We use two-dimensional thermomechanical models to investigate the potential
role of rapid filling of foreland basins in the development of orogenic
foreland fold-and-thrust belts. We focus on the extensively studied example
of the Western European Alps, where a sudden increase in foreland
sedimentation rate during the mid-Oligocene is well documented. Our model
results indicate that such an increase in sedimentation rate will temporarily
disrupt the formation of an otherwise regular, outward-propagating basement
thrust-sheet sequence. The frontal basement thrust active at the time of a
sudden increase in sedimentation rate remains active for a longer time and
accommodates more shortening than the previous thrusts. As the propagation of
deformation into the foreland fold-and-thrust belt is strongly connected to
basement deformation, this transient phase appears as a period of slow
migration of the distal edge of foreland deformation. The predicted pattern
of foreland-basin and basement thrust-front propagation is strikingly similar
to that observed in the North Alpine Foreland Basin and provides an
explanation for the coeval mid-Oligocene filling of the Swiss Molasse Basin,
due to increased sediment input from the Alpine orogen, and a marked decrease
in thrust-front propagation rate. We also compare our results to predictions
from critical-taper theory, and we conclude that they are broadly consistent
even though critical-taper theory cannot be used to predict the timing and
location of the formation of new basement thrusts when sedimentation is
included. The evolution scenario explored here is common in orogenic foreland
basins; hence, our results have broad implications for orogenic belts other
than the Western Alps.
Introduction
The effects of surface processes on orogenic evolution have been intensively
studied over the last 3 decades (e.g., Whipple, 2009). Numerous studies
have shown that erosion can strongly influence the growth of orogenic
hinterland regions, with high erosion rates localizing deformation and
creating a lower, narrower orogenic wedge (Beaumont et al., 1992; Braun and
Yamato, 2010; Konstantinovskaia and Malavieille, 2005; Koons, 1990; Stolar
et al., 2006; Willett, 1999).
Both numerical and analog models also point towards a strong control
exerted by syn-orogenic deposition on the structural development of orogenic
forelands, as sedimentation rates affect the length of both thin- and
thick-skinned foreland thrust sheets, as well as the amount of displacement
taken up by individual faults (Adam et al., 2004; Bonnet et al., 2007;
Duerto and McClay, 2009; Erdős et al., 2015; Fillon et al., 2012;
Malavieille, 2010; Mugnier et al., 1997; Simpson, 2006a, b; Stockmal et al.,
2007; Storti and McClay, 1995). In particular, it has been shown
experimentally that higher rates of syn-orogenic sedimentation result in
longer thin-skinned thrust sheets as well as longer basement thrust sheets
under the foreland fold-and-thrust belt (e.g., Erdős et al., 2015; Fillon
et al., 2012). However, direct comparison of model predictions with
observations from natural case studies (e.g., Fillon et al., 2013) remains
scarce.
The North Alpine Foreland Basin of France and Switzerland developed in
response to continental collision in the Alps during early Tertiary time
(Dewey et al., 1973; Homewood et al., 1986; Pfiffner, 1986). The
stratigraphic infill of this foreland basin has been well documented (e.g.,
Sinclair, 1997; Berger et al., 2005; Kuhlemann and Kempf, 2002; Willett and
Schlunegger, 2010, and references therein) and consists of two major stages:
a Paleocene to mid-Oligocene deep marine (flysch) stage and a mid-Oligocene
to late Miocene shallow marine and continental (molasse) stage (Fig. 1).
Cross section and geological map of the Western Alps,
redrawn after Schmid and Kissling (2000) and Schmid et al. (2017), with the
inset showing the section interpretation of Roure (2008) along part of the
same deep seismic section. SL indicates the location of the Sesia–Lanzo
zone.
During the first stage, exhumation rates of the orogenic hinterland and
deposition rates in the foreland basin were low; hence, the basin remained
underfilled (Allen et al., 1991; Burkhard and Sommaruga, 1998; Sinclair and
Allen, 1992). At the onset of the second stage, both erosion rates in the
Alps and deposition rates in the foreland basin increased (Schlunegger et
al., 1997; Schlunegger and Norton, 2015; Sinclair and Allen, 1992), creating
an overfilled foreland basin. The transition from an underfilled to an
overfilled state coincided with a marked decrease in thrust-front advance
rate (Sinclair and Allen, 1992), but links between the two have remained
speculative.
Here, we use numerical models that build on our previous work (Erdős et
al., 2015, 2014) to test how an increase in sedimentation
rate affects mountain-belt and foreland fold-and-thrust belt evolution. In
earlier work (Erdős et al., 2015) we showed how our model predictions
were consistent with minimum-work theory. Here, we quantitatively compare
our models to critical-taper theory in order to assess the predictions of
this simple but widely used theorem, when including a more complex and
realistic rheology. Our main aim is to explore the potential causal
relationship between a sudden increase in sediment influx and the temporary
slowing of thrust-front propagation, as observed in the North Alpine
Foreland Basin. Such a sediment accumulation scenario is common in foreland
basins (e.g., Allen and Homewood, 1986); hence, the demonstration of a causal
relationship should have a significant impact on our understanding of not just
the North Alpine Foreland, but also the development of similar orogenic systems
around the world.
Numerical method
We explore the potential links between syn-tectonic sedimentation and orogen
structure through the use of 2-D arbitrary Lagrangian–Eulerian
thermomechanical modeling (Erdős, 2014; Thieulot, 2011) coupled to a
simple surface-process algorithm. The numerical experimental setup is very
similar to the one used in our previous studies (Erdős et al., 2015, 2014) and is explained in detail in the Supplement.
The thermomechanical model consists of strain-weakening frictional plastic
materials that allow for the localization of deformation (e.g., Huismans et al.,
2005). Our experiments use a four-layer crust–mantle rheology in which the
upper and lower crust as well as the upper lithospheric mantle undergo
frictional plastic deformation, while the middle crust and lower
lithospheric mantle exhibit power-law viscous creep (Fig. 2). A 3 km thick
pre-orogenic sediment package at the top of the model is separated from the
crystalline crust by a 1 km thick weak layer representing a décollement
horizon (e.g., an evaporite or shale layer). In order to include
self-consistent inherited extensional weakness zones, the model is first
extended, before the velocity boundary conditions are inverted to create a
contractional regime (e.g., Erdős, 2014; Jammes and Huismans, 2012).
(a) Model geometry showing layer thicknesses (including a
close-up of the crust), the position and size of the weak seed (pink
square), the lateral velocity boundary conditions (black arrows along the
sides of the box; note the ± sign), and the initial strength and
temperature profiles of the models. The material properties corresponding to
each layer (including the syn-tectonic sediments) are presented in Table 1.
(b) Frictional plastic strain softening is achieved through a
linear decrease in ϕeff from 15 to 2∘ with a
simultaneous decrease in C from 20 to 4 MPa. (c) Legend for
materials shown in (a).
Mechanical and thermal parameters used in the models for
each material.
The surface-process model includes an elevation-dependent erosion algorithm
as well as a sedimentation rule that fills topography up to a reference base
level at each time step. Both the erosion and sedimentation algorithms are
simple and do not conserve mass; however, the resulting basin-fill
geometries are consistent with observations from natural foreland-basin
systems (DeCelles and Giles, 1996).
The model experiments presented here have sufficiently high resolution (500 m
horizontally and 200 m vertically in the upper-crustal domain) to bridge
the large range of scales from the entire collisional orogen to the
fold-and-thrust belt and the interaction with syn-orogenic deposition.
The initial parameters (crustal setup, convergence velocity) have been
chosen to match conditions likely applicable for the Alpine orogenic system.
The major difference is that the model does not include seafloor spreading
or a lag between breakup and the onset of inversion. The sedimentation–erosion
algorithms have been parameterized to represent moderate rates for both
processes (see Supplement 1).
Model results
We present three model experiments that demonstrate the response of crustal
deformation to sudden temporal changes in syn-orogenic sedimentation. For
Model 1, neither erosion nor sedimentation is included (Fig. 3a–d). In Model 2,
a simple elevation-dependent erosion model is applied together with fixed
base-level sedimentation (Erdős et al., 2015): during each model
time step, basins are filled with sediments to a prescribed base level (Fig. 3e–g).
Model 3 is identical in setup to Model 2, but sedimentation is
initiated 10 Myr earlier and the base level of sedimentation is increased
during the experiment (15 Myr after initiation) to mimic the transition from
an underfilled to an overfilled foreland basin (Fig. 3h–i; see also
animations in the Supplement) as observed in the Western Alps (e.g., Sinclair and
Allen, 1992).
Model results. The material coloring scheme is identical
to that used in Fig. 2. All models are run for 65 Myr: 15 Myr (150 km)
extension followed by 50 Myr (500 km) contraction for a total net
contraction of 350 km. The horizontal scale for panels (a), (d), and (g) is the
same as that of panel (i). (a–d) Model 1 with no surface
processes, showing deformed Lagrangian mesh and isotherms after (a) 15 Myr
(Δx=-150km) and (d) 65 Myr (Δx=350km). Panels (b) and (c) are extracts from panel
(d) showing the small-scale deformation patterns in the foreland
fold-and-thrust belts. (e–g) Model 2 including a simple
surface-process algorithm filling up accommodation space until a base level
of -500m, showing deformed Lagrangian mesh and isotherms after
65 Myr
(Δx=350km). Panels (e) and (f) are extracts from
panel (g) showing the small-scale deformation patterns in the
foreland fold-and-thrust belts. (h–i) Model 3 including a simple
surface-process algorithm with the sedimentation base level changing from
-500 to 0 m at t=45Myr. Panels show deformed Lagrangian mesh and
isotherms after 65 Myr (Δx=350km). (h) An extract
from (i) showing the small-scale deformation patterns in the
pro-wedge foreland fold-and-thrust belt. Note that the polarity of
subduction is randomly oriented for each model. For ease of comparison we
flipped Models 2 and 3 to show them in the orientation that is conventional
for the Alpine cross sections.
Model 1
During the 15 Myr of initial extension, a broad, approximately 200 km wide
asymmetric rift basin is formed in the center of the model domain,
consisting of a number of rotated crustal blocks with mantle material
reaching the surface at two different locations approximately 50 km apart
(Fig. 3a). This is followed by a 15 Myr long inversion period culminating in
subduction initiation and the formation of an uplifted central block
(keystone structure) with a distinct internal structure consisting of a
number of inverted normal faults around a core of uplifted lower-crustal and
lithospheric mantle material (see Movie 1 in the Supplement).
In the third phase of Model 1, deformation migrates into the subducting
plate, building up the pro-wedge (using the terminology as defined by
Willett et al., 1993) initially through the formation of a crustal-scale pop-up
structure, and then primarily through an outward-propagating sequence of
basement thrust sheets (Fig. 3d) with an average thrust-sheet length of 52 km.
We use the term basement thrust sheet when referring to thrust sheets
that cut the crystalline basement (upper crust). Superposed on this
sequence, and often spatially slightly ahead of it, the pre-orogenic
sediments are also deformed, creating a complex thin-skinned fold-and-thrust
belt (Fig. 3b–c; for an extensive description of the interaction of
thin-skinned and thick-skinned deformation, see Erdős et al., 2015).
Deformation in the retro-side of the orogen (defined to be the part of the
wedge situated on the overriding plate) remains comparatively subdued
throughout the model but the initially uplifted central block, which
includes a lower-crustal–mantle lithospheric core, is transported more than
50 km onto the overriding plate.
Model 2, with erosion and sedimentation
The surface-process algorithms in Models 2 and 3 are activated at 45 and
35 Myr (model time), respectively. Consequently, all presented models exhibit
the same behavior during the first two phases described above.
Following the initiation of erosion and sedimentation at 45 Myr in Model 2,
sediment-loaded foreland basins form on both sides of the orogen, with more
intense thin-skinned deformation on the pro-side. The sequence of
outward-propagating basement thrust sheets in the pro-wedge is disrupted as
deformation remains localized on the active frontal basement thrust for 8 Myr,
instead of the 4 Myr observed in Model 1, before stepping out below the
foreland basin 13 km farther than in Model 1 (Fig. 3e–g; Movie 2 in the Supplement).
The effect can be well illustrated by comparing the length and displacement
of basement thrust sheets around the time of the onset of sedimentation
(Fig. 4). Prior to the onset of sedimentation, Thrust A accumulated 6 km of displacement
before Thrust B created a new, 45 km long basement thrust sheet (Fig. 4a). After the
onset of sedimentation, Thrust B remained active for about 8 Myr and accumulated 24 km
of displacement before Thrust C created a new, 83 km long thrust sheet in the
footwall of Thrust B (Fig. 4b).
The evolution of Model 2 around the time of the onset of
sedimentation (and erosion). The material coloring scheme is identical to
that used in Fig. 2. (a) Model 2 at 45 Myr (Δx=150km),
just before the onset of sedimentation. White marks show the length of
the active external basement thrust sheet (thrusting along Thrust B). The length is
measured using the VISU Grid (black grid advected with the materials in the
model); we counted the number of undeformed cells in the top row in the
basement between the old and the new frontal thrust. Red marks show the
amount of displacement along the last abandoned thrust (Thrust A). (b) Model 2
at 53 Myr (Δx=230km) at the time of the initiation of
the first basement thrust sheet after the onset of sedimentation. White
marks show the length of the active external basement thrust sheet
(thrusting along Thrust C). Red marks show the amount of displacement along the just-abandoned thrust (Thrust B corresponding to Thrust B in Fig. 4a). Further towards the
orogenic hinterland the steepened Thrust A is shown (corresponding to
Thrust A in Fig. 4a).
As the model progresses further, upper-crustal blocks in the internal parts
of the orogen that were initially covered with pre-orogenic sediments are
deeply eroded, reaching the surface and bringing the lower-crustal–mantle
lithospheric core of the central block to shallow depths. A small sliver of
mantle lithospheric material eventually reaches the surface along a
back-thrust (Fig. 3f).
We recorded maximum sedimentation rates for 2 Myr intervals (see the
alternating orange and green layers of syn-tectonic sediments in Fig. 3e–i)
throughout the model. After an initial peak of 2.7 kmMyr-1 between 45
and 46 Ma, when the entire available accommodation space is filled up to the
prescribed base level, the maximum sedimentation rates in the pro-foreland
basin stabilize around an average of 0.45 kmMyr-1.
Model 3, with erosion and intensifying sedimentation
The evolution of Model 3 is very similar to that of Model 2, even though
sedimentation and erosion start 10 Myr earlier. Significant differences can
only be seen between the pro-foreland basins, after the base level of
sedimentation is raised (simulated here by an increase in the sedimentation
base level over a 0.5 Myr period) to mimic the transition from an
underfilled to an overfilled foreland basin (Fig. 3h–i; Movie 3 in the Supplement).
The base-level change results in a temporary (approximately 2 Myr long)
increase in the maximum sedimentation rate in the foreland basin (from an
average of 0.45 to 1.1 kmMyr-1 at the location of the
frontal thrust). Subsequently, the maximum sedimentation rate quickly
decreases to its previous (average) value.
As observed in Model 2, the initiation of sedimentation alters the
architecture of the orogenic foreland by creating longer basement thrust
sheets. Similarly, a sudden increase in the sedimentation rate in Model 3
also results in a change in the foreland development. Again, this can be
well illustrated by looking at the deformation pattern around the time of
increase in sedimentation rate (Fig. 5). Prior to the increase in sedimentation
rate, Thrust A accumulated 10 km of displacement before Thrust B created a new, 45 km long
basement thrust sheet (Fig. 5a). After the increase in sedimentation rate,
Thrust B remained active for about 8 Myr and accumulated 22 km of displacement before
Thrust C created a new, 75 km long thrust sheet in the footwall of Thrust B (Fig. 5b).
The subsequent basement thrust-sheet sequence consists of longer thrust
sheets (on average 45 km instead of the previous 40 km) that are active for
longer times (on average 6.5 instead of 4 Myr) compared with the model
behavior before the increase in sedimentation rate (Fig. 3h–i;
Movie 3 in the Supplement).
The evolution of Model 3 around the time of increase in
sedimentation rate. The material coloring scheme is identical to that used
in Fig. 2. (a) Model 3 at 51 Myr (Δx=210km), the
time of increase in sedimentation rate. White marks show the length of the
active external basement thrust sheet (thrusting along Thrust B). The length
calculation method is the same as in Fig. 4. Red marks show the amount of
displacement along the last abandoned thrust (Thrust A). (b) Model 3 at 59 Myr
(Δx=290km) at the time of the initiation of the first
basement thrust sheet after the increase in sedimentation rate. White marks
show the length of the active external basement thrust sheet (thrusting
along Thrust C). Red marks show the amount of displacement along the just-abandoned
thrust (Thrust B corresponding to Thrust B in a). Further towards the orogenic
hinterlands the steepening Thrust A is shown (corresponding to Thrust A in a).
Comparison with critical-taper theory
We attempt to explain the observed behavior of our models at the scale of
the entire wedge in terms of critical-taper theory (Chapple, 1978; Dahlen,
1990; Davis et al., 1983). According to this theory, a wedge will evolve
towards a critical state characterized by being at the verge of brittle
failure both internally and at its base. As a consequence, equilibrium is
reflected by a self-similarly growing wedge with a stable surface slope
(α) and detachment dip (β) (Davis et al., 1983); such a wedge
should react instantaneously to changes in stress regime. Lateral variations
in the structure and surface slope of European Alpine foreland have been
explained using critical-taper theory (von Hagke et al., 2014, and references
therein). However, this purely brittle continuum-mechanical theory has
limited applicability to our model due to the presence of viscous plastic
deformation and strain-weakening materials (Buiter, 2012; Simpson, 2011).
Simpson (2011) argued that an elastic–plastic wedge is often well below the
critical stress threshold locally. Hence, we explore here whether the
large-scale deformation of our model orogens exhibits a behavior that is
consistent with critical-taper theory predictions.
When we consider a brittle Coulomb wedge, a sudden increase in sedimentation
rate will result in the filling up of the previously unfilled (or
underfilled) foreland basin, reducing α significantly while
moderately increasing β due to the loading of the basin.
Critical-taper theory predicts that such a sudden change in the taper
angles, without a simultaneous modification of the mechanical properties of
the wedge or the basal detachment, should drive the wedge towards a
subcritical state. Subsequently, the wedge needs to deform (thicken)
internally to increase its taper angle until it reaches critical state once
again (see also Willett and Schlunegger, 2010).
We analyze five models to assess whether our pro-wedges replicate the above
predictions of critical-taper theory. In order to isolate the potentially
tangled effects of erosion and sedimentation, we include in this analysis a
model with erosion but no sedimentation (Model 1.1) and one with
sedimentation but no erosion (Model 2.1). We define the wedge as the zone
between the surface trace of the frontal (thin-skinned) thrust and the
lower-crustal indenter of the overriding plate (denoted S point in Fig. 6).
The basal slope β is calculated using the top of the lower crust as a
reference horizon. We acknowledge that these definitions are arbitrary and
in some cases at odds with assumptions of critical-taper theory (i.e., the
top of the lower crust separates the ductile middle crust and the brittle
lower crust), but these definitions allow for a consistent derivation of
α and β values for each time slice in every model.
Example of α and β sampling routine.
S point: internal limit of the wedge considered for critical-taper analysis,
located at the tip of the lower-crustal indenter of the overriding plate.
Wedge tip: the outer tip of the wedge considered for critical-taper analysis, located
at the tip of the orogenic deformation zone. Red dots: elevation sampling points
along the wedge for a given sampling interval. For each sampling interval,
α is first calculated for every adjacent point (e.g., α11, α12) before we calculate the mean (α1)
of these local, individual α values for the entire wedge. The
process is then repeated for all sampling intervals (e.g., α21). Blue dots: depth sampling points along the wedge for a given sampling
interval. β is calculated in the same manner as α (described
above).
Due to the complexity of the surface topography (and to a lesser extent
the basal décollement), representing the entire wedge with a single
α–β pair is notoriously difficult. In this study, we
calculated multiple sets of α and β values along the wedge
using a range of different sampling intervals for every time slice of the
model (e.g., Fig. 6). Subsequently we calculated the mean α and
β values for each sampling interval and visualized the resulting mean
of these sampling intervals using box plots (see Fig. 7). This analysis
allows us to identify temporal trends that are persistent through a range of
characteristic length scales. We tested over 100 different
sampling intervals from 2.5 to 100 km and decided to use a subset of 41
of these, ranging from 10 to 30 km, to create the plots for this study. Note
that the trends described here were also present at the higher and lower
ends of the sampling scale.
Plots of α+β vs. model time for Models 1 (a), 2 (b),
and 3 (c). For each time slice, the
α and β values were determined using a range of sampling
intervals. The box plots present the average α+β, α,
and β values of these individual sampling intervals calculated for
the entire wedge. On each box, the central mark is the median, the edges of
the box are the 25th and 75th percentiles, and the whiskers extend to the most
extreme data points considered not to be outliers. The outliers are plotted
individually.
For brevity, we only discuss the implications of the above detailed
critical-taper analysis. The individual α, β, and α+β
vs. model time plots and their detailed interpretations can be
found in the Supplement, along with a detailed description of
Models 1.1 and 2.1. Generally, the models without sedimentation conform to
the predictions of critical-taper theory. After an initial mountain-building
phase, α+β stabilizes at a roughly stable level and is
only slightly perturbed around individual basement thrusting events (see
Fig. 7a). Erosion slightly increases α and reduces β, keeping
α+β at a constant value. The increase in α is a
result of the development of a narrower and steeper wedge with a narrower
foreland basin. Conversely, the decrease in β is partly due to
decreased topographic loading: models with erosion do not produce topography
higher than 6 km, while models without erosion can grow topography as high
as 8 km.
When sedimentation is included in the models, the behavior is considerably
more complex and the importance of the initiation of new thin-skinned
frontal thrusts becomes more pronounced (Fig. 7b and c). As the orogenic
foreland – and hence the wedge itself – grows wider, the crustal load
exerted by the orogen grows as well. The loading increases β until
the deformation moves to a new frontal thrust, further widening the wedge
and incorporating a previously undeformed, gently dipping basement, which
instantaneously reduces β. These cycles in β are superimposed
on top of a long-term decreasing trend, likely resulting from the wedge
becoming larger, warmer, and easier to deform over time. In the meantime, the
wide and low-relief orogenic foreland thrust belts generally decrease
α to very low (0.5–2∘) values.
The observed cyclic behavior, in which the deformation periodically
migrates to a new frontal thrust, is similar to the “punctuated thrust
deformation” described by Hoth et al. (2007) and Naylor and Sinclair
(2007), whereby the position of the deformation front fluctuates as
successive thrusts are gradually incorporated into the wedge. This discrete,
punctuated behavior causes the wedge to oscillate around a critical-taper
value rather than staying in complete equilibrium through time. Here we have
shown, moreover, how erosion and sedimentation influence this behavior
consistent with the predictions of critical-taper theory.
We have created animations showing the temporal and spatial (along-profile)
variations of α, β, an arbitrary metric of the shallow strain
rate, and the topography for Models 1 and 2 (see Movies 6 to
9 in the Supplement). Our aim with this exercise was to establish whether the changes in
topography (α, β) are driven by strain-rate changes or the
other way around. A key observation here is that the evolution of α
in Model 1 (and to a lesser extent in Model 2) shows a particular pattern: a
new thrust is activated after α of the region around the active
fault reaches ∼10∘. After the new thrust is
activated, this high α rapidly decays. This suggests that α≈10∘ can locally be seen as a critical value, which
triggers the formation of a new frontal thrust. This new thrust is generally
activated close to the tip of the active thin-skinned deformation front.
When sedimentation is included (Model 2), the high-α regions are
more persistent. We argue that, since the sediments are stifling the
foreland basin, there is significantly less room for thin-skinned
deformation that would otherwise create a gentler slope around the surface
trace of the basement thrusts. This results in negative-α basins
sliding between thick-skinned thrusts on top of the décollement. Our
thermomechanical models are therefore in agreement with the analytical
results shown by Willett and Schlunegger (2010).
Discussion
The first-order evolution of all three presented models is similar,
regardless of the imposed erosion–sedimentation scenario. First, an
asymmetric rift is formed with a wider and a narrower passive margin
consisting of rotated upper-crustal fault blocks on either side of an
upwelling sublithospheric mantle. This rifting phase is followed immediately
by the inversion of the large normal faults. After full inversion, a central
keystone structure is uplifted, with a crustal-scale thrust on either side
of it. As the rift was asymmetric, the keystone structure and its base are
also asymmetric. The subduction interface is consistently formed in the
basement of the initially wider passive margin. After the polarity of
subduction is established, a new basement thrust is formed in the subducting
pro-wedge lithosphere on average every 3.1 Myr (in the case of Model 1) in an
outward-propagating sequence. As the initial model setup is completely
symmetrical, the orientation of the initial asymmetric rift and, through
that, the polarity of the subduction are decided randomly. The main
differences between the models are the position and timing of thrust
activations.
The step-like migration of the deformation front and, to a lesser extent,
the distal edge of the foreland are present throughout all our model
experiments, but are enhanced when a change in the sedimentation history
occurs. In Model 2, the distal edge of the foreland basin advances rapidly
after the onset of sedimentation, while the basement deformation front
remains stationary (Fig. 8b). After this transitional period, lasting about
2 Myr, a new propagation order is established with longer basement thrust
sheets (on average 46 instead of 40 km) that stay active for longer times
(on average 7 instead of 4.5 Myr). In Model 3, two such transitional
periods can be observed (Fig. 8c): one at the onset of sedimentation (35 Ma;
see caption) and one at the increase in sedimentation rate (20 Ma; see
caption). During this latter transition, the distal edge of the foreland
basin rapidly advances again (approximately 150 km in 2.5 Myr), while the
outermost basement thrust remains active for 4 Myr longer than the previous
frontal thrusts (7.5 Myr instead of the previous 3.5 Myr).
Thrust-front propagation and sediment onlap on the distal
edge of the foreland basin vs. time (a) in the Western Alps
(redrawn after Sinclair, 1997) (b) derived from Model 2 and
(c) derived from Model 3. The thin dashed line in (b) and
(c) shows the thrust-front propagation pattern of Model 1. Note
that in (b) and (c) the time axis of the models is
reversed from Myr (forward model time) to Ma (time before “present”) to
fit the original axis of the Western Alps.
Conceptual figure showing the difference between the
evolution of a mountain belt with and without intensive late-stage
sedimentation. The cartoons are generalizations of our model results,
depicting them after the same amount of convergence.
In general, the location of a newly initiated in-sequence basement thrust
corresponds to the point at which the total work needed to slide on the viscous
mid-crustal weak zone and to break through the upper crust is lower than the
work needed to maintain deformation on the existing thrust front (Erdős
et al., 2015; Fillon et al., 2012; Hardy et al., 1998). Upon initiation (or
increase) of sedimentation in the foreland basin, the work required to
create a new basement thrust is suddenly increased as the sediments
effectively expand the thickness of the rock column overlying the
mid-crustal weak zone (Erdős et al., 2015). This increased resistance
against the formation of a new basement thrust breaks the previous cyclic
behavior and delays the propagation of the basement deformation front into
the foreland basin.
Comparison with the Alps
The models presented here capture a number of first-order features of the
Western European Alps (Schmid and Kissling, 2000; Schmid et al., 2017) (Fig. 1),
including the following: (a) a major step in Moho depth between the European and
Adriatic (or Apulian) plates; (b) strong decoupling between the upper and
lower crust, with the lower crust under-thrusting and subducting with the
mantle lithosphere; (c) stacking of basement thrust sheets in the central
part of the orogen; (d) shallow emplacement of lithospheric mantle material
in the retro-wedge, with a sliver of mantle material reaching the surface
(Fig. 3f, g, i), loosely resembling the Ivrea body and Sesia–Lanzo zone,
respectively; and (e) a generally asymmetric orogen with deformation
stepping out much further on the pro-side than on the retro-side. The
presence of a weak décollement below the pre-orogenic succession in this
model is also characteristic of the western Alpine foreland and allows for
the coexistence of thin-skinned and thick-skinned tectonics (see, e.g.,
Erdős et al., 2015), a feature that is much less prominent in the
Eastern Alps, where the décollement is absent (Schmid et al., 2004).
The initial extensional phase allows for the creation of physically
self-consistent inherited structural weakness zones, as observed in most
orogens. After extending the model for 15 Myr, the continental lithosphere
has effectively ruptured, creating two small separate ocean basins that
mimic the pre-orogenic presence of the Piemont–Ligurian and Valais basins in
the Alpine domain (Stampfli et al., 2001). It must be pointed out that
running the models further in extensional mode in this setup is not viable
because there is no built-in mechanism for the creation of oceanic
lithosphere. The effects of a thermal relaxation phase were not explored
either, as potentially important mechanisms like strain healing are not yet
implemented in the model.
The basement under the pro-foreland basin is rather smooth, dipping on
average 3∘ towards the orogen at the time slice captured in Fig. 4b
(which corresponds best to the present state of the North Alpine Foreland
Basin). This value is in good agreement with those inferred from the
interpretation of seismic reflection lines (Burkhard and Sommaruga, 1998;
Sommaruga, 1999).
The increase in sedimentation in Models 2 and 3 links basement thrust-front
propagation and the onlap of sediments onto the foreland, as observed in the
North Alpine Foreland Basin (e.g., Sinclair, 1997; Fig. 5). Both deposition
scenarios lead to longer frontal basement thrusts that remain active for a
longer period before a new basement thrust is formed. This suggests that
increased sedimentation, which resulted from the increasing relief and
changing climate in the Alpine hinterland (Schlunegger et al., 1997;
Schlunegger and Norton, 2015), was a significant factor in the mid-Oligocene
stalling of thrust-front advance observed in the western section of the
North Alpine Foreland. Note that this behavior is not observed further east
along the foreland where the amount of orogen-perpendicular shortening is
less and the decoupling salt layer is absent from the foreland basin (Schmid
et al., 2004). This could well limit the distance to which the thin-skinned
deformation of the foreland fold-and-thrust belt can reach.
The shortening rates, timing of orogenesis, and transition from an
underfilled
to an overfilled basin in the case of Model 3 are based on observations in the
northern foreland of the Western Alps. The timescales of thrust and basin
evolution of the models are comparable to those of the western Alpine
system. The jump in thrust-front position is of the order of 100 km both
in the model and nature, but the stagnation in the Alps lasted about twice as
long as observed in the models.
We also note that the stepwise behavior shown by Sinclair (1997) is present
in our models even if there is no change in the deposition scenario applied.
However, we argue that an increase in the amount of material deposited in
the foreland basin will necessary result in stalling of the basement
thrust-front propagation, while it will also allow for the distal edge of the
foreland basin to migrate further onto the downgoing plate.
Implications for other mountain belts
An early synthetic stratigraphic model of foreland-basin development
(Flemings and Jordan, 1989) showed that peripheral orogenic foreland basins
have a tendency to evolve from an underfilled into an overfilled state.
Numerous studies focusing on the stratigraphic infill of natural foreland
basins (e.g., Allen et al., 1991; DeCelles and Burden, 1992; Quinlan and
Beaumont, 1984) have demonstrated the merits of this model. Moreover, as the
internal part of the orogen grows, more surface area reaches higher
elevations, resulting in a potential increase in erosion rates and,
consequently, sediment flux into the foreland basin (Simpson, 2006a, b;
Sinclair et al., 2005). Hence, the orogenic foreland-basin evolution
scenario described in this study should be applicable to a wide range of
orogens around the globe. A prime example may be the southern Pyrenean
(pro-)foreland fold-and-thrust belt, where a middle Eocene increase in
sedimentation rate was accompanied by stalling of the thrust front (Sinclair
et al., 2005). Based on their stratigraphic models, Flemings and Jordan
(1989) proposed that changes in the rate of thrust loading, climate, or
source-rock lithology (all present in their models through surface-process
transport coefficients) can cause the shift from underfilled to overfilled
basins. Our model results imply that there is a strong feedback between
these potential controls and the state of the basin fill.
Conclusions
The thermomechanical models presented here provide first-order insights
into the intricate relationship between changing sedimentation rates and
deformation patterns in orogenic forelands (Fig. 9). Our models show that
a sudden increase in sedimentation rate disrupts thrust-front and
foreland-basin propagation patterns. The outermost basement thrust remains
active for a significantly longer time and accumulates more deformation than
previous thrusts developed during periods of lower sediment input, before
deformation steps out again under the sediment-loaded foreland basin. After
determining α and β values for each model and examining their
evolution over time, we conclude that they are broadly consistent with
predictions from critical-taper theory, despite the more complex and
realistic rheology included in our models. However, when sedimentation is
included, critical-taper theory cannot be used to predict the timing and
location of the formation of new basement thrusts.
The results are in good agreement with observations from the Western Alps
and the North Alpine Foreland Basin, where deformation remained relatively
stable for an extended period of time after the foreland basin shifted from
an underfilled to a filled–overfilled state. They should also be applicable
to other orogens around the globe.
Data availability
The data sets for this article are available as video
supplements in the Supplement and via the following DOIs: evolution of
Model 1 (10.5446/38571; Erdős et al., 2018a); evolution of Model 2
(10.5446/38572; Erdős et al., 2018b); evolution of Model 3
(10.5446/38576; Erdős et al., 2018c); evolution of Model 1.1
(10.5446/38573; Erdős et al., 2018d); evolution of Model 2.1
(10.5446/38574; Erdős et al., 2018e); α, shallow strain
rate,
and topographic evolution of Model 1 (10.5446/38575; Erdős et al.,
2018f); β, shallow strain rate, and topographic evolution of Model 1
(10.5446/38578; Erdős et al., 2018h); α, shallow strain
rate,
and topographic evolution of Model 2 (10.5446/38577; Erdős et al.,
2018g); β, shallow strain rate, and topographic evolution of Model 2
(10.5446/38579; Erdős et al., 2018i).
The supplement related to this article is available online at: https://doi.org/10.5194/se-10-391-2019-supplement.
Author contributions
ZE and RSH designed the experimental setup.
ZE ran the model experiments and all three authors contributed to the
interpretation of the results. ZE prepared the paper with contributions
from both coauthors.
Competing interests
The authors declare that they have no conflict of
interest.
Acknowledgements
We thank Fritz Schlunegger for his constructive comments at an early stage
of this project and Stefan Schmid for providing the latest version of the
Alpine cross section assembled by his group. We also thank Mary Ford and
Christoph von Hagke for their constructive feedback on a previous version of
this paper.
Review statement
This paper was edited by Bernhard Grasemann and reviewed by Xiaodong Yang and one anonymous referee.
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