of the Po Plain (Fig. 1). The area between the Western
Alps (including the Ligurian Alps) and the Northern
Apennines has experienced a complicated tectonic
history related to the oblique collision between the
European and African plates (e.g. Gueguen et al.,
1998; Laubscher et al., 1992; Schmid and Kissling,
2000; Schumacher and Laubscher, 1996). During Tertiary times this area was affected by N–S directed
convergence related to an important phase of intracontinental shortening (Platt et al., 1989). The arcuate
shape of the Western Alpine belt was accentuated after
35 Myr by postcollisional WNW directed motion and
anticlockwise rotation of the Adriatic microplate, associated with wedging of lower crustal slices (Schmid
and Kissling, 2000; Jimenez-Munt et al., in press).
Continental collision was responsible for the forma-
Fig. 1. A) Simplified tectonic map of the western Mediterranean region modified after Brunet et al. (2000); the inset box refers to Fig. 1B; B)
Structural map of the Tertiary basins of the Alps/Apennine junction (modified from Biella et al. (1997)). A-R.: Aiguilles Rouges Massif; DM:
Dora Maira; AGM: Argentera Massif; LA: Ligurian Alps; GP: Gran Paradiso; TH: Torino Hill; TPB: Tertiary Piedmont Basin; AM: Alto
Monferrato; M: Monferrato; BG: Borbera Grue. IL: Insubric line; RFDZ: Rio Freddo deformation zone; VVL: Villalvernia–Varzi line; VGT: Val
Gorrini thrust; SVZ: Sestri-Voltaggio zone; SF: Saluzzo thrust-fold. C) Location of the profiles used in this study. ECORS-CROP-Alp: seismic
profile; A–B: gravity profile after Miletto and Polino (1992) used in the crustal section of Fig. 2; profile C–D (seismic interpretation) after
Cassano et al. (1996) used in Fig. 2 and for modelling; profile D–E (seismic interpretation) after Pieri and Groppi (1981) used as support for
profile C–D of Cassano et al. (1996) in Fig. 2; Profile 1: seismic interpretation after Pieri and Groppi (1981) used for models I (see Fig. 4);
Profile 2 after stratigraphic reconstruction of Gelati et al. (1993) used for Model II (see Fig. 4). Dashed-dot line corresponds to the profile used
for Model I in this study.
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Fig. 2. A) Modified crustal scale interpretation of the ECORS-CROP-Alp profile (see Fig. 1 for location) from Stampfli and Marchant (1997) and Schmid and Kissling (2000) with
focus on the TPB infill (rectangle) showing the complicated tectonic setting of the study area (Fig. 2B). The ECORS-CROP-Alp profile has been integrated with the gravity model of
Miletto and Polino (1992) (A–B in Fig. 1) and with the profile of Cassano et al. (1996) (C–D in Fig. 1C).
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
tion of the Alpine/Apennine thrusts belts (e.g. Platt et
al., 1989; Polino et al., 1995) and related foredeeps
during Paleogene–Neogene times (e.g. Boccaletti et
al., 1999; Castellarin, 2001).
The TPB is bounded to the south by the Ligurian
Alps, which form the easternmost segment of the
Western Alpine Arc, to the west by the rest of the
Western Alpine domain, to the north and north-east by
the Po Plain and to the east by the Northern Apennine
(Figs. 1 and 2). In the south, the TPB unconformably
covers the northern margin of the Ligurian Alps collisional nappe stack. However, the present day southern limits of the TPB sediments are erosional and,
consequently, it is unknown how far the basin previously extended above the Ligurian Alps. In the west
the TPB is flanked by Plio-Quaternary sediments,
which in turn are bounded to the west by metamorphic
units belonging to the internal Western Alpine Alps.
The TPB succession dips northward underneath the
younger clastic sediments of the Po Plain (Dalla et al.,
1992; Schumacher and Laubscher, 1996), with the
exception of its easternmost part where it is truncated
by the Villalvernia–Varzi left-lateral strike-slip tectonic line (VVL in Fig. 1; Di Giulio and Galbiati, 1995
and references therein).
TPB sediments were thus deposited unconformably on both Alpine and Apennine units (Dela Pierre
et al., 1995; Piana, 2000; Piana and Polino, 1995;
Polino et al., 1995; Roure et al., 1990). The boundary
between these two units is roughly represented by two
structures: the N–S trending Sestri Voltaggio zone
(SVZ in Fig. 1) and the Villalvernia–Varzi left-lateral
strike-slip tectonic line (Boccaletti and Guazzone,
1970; Cortesogno and Haccard, 1979; Di Giulio and
Galbiati, 1995; Elter and Pertusati, 1973; Gelati and
Pasquaré, 1970; Haccard and Lorenz, 1979). The
Sestri Voltaggio zone formed before the Oligocene
and separates the ophiolitic metamorphic units of the
Voltri Group from the non-metamorphic flysch units
belonging to the Apennine domain. The Villalvernia–
Varzi left-lateral strike-slip tectonic line marks the
tectonic boundary between the Alpine and the Appennine domain prior to the late Oligocene and has
accommodated ongoing deformation in the Neogene
(Cavanna et al., 1989; Di Giulio and Galbiati, 1995;
Elter and Pertusati, 1973; Hoogerduijn Strating et al.,
1991; Laubscher et al., 1992). Subsidence and clastic
sedimentation started in the Oligocene with a south-
201
Fig. 3. A) Location of stratigraphic transects (Gelati et al., 1993)
used to construct subsidence curves (after Carrapa et al., 2003a); B)
subsidence rates modified after Carrapa et al. (2003a).
ward transgression, and continued until Pliocene time
(Gnaccolini, 1998 and references therein) with strong
sedimentation and subsidence rates during the middlelate Miocene (Carrapa et al., 2003a; see Fig. 3).
3. The structural evolution of the TPB
The Oligocene to early Miocene TPB basin was
lying between the subsiding Po Plain in the N (Clari et
al., 1995) and the extensional Ligurian-Provençal
Basin (present day Ligurian Sea) in the SW. Anisotropy of magnetic susceptibility (AMS) data measured
in TPB sediments together with structural evidence
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show that during this time span the TPB experienced
NE–SW shortening (Carrapa et al., 2003a). Paleostress analysis of small-scale normal faults detected in
Rupelian to Tortonian sediments suggests a fairly
homogeneous N–S extension over the entire basin
(Carrapa et al., 2003a) but the temporal relationship
between N–S extension and shortening is not clear. In
the same time span, evidence of thrusting towards the
southern Alpine foreland is widespread in the Western
(Schmid and Kissling, 2000) and Central Alps (Bernoulli et al., 1989). In particular, during the Oligocene
the Central Alps were affected by back-thrusting responsible for the formation of the Gonfolite Lombarda
(Bernoulli et al., 1989) which is the equivalent of the
northern Swiss Molasse Basin and can be possibly
considered the equivalent of the Tertiary Piedmont
Basin (Fig. 1). Also, during the Oligocene–early
Miocene, thrusts in the Ligurian Alps over the
TPB sediments are documented (Piana et al., 1997;
Hoogerduijn Strating et al., 1991) suggesting that
shortening in the belts was active during the TPB
earliest stages. During middle Miocene times, NE–
SW directed compression and limited shortening
remained active, producing syn-sedimentary structures, but no significant extensional features have
been detected (Carrapa et al., 2003a). In general,
evidence of Oligocene–Miocene back-thrusting with
E directed senses of shear has been detected in the
Western Alps (Bucher et al., 2003; Platt et al., 1989;
Roure et al., 1990; Laubscher, 1991; Choukroune et
al., 1990).
During late Miocene times, the western Po Plain
and the NW Apennines were undergoing ~N–S to
NE–SW directed shortening (Laubscher et al., 1992;
Schumacher and Laubscher, 1996; Boccaletti et al.,
1985). Serravallian and older sediments of the TPB
experienced NW–SE directed compression, which is
interpreted to be younger than the NE–SW directed
event, and therefore is post Tortonian (Carrapa et al.,
2003a). From middle Miocene to Pliocene times,
shortening formed the Saluzzo thrust-fold to the
west of the TPB (Fig. 1; Pieri and Groppi, 1981)
and was responsible for northward shift of the Alto
Monferrato thrust front (Falletti et al., 1995) to the
east of the TPB. Post Miocene uplift is responsible
for the present day TPB morphology (e.g. Lorenz,
1984) characterised by gentle hills up to 800 m, in
contrast to the flat morphology of the rest of the Po
Plain. By this time, Alpine back-thrusting was no
longer active and the Po Plain was being thrust
under the Northern Apennine (e.g. Castellarin,
2001). The preceding discussion clearly shows that
the evolution of the TPB has been mainly regulated
by contractional tectonic movements, with extension
playing only a minor role.
4. Possible subsidence mechanisms in the
formation of the TPB
Extension related to opening of the Ligurian-Provençal Basin has been proposed as mechanism responsible for the first period of subsidence by Mutti et al.
(1995). Subsidence in stretching basins is due to two
different mechanisms: a) fault-controlled initial subsidence caused by mechanical stretching of the upper
brittle layer of the lithosphere; b) thermal subsidence
caused by the cooling and contraction of the upwelled
asthenosphere. The stretching stage and the related
isostatic flexural response influence geometry, depth
and size of the basin (e.g. Cloetingh et al., 1992, Kooi
and Cloetingh, 1992, Mckenzie, 1978). Both structural
and thermal indicators in the TPB are inconsistent with
the two modes of subsidence related to stretching
basins. Low vitrinite reflectance (VR) values are between 0.2% and 0.7% and thermal alteration indices on
palinomorphs (T.A.I.) are always less than 2 for the
oldest TPB sediments (Molare Formation). These data
suggest temperatures generally lower than 100 8C (e.g.
Robert, 1988). Assuming a minimum basin depth of 4
km (present day maximum thickness) a maximum
palaeo-geothermal gradient of 25 8C/km is obtained;
this can be considered as a dnormalT gradient (in the
sense of Robert (1988)), which is typical of a foreland
basin setting (Allen and Allen, 2005) and suggest a cold
(flexure-induced) mechanism for the TPB subsidence.
Supporting data came from Apatite Fission Track Thermochronology on pebbles from the Oligocene Molare
Formation (Barbieri et al., 2003), which indicate that
these sediments did not experience temperatures higher
than 120 8C since the Oligocene. Therefore, structural
and thermal evidence strongly suggest that neither
brittle extension nor thermal subsidence, possibly related to an early stretching phase responsible for the
subsequent opening of the Ligurian Sea, is a likely
mechanism for the TPB formation.
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
Shortening driven by back-thrusting was reported
by Roure et al. (1990) within the south-western
Alpine orogen, and it has been proposed to be
responsible for the increase of subsidence in the
TPB during Miocene times (Gelati et al., 1993).
Lithospheric flexural loading due to emplacement
of thrust sheets is the mechanism responsible for
subsidence in foreland basins that are dynamically
linked to associated orogenic belts (e.g. Allen and
Allen, 2005; Beaumont, 1981; Dickinson, 1974;
Price, 1973). The deflection on the foreland plate
depends on a number of factors: a) flexural rigidity
of the flexed lithosphere (Mcnutt et al., 1988); b)
nature of the loads (topographic/thrust loads, horizontal and vertical forces, bending moments, sediment and water loads) (Royden, 1988; Sheffels and
MCnutt, 1986; De Celles and Giles, 1996); c) preexisting heterogeneities (e.g. Mugnier and Vialon,
1986); d) and the direction of the subduction zone
(Doglioni, 1994). The main architectural features
predicted by foreland basin models are: a) the
wedge-shaped geometry of the sedimentary units,
thick near the orogenic belt and thinning onto the
foreland; b) the lateral shift of the greatest thickness
of different formations, due to the movement of the
thrusts (the area of greatest subsidence shifts
through time); this results in shifting of the pinch
out toward the peripheral bulge; c) the increasing
dip of the basin infill with depth. However, if the
thrust system does not move towards the foreland
but rather towards the wedge, such as in the case of
out-of-sequence thrusting, then different features can
be expected (Garcia-Castellanos et al., 1997). The
high present-day elevation and rapid Cenozoic exhumation of the Alpine orogen (e.g. Hurford et al.,
1991; Carrapa et al., 2003b), coupled with strong
Oligo-Miocene subsidence in the TPB (e.g. Carrapa
et al., 2003a), suggest that the TPB could be the
result of flexural subsidence from orogenic loading.
The presence of Oligocene–Miocene back-thrusts
(Gelati et al., 1993) in the Western Alps points to
a link between thrust activity and TPB subsidence.
Further, supporting evidence for such scenario
comes from 40Ar / 39Ar and apatite fission track
thermochronology data, which show that the Western Alps (including the Ligurian Alps) underwent
rapid exhumation before and during the time of
TPB sediment deposition (Barbieri et al., 2003;
203
Carrapa et al., in press). Very high Oligocene cooling rates related to fast exhumation have been
detected in the Ligurian Alps (south westernmost
segment of the Alps) when sedimentation in the
TPB had already started (Barbieri et al., 2003).
These data clearly show a link between contemporaneous exhumation and loading of the Alpine belt
(margin of the TPB) and the subsiding TPB.
Strike-slip mechanisms have been considered as
important in the formation of the TPB by some
authors (Piana, 2000; Schumacher and Laubscher,
1996). However, the only strike-slip structure that
could be of regional importance is the Villalvernia–
Varzi left-lateral strike-slip tectonic line (VVL in
Fig. 1) but it was active mainly prior to late
Oligocene times and during early Miocene times
and was inactive during the middle Miocene,
which is the period of strongest subsidence
(Cavanna et al., 1989; Di Giulio and Galbiati,
1995; Elter and Pertusati, 1973; Hoogerduijn Strating et al., 1991; Laubscher, 1991). Therefore, the
absence of major active strike slip faults in the TPB
excludes strike-slip movements as a possible mechanism for the TPB subsidence.
In the following, flexural loading will be tested via
numerical modelling in order to further examine and
quantify the feasibility of such a tectonic scenario.
This study is the first to attempt a quantitative validation of the mechanism(s) responsible for the vertical
movements that have occurred in this tectonically
complex area.
5. Testing flexural subsidence
Two different flexural models for the TPB will
be tested below, the first (hereafter referred to as
Model I) consists of a static, pure-elastic model
responding to the present topography and additional
hidden loads (after Zoetemeijer, 1993). This model
provides a first check on general relationships between basin shape, thrust loading and lithospheric
rigidity. The second flexural model (hereafter referred as Model II) calculates deflection and sediment geometry through time as a result of viscoelastic flexural response to the kinematics of thrust
stacking and surface erosion and deposition (after
Garcia-Castellanos et al., 1997).
204
SE
NW
Profile 1
Pellice
2500
2000
m
Saluzzo Maira
1500
Stura Tanaro
Bormida Sp.
Bormida Mill.
1000
Savona
Western A
lps
0
SALUZZO
STURA
MAIRA
TANARO
235km
BORMIDA
Pro
Sl
file
1
file
Mid.-Upp. Pliocene
Lower Pliocene
Upper Miocene
Middle Miocene
Lower Miocene
2
Pro
Basement + Oligocene
5km
0km
148km
SW
NE
transect3
2500
transect2
transect1
m
2000
transect5
transect4
Profile 2
Tortona
1500
1000
500
250
0
170km
Sea
level
5km
Upper MioceneLower Oligocene
Basement
0km
36km
Fig. 4. Profiles 1 and 2 and related cross-sections. The cross-section in profile 1 corresponds to the seismic interpretation (line I) of Pieri and Groppi (1981), (the deeper horizon
corresponds to the base of the Miocene); the cross-section in profile 2 is based on the stratigraphic reconstruction of Gelati et al. (1993) (see also Carrapa et al., 2003a).
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
500
250
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
5.1. Pure elastic model (Model I)
This model calculates deflections for distributed
loads on a plate with variable thickness and rheology
(Zoetemeijer, 1993). The program can take into account different boundary conditions (continuous vs.
broken plate) and lithospheric stresses. The model
allows the insertion of extra forces such as bending
moment and vertical shear forces. These forces indirectly could represent the effect of slab pull even
though they underestimate the positive contribution
to the gravity anomaly given by the density contrast of
a slab pull (Zoetemeijer, 1993).
Two profiles have been modelled with a pure
elastic flexural model (Figs. 1C and 5): one NW–SE
along the Dora Maira Massif (profile 1: FG in Fig.
1C) and the other SW–NE along the Ligurian Alps
(profile 2: dashed line in Fig. 1C). These two profiles
have been chosen because both the Dora Maira and
the Ligurian Alps were affected by vertical movements during the TPB evolution (Carrapa et al.,
2003b). Profile 1 starts from Grenoble (NW) and
ends in Savona (SE). Line 1 (from Pieri and Groppi
(1981)) has been used to compare with the calculated
deflection produced by Model I (FG in Fig. 1C). Note
that the base of line 1 corresponds to the base of the
Miocene because of the lack of seismic data from the
Oligocene. Profile 2 starts from St. Martin (SW) and
ends in Tortona (NE) (Fig. 1C). Basement depths
deflection (km)
Model I-Profile 1
5
0
-5
Te15km
Te25km
-10
-15
0
200
100
distance (km)
300
Test 1
Test 2
basement deflection (broken plate at 100 km)
basement deflection (broken plate at 100 km+vertical force)
topography
Miocene base (from seismic interpretation)
data affected by thrust deformation
Fig. 5. Model I applied to profile 1 (Fig. 4). Broken plate at 100 km,
fixed topography, former passive margin configuration with a
palaeo-water depth of 50 m (no data are available on pre-Oligocene
water depths). Open dots indicate that the basement is affected by
post thrust-deformation (see Fig. 4, profile 1 after Pieri and Groppi
(1981)) and therefore these points cannot be considered as representative of the original TPB geometry.
205
obtained from the stratigraphic reconstruction of Gelati et al. (1993) (see also Carrapa et al., 2003a) have
been used to test the calculated deflection produced by
Model I.
Adopted density values are 2400 kg/m3 for sediment; 2600 kg/m3 for the load; 2800 kg/m3 for the
crust; and 3300 kg/m3 for the mantle (e.g. Zoetemeijer et al., 1993 and references therein). The load
density is representative of the orogen (constituted
by imbricated thrust sheets) while the density of the
sediments is representative of the unconsolidated
sediments deposited in the foreland basin. Forward
modelling has been conducted using Te values of 15
and 25 km (e.g Mcnutt et al., 1988; Royden, 1993;
Stewart and Watts, 1997). Test 1 on profile 1 (Fig. 5)
is based on the assumption of a plate broken at
x = 100 from the left edge of the belt. This assumption is necessary in order to reproduce the magnitude
of subsidence observed in the basin. The response to
surface loading in the case of a broken lithosphere is
about twice as large as that for a continuous lithosphere (Sheffels and MCnutt, 1986). The position of
the broken plate coincides with the margin of the
Adriatic indenter (see Laubscher, 1991, Stampfli and
Marchant, 1997) and the Insubric Line, assuming its
continuation toward the south (Schumacher and
Laubscher, 1996). Test 1 is unable to reproduce the
TPB subsidence and basement geometry (Fig. 5).
However, numerous previous modelling studies
(e.g. in the Apennines) show that topographic load
alone is often not enough to explain the observed
subsidence in foreland basins and extra forces
(dhidden loadT in the present applied model) are
used in flexural models (Royden and Karner,
1984). On the basis of these results, an additional
dhidden loadT, represented by a vertical shear-force
of 1.5 1012 N/m in the left model boundary at
x = 100 km, has been added in Test 2 on profile 1
(Fig. 5). Test 2 shows that the maximum depth of the
basal Miocene can be reproduced even though the
general geometry of that horizon cannot be
explained, probably because of the effect of post
thrust deformation (Profile 1 in Fig. 5).
Test 1 on profile 2 has been calculated assuming
that the plate is broken 50 km NW from the edge of
the belt, using a Te of 15 km and 25 km, respectively
(Fig. 6) and shows that the shape of the basin, as for
profile 1, cannot be reproduced, and that Model I can
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
deflection (km)
206
Model I-Profile 2
5
0
-5
-10
-15
Te15km
Te25km
0
100
200
300
distance (km)
Test 1
Test 2
basement deflection (broken plate at 50 km)
basement deflection (broken plate at 50 km+vertical force)
topography
Oligocene base (from stratigraphic reconstruction)
Fig. 6. Model I applied to profile 2 (Fig. 4). Broken plate at 50 km,
fixed topography, former passive margin configuration with a
palaeo-water depth of 50 m. Black dots indicate the observed
basement depth, as derived from the stratigraphic reconstruction
of Gelati et al. (1993) (see also Fig. 4, profile 2).
fit only one data point in the stratigraphy (Fig. 6).
Furthermore, there is almost no difference between
the subsidence obtained by applying the two different values of Te. Test 2 on profile 2 has been
calculated assuming the same boundary conditions
as test 1 but with an additional vertical force of
1.5 1012 N/m. Test 2 (Fig. 6) shows that the maximum deflection calculated by Model I can fit the
basement depth and that the best fit is obtained with
a Te of 25 km.
The tests presented above show a good correspondence between the area of maximum subsidence in
the basin and the area of maximum deflection generated by Model I. These results suggest that the belt
located NW of the TPB was deformed and uplifted
during the TPB formation, loading the basin. However, the potential deformational effect of blind
thrusting and compressional stresses could have
been responsible for the final TPB geometry. In
order to further explore this possibility, and to better
understand the kinematics of subsidence, Model II is
designed to account for non-instantaneous loading by
thrust deformation.
5.2. Visco-elastic model (Model II)
Foreland basin models generally predict a wedgeshaped geometry to the sedimentary units, with a
migration of the main subsiding area towards the
foreland through time and an increase of the dip of
the basin infill with depth (e.g. Allen and Allen,
2005; Royden, 1988; Sheffels and MCnutt, 1986;
De Celles and Giles, 1996). However, if out of
sequence (back)thrusts propagate towards the
wedge, instead of towards the foreland, or if the
anelastic properties of the lithosphere are accounted
for, then different features may be expected (GarciaCastellanos et al., 1997).
For Model II we use the program tAo (GarciaCastellanos et al., 1997, 2002), which allows for a
coupled simulation of thrust loading, surface processes (erosion and transport) and regional isostasy
(visco-elastic flexure), among other processes not
required in the present study. The source code is
open and available for public use at http://cuba.ija.
csic.es/~danielgc/. At each time step of the model
evolution, these three processes are calculated using
the finite difference technique, according to a forward modelling scheme. A cinematic, vertical-shear
approach is assumed for the movement of the hanging wall of each thrust (i.e., the units preserve their
vertical thickness during tectonic transport; for further details we refer thereafter to Garcia-Castellanos
et al. (1997). Flexural vertical motions of the lithosphere are calculated via a visco-elastic thin-plate
approach accounting for intraplate horizontal stresses. At each time step the distribution of incremental
vertical deflection Dw(x) is given by:
Dwð xÞ ¼ wVð xÞ þ BwW
BtDt
where Dt is the time step of the calculations (0.1 Myr
in this study) wV(x) and wU(x) are the elastic and
viscous components of deflection respectively that
result from solving the following differential equations
(modified after Nadai, 1963; Turcotte and Schubert,
1982):
d2
dx2
Dð xÞd 2 wVð xÞ
x2
Fx
d 2 wVð xÞ
þ qm g wVð xÞ ¼ Dqð xÞ
dx2
1
0
W
B2 Bw
B2 BwW
B2 @
BwW qð x; t Þ qm gwð x; t Þ
Bt
Bt A
¼
Dð xÞ
þ qm g
Fx
Bt
s
Bx2
Bx2
Bx2
where x is the position along the section, D(x) is the
rigidity of the lithosphere, F x is the horizontal intraplate force, g is the mean gravitational acceleration, q m
is the density of mantle, Dq is the increase of load
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
related to thrusting and erosion/sedimentation (measured in N m 2), q is the total cumulative load, w is the
total cumulative deflection, and s is the viscous relaxation time.
Rigidity D is related to the equivalent elastic thickness of the thin flexural plate Te, and relaxation time s
is related to viscosity l following these expressions:
ETe3
2d ð1 þ vÞd l
s¼
D¼
E
12ð1 v2 Þ
where E is Young’s modulus (7d 1010 Nd m 2), and m
is Poisson’s coefficient (0.25 for elastic flexure; 0.5
for visco-elastic flexure). Assuming zero moment and
vertical shear stress as boundary conditions of the
model (at x = 120 km and x = 200 km), this set of
equations allows calculation of the evolution of vertical deflections as a function of upper crustal mass
redistributions related to tectonic deformation and
erosion/sedimentation.
Sedimentation is applied assuming a laterally constant rate where accommodation space is available
(below sea level). Erosion is assumed to be proportional to elevation above sea level with a constant rate.
For the sake of simplicity, deformation of sediments is
not accounted for in this paper (they shift only vertically to accommodate the thrusting units, preserving
their thickness).
Fig. 7 shows the preferred model run among those
tested to examine the effect of out-of-sequence backthrusting on the stratigraphic features in profile 1 (the
best constrained profile). Two thrusts are defined, the
first being active from 33.7 Ma (Early Oligocene;
beginning of sedimentation in the TPB) until 23.8
Ma (base Aquitanian), and the second (further towards
the mountain belt) from 23.8 to 7 Ma (top Tortonian).
Both faults have a slip rate of 3 mm/yr. A maximum
sedimentation rate of 100 m/Myr is assumed under sea
level. In this run, erosion in the continental areas is
100 m/Myr for every km of topography. The initial
model surface (at 33.7 Ma) is flat.
Results (Fig. 7) show remarkable similarity between the model geometry obtained for the present
(after 33.7 Myr of tectonic shortening, isostasy, and
surface transport) and the stratigraphic features observed in the TPB (Fig. 4). In particular, Model II
reproduces the general shape of the basin and the
clastic infill, with onlap contacts and a thicker Miocene sequence towards the Western Alps (e.g. Saluzzo
207
Basin; Fig. 1) and tilting with erosional truncation
towards the Ligurian Alps (see Fig. 4, profile 1 for
comparison). This geometry results from viscous
stress relaxation within the lithosphere, which progressively reduces the basin width and increases its
depth and forebulge uplift (Garcia-Castellanos et al.,
1997). This feature cannot be reproduced assuming
pure elastic lithospheric flexure. The required relaxation times are in the range of 2–4 Myr. These values
are comparable with values derived in previous studies (e.g. Walcott, 1970; Lambeck, 1983; Turcotte and
Schubert, 1982; Garcia-Castellanos et al., 2002) ranging from 0.1 to 50 Myr, and imply reasonable values
of effective lithospheric viscosity of 1.5–3 1024
Pad s (common values obtained from lithospheric deformation modelling range between 1023 and 1025
Pad s; e.g. Walcott, 1970; Kukacka and Matyska,
2004; Marotta and Sabadini, 2003). In agreement
with these results, Jimenez-Munt et al. (in press)
have recently obtained viscosity values ranging from
1025 in the undeformed Adria domain to 1022–1023
Pad s in the Alps, using planform thin-sheet modelling
techniques that suggest a compressional to strike-slip
regime during most of the post-collisional history of
the TPB. To compensate for viscous stress relaxation,
a larger Te of 35 km, compared to the pure elastic
model, is required for this visco-elastic plate model.
The maximum subsidence produced by Model II (2.7
km; Fig. 7) is lower than that observed in the basin
(~3.5 km; Profile 1) because the latest Miocene–Pliocene TPB evolution is not considered in Model II.
From the late Miocene to the present, the TPB subsidence is probably the result of Apennine-related
movements responsible for the thrust of the Po Plain
under the Northern Apennine (e.g. Falletti et al.,
1995). For this reason, the infill predicted by the
model during this interval it is not fully comparable
with the present situation.
Far-field compressional stress also has a significant
effect on the predicted subsidence. The presence of a
compressional regime during basin development is
well supported by geological evidence (e.g. Carrapa
et al., 2003a), although the magnitude of the compression is difficult to constrain. Flexural studies of
basin formation usually yield values of 1012–2 1013
N/m (e.g., Van Wees and Cleotingh, 1996; Ershov,
1999; Garcia-Castellanos et al., 2002). A value of
5 1012 N/m has been chosen for Model II to test
208
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
Model II
eros./sed. (km) deflection(km)
-50
0
no compression
4
del II)
sion (Mo
compres
a)
8
erosion
sedimentation
-50
elevation
(km)
100
0
4
2
0
-2
0
-10
2
elevation (km)
50
Ma
23.8-7.0
0
33.7-
50
100
a
23.8 M
c)
ed
erod
d)
0
0
3
-2
-4
b)
7
11
15
19
24
23
28
32
forebulge uplift
flexural
subsidence
NW
-50
0
50
s
scou
ic+vi
elast in tilting
bas
SE
100
distance (km)
Fig. 7. Final stage (t = 0 Ma, present) of Model II. a) flexural vertical movements calculated for Model II (solid line) on profile 1 (Fig. 4) and
those predicted for an equivalent model without far-field compression but producing a similar final topography by reducing shortening rate
(dashed line). b) accumulated erosion (positive) and sedimentation (negative) at each location (solid line) (in equivalent meters of mother rock
density (2600 kg/m3)). c) true-scale final geometry of Model II. d) vertically exaggerated geometry of the basin infill in Model II. Note that the
absence of a compressive force would reduce the basin depth and tilting.
the effect of an external force on basin formation. In
the absence of this force, the predicted subsidence
under the thrusting blocks is reduced (Fig. 7a), requiring slower shortening rates in order to keep the
maximum topography in the model close to 2000 m.
In turn, this reduction in shortening rate results in a
shallower basin, therefore providing a worse fit to the
basin geometry. Although the results do not allow
more precise constraints on the force magnitude,
they do demonstrate that including such a compressive force better reproduces the basin geometry. Horizontal compression also significantly increases the
uplift along the external margin of the basin (forebulge), enhancing the tilting of the sediments towards
the Ligurian Alps discussed above. Note that no
additional (hidden) load has been applied in Model
II (in contrast to Model I), indicating that this additional load, if it exists, has a relatively small impact
on subsidence, and that the bulk of the subsidence can
be attributed to the effect of thrust loading, flexural
lithospheric response, and far-field compression. Viscous stress relaxation in the lithosphere is required in
our model to reproduce the observed basin tilting and
sedimentary infill geometry.
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
6. Discussion and conclusions
The location of the TPB in the area between the
junction of the Western Alps and the Northern Apennine is one of marked geodynamic complexity. Our
results suggest that this complexity precludes attribution of the generation of accommodation space to any
unique basin subsidence mechanism. Previous studies
have proposed both extension and compression as
basin formation mechanisms. Here we show that a
flexural subsidence model in combination of far-field
compressional stresses can explain the TPB evolution.
Results from flexural modelling show that both the
basement depth and sedimentary infill geometry can
be explained as a lithospheric flexural response to a
combination of thrust loading and far-stress field
compression. Support for this hypothesis comes
from Oligocene back-thrusting in the Western Alps
involving ~30 km shortening (Schmid and Kissling,
2000) and from Miocene syn-sedimentary shortening
in the TPB, which can be related to coeval western
alpine back-thrusting identified by Roure et al. (1990).
A phase of basin uplift associated with shortening in
post late Miocene time is indicated by the presence of
a north-west vergent thrust in the Saluzzo Basin and
by erosional truncation of the stratigraphic sequence
towards the Ligurian Alps (Fig. 4, profile 1). From the
late Miocene on, the TPB may have been influenced
209
by compressional which affected the NW Apennine
and was responsible for the formation of the Apennine
thrust front (e.g. Boccaletti et al., 1985; Falletti et al.,
1995). In late to post Miocene time, general uplift
affected both the basin and its margins (e.g. Dalla et
al., 1992).
Results of our model II suggest that sediment tilting and erosional truncation towards the Ligurian
Alps are related to basin tilting due to visco-elastic
relaxation of flexural stresses in the lithosphere (Beaumont, 1981; Garcia-Castellanos et al., 1997, 2002)
probably enhanced by far-field compression (Fig. 7).
Tilting and erosional truncation in the external,
uplifted zone of the basin is consistent with a viscoelastic lithospheric rheology (e.g. Beaumont, 1981).
Far-field compressional forces, which cause folding of
the lithosphere, also contribute to this basin infill
geometry.
The Oligocene–late Miocene geodynamic evolution of the TPB appears to be regulated by the complex interaction between vertical movements within
the Alpine orogen and basin subsidence under prevalent shortening. Although the basic trends of the basin
evolution are captured in a lithospheric flexural model
loaded by back thrusting, our results suggest that
vertical movements related to far-field compressional
stresses (i.e. crustal/lithospheric folding), and possibly
subcrustal additional vertical forces (e.g. slab pull) are
Fig. 8. Schematic palaeo-geographic/tectonic reconstruction of the Oligocene to Present TPB evolution. Black arrows indicate NE–SW
shortening active in middle Miocene and NW–SE shortening active in post Tortonian time. L: Ligurian Alps; TPB: Tertiary Piedmont
Basin; TH: Torino Hill; VVL: Villalvernia–Varzi line. Grey arrows indicate that NE–SW shortening was possibly still active after middle
Miocene time. +/ indicate subsiding/uplifting areas respectively (after Carrapa et al., 2003a; Bigot-Cormier et al., 2002 and Foeken et al.,
2003).
210
B. Carrapa, D. Garcia-Castellanos / Tectonophysics 406 (2005) 197–212
also present. While the rest of the Po Plain formed
mostly as a flexural response to the generation of
topography in either the Alps or the Apennines (Barbieri et al., 2004; Bertotti et al., 1997), the TPB is the
result of the interaction between shortening and farfield stresses in the Western Alps, suggesting a complex 3D interaction between multivergent compressional tectonics (Fig. 8).
Acknowledgments
This study benefited from fundamental scientific
input by Giovanni Bertotti and Sierd Cloetingh. Reini
Zoetemeijer, François Roure, Phillip Allen and Hugh
Sinclair and an anonymous reviewer are greatly
thanked for their constructive reviews. Marlies ter
Voorde and Jorge Gaspar-Escribano are thanked for
their helpful advice and contribution to modelling.
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