Tectonophysics 463 (2009) 109–129
Contents lists available at ScienceDirect
Tectonophysics
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / t e c t o
Regional isostatic response to Messinian Salinity Crisis events
Rob Govers ⁎, Paul Meijer 1, Wout Krijgsman 2
Faculty of Earth Sciences, Utrecht University, Netherlands
a r t i c l e
i n f o
Article history:
Received 4 July 2007
Received in revised form 29 August 2008
Accepted 16 September 2008
Available online 25 September 2008
Keywords:
MSC
Paleogeographic reconstruction
Mediterranean tectonics
Messinian erosion surface
3D Flexure Shareware
a b s t r a c t
The salinity crisis of the Mediterranean during Messinian time was one of the most dramatic episodes of
oceanic change of the past 20 or so million years, resulting in the deposition of kilometer thick evaporitic
sequences. A large and rapid drawdown of the Mediterranean water level caused erosion and deposition of
non-marine sediments in a large ‘Lago Mare’ basin. Both the surface loading by the Lower Messinian
evaporites, and the removal of the water load resulted in isostatic/flexural rebound that significantly affected
river canyons and topographic slopes. We use flexure models to quantitatively predict possible signatures of
these events, and verify these expectations at well-studied margins. The highly irregular shape of the
reconstructed basin calls for a three-dimensional model. Near basin margins, plate-bending effects are most
pronounced which is why flexure is particularly important for a relatively narrow basin like the
Mediterranean. We focus on one specific sea level scenario for the Messinian Salinity Crisis, where most
of the evaporite load was deposited during a sea level highstand, followed by a rapid desiccation. Evaporite
loading at current sea level is expected to cause subsidence of the deep basins by hundreds of meters and
simultaneous uplift of continental parts of the margins. Differential uplift may lead to significant slope angle
changes and thus gravity flows. The relative scarcity of Lower Evaporite sequences along the margins may be
a result of these phenomena. Normal faulting of Lower Evaporite and older sediments and rocks is expected
on the margins. Desiccation enhances erosion of the freshly exposed continental shelf and slope. Subsidence
and riverbed sedimentation occurs on the continental margins, and significant uplift towards the basin
center. Reverse faulting is predicted at the margins. Finally, regional isostatic uplift following Zanclean
flooding is predicted to destabilize margin slope deposits, and to cause marginal uplift, river down-cutting,
and normal faulting.
© 2008 Elsevier B.V. All rights reserved.
1. Introduction
Surface processes act to move loads from one place to another. The
solid Earth responds by rock uplift or subsidence to this redistribution
of surface mass. These vertical (isostatic) motions in turn affect the
surface processes, for instance by increasing sedimentation where
subsidence occurs. From the perspective of the solid Earth, it is
irrelevant whether the surface loads are water, sediments or rocks; the
only thing that counts is the weight. Obviously, this is quite different
from a surface process perspective. For instance, an increase of
seawater depth at some location is quite different from sedimentation
with the former leaving quite different footprints in the geological
record than the latter.
One of the most pronounced natural experiments with surface
processes has been occurring in the Mediterranean realm. Here, the
deposition of thick evaporitic sequences and the drying up of most of
⁎ Corresponding author. Tel.: +31 30 253 4985; fax: +31 30 253 5030.
E-mail addresses: [email protected] (R. Govers), [email protected] (P. Meijer),
[email protected] (W. Krijgsman).
1
Tel.: +31 30 253 5091; fax: +31 30 253 5030.
2
Tel.: +31 30 253 1672; fax: +31 30 253 1677.
0040-1951/$ – see front matter © 2008 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2008.09.026
the basin (desiccation) has been established by numerous studies.
These events occurred during the Messinian stage of the Late Miocene,
which is why they are commonly referred to as the Messinian Salinity
Crisis (MSC). These giant evaporite bodies were discovered in the early
1970's by shipboard scientists of the Deep Sea Drilling Project (Hsü
et al., 1973). Numerous papers and working hypotheses have
subsequently contributed to a better understanding of the complex
and enigmatic scenario of rapidly changing depositional environments that governed this MSC. Most current interpretations converge
to a scenario of progressive isolation of the Mediterranean in a twostep model (Clauzon et al., 1996; Krijgsman et al., 1999; Roveri et al.,
2001; Rouchy and Caruso, 2006), but the precise course of events is
still a matter of serious debate.
Astronomical tuning of Messinian sedimentary sequences resulted
in a high-resolution time frame indicating that the MSC started at
5.96Ma (Krijgsman et al., 1999) with the deposition of the Lower
Gypsum unit (Lower Evaporites) all over the Mediterranean at
marginal settings. Field evidence and quantitative analyses suggest
that Mediterranean sea level remained at the level of the Atlantic
during Lower Gypsum formation which lasted at least until ~ 5.6Ma
(Krijgsman et al., 1999; Krijgsman and Meijer, 2008). The Lower
Evaporites are classically described from Sicily to comprise evaporitic
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R. Govers et al. / Tectonophysics 463 (2009) 109–129
limestone, gypsum and massive halite (Decima and Wezel, 1973), but
also consist of detritic mass deposits, along with re-sedimented
evaporites (Roveri and Manzi, 2006). The detailed correlation of the
Sicilian sequences to the seismic profiles of the abyssal Mediterranean
is still a matter of serious controversy (e.g., Roveri et al., 2006). The
Sicilian lower evaporites are overlain, with an erosional unconformity,
by the Upper Evaporitic sequences (~ 5.5–5.33Ma) which comprise
gypsum as well as brackish-water sediments of so-called “Lago Mare”
facies. These Upper Evaporites are generally thought to have been
deposited at low sea level when the Mediterranean was isolated from
the Atlantic (Hsü et al., 1973; Orszag-Sperber, 2006). Both the timing
and amount of sea level drawdown are controversial. The MSC ended
at 5.33Ma with the Zanclean (Pliocene) flooding of the Mediterranean,
which re-filled the basin within a few thousand years (Blanc, 2002).
The cause for re-establishing an ocean connection is speculative and
quite enigmatic.
Like today, the Mediterranean during the Miocene was most likely
an evaporation basin, meaning that in order to preserve sea level
evaporative water losses needed to be balanced by inflow of water
from rivers and the Atlantic Ocean. During the MSC, as tectonics acted
to cork the strait connections with the Atlantic Ocean, water influx
into the Mediterranean basin no longer sufficed to maintain sea level.
This resulted in a near-synchronous desiccation of the entire
Mediterranean basin within a couple of thousand years (Meijer and
Krijgsman, 2005). Krijgsman and Meijer (2008) and Govers (submitted
for publication) present evidence that this occurred during the last
stage of deposition of the massive halite (top of the Lower Evaporites).
Most workers today agree that there was only one (major) drawdown
event. Desiccation resulted in hill-slope instabilities and erosion,
leading to the Intra-Messinian Unconformity (or “Messinian Erosion
Surface”) between the top of the Lower Evaporites and the Upper
Evaporites. The amount of sea level drop is only incompletely
constrained by observations. Seismic images offshore show that the
base of erosional canyons and continental fans are in continuity with
the Messinian evaporites (Lofi et al., 2005), giving estimates of a paleoshoreline at − 2500m. Estimating the exact amount of sea level drop,
however, requires that the isostatic uplift is quantified, once postMessinian sediments and the sea water column have been removed,
as well as the isostatic subsidence because of evaporite loading — as is
being done in flexural backstripping.
Observations and discussions on the MSC in the decades following
its discovery (Ryan, 1969; Hsü et al., 1973) have focused on the
involved surface processes. Norman and Chase (1986) were the first to
recognize the potential impact for the solid Earth of these surface
events, and the possible feedbacks of basement vertical motions to
river discharge and incision. Similar, but more detailed studies of
basement motions of the Gulf of Lions/Provence region and incision of
the Rhone river where conducted by Gargani (2004a,b). In view of the
strong coupling between surface processes and basement isostatic
motions, we feel that the relative lack of attention for the solid Earth
role represents a hiatus in our understanding of the MSC. Therefore, it
is the purpose of this paper to estimate the solid Earth response to
MSC events, and to predict how surface processes likely were affected.
Based on our new three-dimensional approach, we show that
previous work substantially overestimated the regional isostatic
deformation. We predict hundreds of meters of basement motions
in most of the Mediterranean realm. We pay particular attention to
possible errors in the model, and how they affect our predictions.
In the first part of this paper we present a paleogeographic
reconstruction of the Mediterranean basin, along with the most
pertinent observations and arguments that we used to build it. In the
second part we present our estimate of the “effective elastic thickness”
during the MSC, which is a mechanical parameter in our isostatic
calculations. In the third part we briefly introduce the differential
equation, the numerical method, and boundary conditions. In the
fourth part we investigate the response to two different Lower
Evaporite scenarios. In part five we present results of different
desiccation scenarios. In part six we discuss the model results, and
their uncertainties, along with expected and observed responses at
well studied Messinian localities.
One consequence of the (still ongoing) tectonic activity in the
Mediterranean domain is that its horizontal extent and bathymetry
was different during the Messinian (e.g., Jolivet et al., 2006). As our
model predictions are sensitive to basin topology, we start by
presenting a reconstruction in the next section. We will argue below
that the western and eastern Mediterranean had a geometry that was
very similar to the modern geometry, and that important changes took
place in the central portion. We recognize, however, that available
observations leave room for alternative reconstructions, particularly
in the central Mediterranean. We will deal with these uncertainties by
also calculating the isostatic response in a Current geometry.
The sensitivity of the model calculations to (potential errors in) the
reconstruction will thus become clear as the difference between
the two models predictions; one for the reconstructed basin, and the
other for the current basin.
2. Paleobathymetric reconstruction
Fig. 1a shows our reconstruction of the Mediterranean basin for the
Messinian. Realizing that significant portions of the region were little
affected after the Messinian, we took the Current topography (and
bathymetry) and modified it where changes did occur. We used rigid
block rotations to make alterations, taking the topography within a
block and rotating it into its new position. The limitation of such an
Euler pole approach is that it does not capture block-internal strains.
However, well-dated total strain data are too scarce to sufficiently
constrain better reconstructions everywhere. The time of reconstruction is 5.65 Ma, midway between the beginning (5.95 Ma) and end
(5.33 Ma) of the Messinian salinity crisis.
The eastern Mediterranean's Ionian and Levantine basins are relics
of the Mesozoic oceanic Tethys (e.g., Hsü and Bernoulli, 1978). By
contrast, the western Mediterranean (including the Tyrrhenian basin)
and the Aegean Sea basin are young basins. The largest part of the
post-MSC motions within the Mediterranean basin differed substantially from the slow, overall convergence between Nubia and Eurasia
(Sella et al., 2002).
2.1. The Mediterranean basin west of Corsica–Sardinia
Our choices of block boundaries and reconstruction poles are based
on the following arguments.
1. Sicily and Corsica reached their Current positions relative to stable
Europe well before Messinian time. This inference is based on
paleomagnetic findings (Montigny et al., 1981; Pares et al., 1992;
Vigliotti and Langenheim, 1995) and on the end of volcanism on
Sardinia (Beccaluva et al., 1985).
2. Opening of the Valencia Trough was accompanied by rotations of
the Balearic Islands. Available paleomagnetic observations indicate
that rotations ended in the Late Miocene to Pliocene (Montigny
et al., 1981; Freeman et al., 1989; Pares et al., 1992). Watts and Torne
(1992) and Negredo et al. (1999) use geohistory analysis to date
Valencia rifting before 8Ma (but the data allow some extension
after that). The age of the youngest calc-alkaline volcanics (Lonergan and White, 1997) is consistent with our assumption that
extension ended well before the MSC.
3. Docking of Kabylian terranes onto the Algerian/Tunesian margin
was finalized during the Tortonian (Lonergan and White, 1997;
Carminati et al., 1998).
4. Post-MSC motion and deformation of the Alboran domain relative
to Iberia was small. Paleomagnetic rotations in both the Internal
and External zones of the Betic Cordillera are widely held to have
R. Govers et al. / Tectonophysics 463 (2009) 109–129
111
Fig. 1. (a) Reconstructed geometry and bathymetry (in meters) of the Mediterranean region during the Messinian period. (b) Effective elastic thickness (km) of the Mediterranean
region during the Messinian.
ended during the Middle Miocene (Platzman, 1992; Allerton et al.,
1993; Allerton, 1994; Lonergan and White, 1997; Calvo et al., 2001).
Stratigraphic backstripping results suggest an even earlier end of
the (Burdigalian/Langhian) (Watts et al., 1993). We agree with the
conclusion of Lonergan and White (1997) that “shortening in the
External Zones and extension in the Alboran Sea had terminated by
the Late Tortonian, suggesting … that the Alboran Sea basin had
achieved its present-day dimensions”. The 15–7Ma age of calcalkaline volcanics in the Alboran Sea and the Betic-Rif Cordillera
(compilation in Lonergan and White, 1997) is consistent with this
interpretation.
Consequence of these starting points is that the shape of the
reconstructed basin west of Corsica and Sardinia is nearly identical to
the current shape.
To constrain the pre-MSC basin depth in our reconstruction, we
note that the post-Miocene sediment thickness is mostly less than ~
1500m, and we estimate the Messinian evaporites to be 1500m thick.
A backstripping calculation shows that these sediments would have
left the basin shallower by ~ 900m since the onset of the MSC. The
other aspect to account for is the ageing of the oceanic sea floor, which
resulted in basement sinking. The average age of the Mediterranean
basin west of Corsica and Sardinia at 5.96Ma was around 14Myr
(Faccenna et al., 1997). A half-space cooling model predicts about
300m of isostatic sinking since the start of the MSC. The net outcome
is that, relative to the current bathymetry, we estimate this basin to
have been b 600m deeper at the onset of the MSC. Our estimate is
consistent with the findings of Ryan (1976) and with paleodepth
estimates as compiled by Dercourt et al. (2000). Below, we address the
impact of uncertainties in our paleodepth estimates.
2.2. The Mediterranean basin east of Adria
1. For our reconstruction of the Aegean region and the Hellenic
Trench, we backrotated the Peloponnesus using the results of Van
Hinsbergen et al. (2005) and Reilinger et al. (2006).
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R. Govers et al. / Tectonophysics 463 (2009) 109–129
2. Post-Miocene westward extrusion of Anatolia south of the North
Anatolian Fault was removed using the pole of McClusky et al.
(2000), which is consistent with paleomagnetic data of Kissel et al.
(2003).
3. We removed the post-Messinian back-arc extension of the Sea of
Crete based on references in Van Hinsbergen and Meulenkamp
(2006). Pliocene opening of the Sea of Crete occurred synchronous
and parallel with sinistral strike slip along the Pliny and Strabo
trenches (Le Pichon and Angelier, 1979; Van Hinsbergen et al., 2007).
4. High-resolution tomography indicates the existence of a NE
dipping slab beneath the Isparta triangle (Wortel et al., 2006;
previously believed to be detached by McKenzie 1978). We
interpret this to have resulted from convergence between the
Tauride Mountains and central Anatolia. The Late-Miocene–
Pliocene Aksu thrusting phase in the Isparta Triangle is interpreted
as a surface manifestation of this convergence (e.g., Poisson et al.,
2003). In our reconstruction of the Mediterranean Basin (Fig. 1),
this shows as a NW–SE trending basin within Anatolia that is an
artifact of our block rotation method. There is no point in trying to
remedy this for the purpose of the work in this paper.
5. Back rotation of the Africa margin increases the north–south extent
of the Eastern Mediterranean. The oceanic lithosphere (roughly
south of Crete and Cyprus) has a Mesozoic age, so that postMiocene isostatic sagging due to ageing was insignificant.
6. As the Nile Delta is largely a post-Messinian feature (Salem, 1976;
Ross and Uchupi, 1977), we used the results of Segev et al. (2006) to
estimate the Messinian bathymetry of the easternmost part of the
basin.
7. The oldest sediments in the Mediterranean Ridge are of Early
Miocene age (Robertson and Kopf, 1998). About half of the current
sedimentary volume was deposited after the beginning of the MSC
(Kopf et al., 2003). The spatial distribution of pre-MSC sediments is
however not well constrained. For reconstructing the paleobathymetry, we therefore took the simplest possible assumption, that
these sediments had a uniform thickness.
2.3. The central portion of the Mediterranean basin
Reconstructing the Central Mediterranean basin (from Corsica–
Sardinia to the Adriatic) we assumed the following;
1. The Apulian Platform remained in its current position. Structural
studies in the Gargano Promontory, in the Adriatic Sea north of
Apulia (Bertotti et al., 1999; Casolari et al., 2000), have shown that
normal NE–SW compression and shortening were interrupted in
the latest Messinian–middle Pliocene by NW–SE extension followed by NW–SE directed contraction. The net outcome is limited
post-Miocene horizontal deformation in the Gargano basin.
Bertotti et al. (2001) estimate the paleo-waterdepth of what
today constitutes the foredeep at 2–3km. Vertical deformation of
the foredeep began in the Pliocene. The Adriatic crust consists of a
crystalline basement (Finetti et al., 1987) that is covered by a thick
sequence of post-Devonian sediments (Ricchetti et al., 1988), and is
most often considered continental. Di Stefano et al. (submitted for
publication) use high-resolution tomography to trace the intersection line of the Adriatic (and Apulian) Moho surface and the
Tyrrhenian Moho surface. They find that this line follows the
culmination of the Central and Southern Apennines. We interpret
this as the location of the Adriatic passive margin, west of which
the Adriatic oceanic basin existed during the Messinian.
2. That the opening of the Sicily Channel occurred after the Messinian,
based on structural studies (Reuther and Eisbacher, 1985; Gardiner
et al., 1995; Catalano et al., 1995) and age of most of the volcanism
(Rotolo et al., 2006).
3. Like Dercourt et al. (1986), we assume that a deep, Tethyan ocean
extended east of the Apenninic subcontinent. This assumption is
motivated by a geodynamic argument (Royden, 1993), that
retreating subduction boundaries likely cease shortly after entry
of the Adriatic passive margin into the subduction zone. Pliocene–
Quarternary rotation of the Apennines is taken to be the result of
subduction and rollback of the oceanic Adriatic slab. The paleobathymetry of this oceanic domain is taken to be similar to that of
the eastern Mediterranean oceanic domain.
4. That the opening of the Tyrrhenian basin, including Marsili and
Vavilov and Paola sub-basins, occurred mostly after the Messinian.
This is based on the following arguments; Nicolosi et al. (2006)
demonstrate that Marsili opened between 1.6 and 2.1Ma; Vavilov
basin formation occurred after the MSC (Kastens and Mascle, 1987;
site 651 and 655, core data reveal lower to middle Pliocene oceanic
crust and sediments). Rifting and the onset of the Tyrrhenian Basin
commenced during the Tortonian (Kastens et al., 1988). Most of the
extension and opening of the Tyrrhenian Basin had yet to occur at
the onset of the MSC, as documented by the post-MSC age of both
the Marsili and Vavilov basins.
5. To reconstruct the shape of the overriding Apenninic crust, we note
that available data insufficiently constrain such attempt. We
therefore treat the Apennine crust as a single tectonic unit, in
spite of geodynamic objections to this choice. We assumed that
narrowing of the peninsula due to Apennine thrusting is balanced
by widening due to the development of foreland basins. The
reconstructed shape of the Apenninic subcontinent is therefore
similar to its current shape.
6. That Calabria constituted the overriding crust that independently
tracked roll-back of the Ionian slab from its pre-Messinian position
close to Sardinia (Bouillin et al., 1986).
3. Effective elastic thickness
A critical element in our prediction is the thickness of the flexing
plate. This is typically not identical to the thickness of the lithosphere, as
the lithosphere consists of high- and low-viscosity layers (e.g., Burov and
Diament, 1995). For our first order assessment of the flexure, we assume
that the lithosphere is made up of two types of elastic-viscous layers; 1.
Layers in which the viscosity is so high that no relaxation occurs over the
time span of the MSC and which therefore are elastic, and 2. Lowviscosity layers in which deviatoric stresses have fully relaxed and which
therefore do not support surface loads. In this simplified view, the
accumulative thickness of all elastic layers (the effective, or equivalent,
elastic thickness) is relevant for the bending problem; the effective
elastic thickness (EET) is consequently less or equal than that of the
lithosphere (Watts, 2001). Our view is simplified in two respects; first,
the shallow lithosphere has a finite (brittle) strength. Secondly, part of
the elastic-viscous layers have a Maxwell time that is similar to time
scale that we consider here, i.e., between a few thousand and 0.5Myr. As
a physical process, the development of lithospheric flexure is thus
expected to be time dependent. However, we are not aware of
observations that document this time evolution for the Mediterranean
MSC, which is why we adopt a simple elastic model that aims to predict
the flexural response at a few thousand years after loading, when most
of the transient signatures have vanished.
EET's are commonly inverted from an observed lithospheric
response to (known) surface loads, like from sediments or volcanoes.
As it is the purpose of this paper to estimate lithospheric response
from surface loads, we cannot use this approach. We therefore use a
similar procedure as for our reconstruction in Fig. 1, i.e., we take EET
“measurements” (from Bouguer coherence) for the Current Mediterranean basin, and modify it only in parts where significant changes
occurred since the Pliocene. These “gravimetric” EET's represent the
long-term strength of the lithosphere and are thus consistent with our
elastic model approach.
Fig. 1b shows our estimate of the effective elastic thickness, based
on the Current EET model of Perez-Gussinye and Watts (2005). The
R. Govers et al. / Tectonophysics 463 (2009) 109–129
reconstruction shows the signature of Miocene back-arc extension
due to slab roll-back in the western Mediterranean (Malinverno and
Ryan, 1986) in that the effective elastic thickness is representative for a
15Myr oceanic lithosphere (10km). The eastern basin is presumed to
be the last remnant of the Mesozoic Neo-Tethys ocean (Catalano et al.,
2001; Hafkenscheid et al., 2006). In the southeast Mediterranean, a
surface heat flow of about 50mW/m2 (Eckstein, 1978) is indicative of
approximately 80Myr old lithosphere. As a consequence, the effective
elastic thickness of the lithosphere was relatively large — we assume
30km (Caldwell and Turcotte, 1979; ~ 600°C isotherm). Consequences
of uncertainties in these EET estimates will be addressed in the last
part of this paper.
4. Model setup
The response of the entire planet to surface (un)loading is elastic at
first, followed by viscous relaxation. Within the first few thousand
years after the loading, stress relaxation occurs particularly in the low
viscosity asthenosphere. This effectively transfers the support for the
surface load to the more viscous layers, i.e. to the lithosphere. This
process dominates in post-glacial rebound and takes place on time
scales of a few thousand years. On longer time scales, surface loads are
eventually carried in part by the strength of the lithosphere (regional
support or flexure) and by isostatic mantle pressures. The purpose of
the present study is to compute this eventual lithospheric bending,
which thus reflects the solid Earth response a couple of thousand
years after a change in surface loading.
The governing differential equation is typically (e.g., Zienkiewicz
and Taylor, 1991) derived assuming small vertical deformation of a
thin (Kirchhoff) plate. For linear isotropic elasticity and plane stress
assumptions, a fourth-order differential equation describes the
vertical displacement w that results from flexure in a Cartesian
coordinate system:
!
A2 w
A2 w
A2
D 2 þD 2 þ 2
Ax
Ay
Ay
A2
Ax2
2
2
þðρa −ρi Þgw ¼ q þ Pxx
2
ð1Þ
A2 w
A2 w
A2 w
þ Pyy 2 þ Pxy
AxAy
Ax2
Ay
x and y represent the two horizontal coordinate directions parallel to
the plate, z is the vertical coordinate direction. Poisson's ratio is
represented by ν (0.25), E is Young's modulus (50GPa), ρa is the
density of the asthenosphere (3200kg/m3), ρi the density of the
material that fills the flexural moat, g gravity acceleration (9.8m/s2),
and q represents the surface load. Pxx, Pyy, and Pxz are boundary
tractions, integrated over the thickness T of the plate (T = EET):
Z
Pxx u
T=2
−T=2
Z
σ xx dz;
Pyy u
T=2
−T=2
Z
σ yy dz;
Pxy u2
T=2
−T=2
σ xy dz
and D is the flexural rigidity (or bending stiffness):
Du
GFLEX Fortran source files are available via our website at http://www.
geo.uu.nl/Research/Tectonophysics/.
Plate boundaries constitute important tectonic elements of the
Mediterranean region. Flexure can be substantially affected by the
presence of major faults (Vening Meinesz, 1950), particularly when
these faults are near vertical. Faults are irrelevant for flexure when
they have a near-horizontal dip. As most Mediterranean plate
boundaries are subduction contacts, i.e., gently dipping faults in the
crust, we do not incorporate the plate boundaries in the model.
In the flexure models that we investigate here we do not include
any boundary tractions. We have no doubt that such tractions did
actually act on this tectonically active region, but given the expected
complexity of stress patterns (see Jimenez-Munt et al. (2003) for the
Present) which includes Europe–Africa convergence simultaneous
with back-arc extension and roll-back, we currently feel that adding
such tractions would increase the complexity of the models but not
their accuracy. Anyway, Cloetingh (1986) showed that boundary
tractions have a second-order effect on the flexural response.
Version 1.1 of GFLEX that we used here works with a uniform
restoring pressure (ρa − ρi)gw. Particularly the assumption of uniform
infilling material is inaccurate. We investigated the impact of this
assumption on our results in an iterative procedure where we added
or removed (image) surface loads depending on the flexed topography. We found that the results were nearly identical.
In the model calculations, the equidistant geographical load grid is
first converted to a Cartesian grid with uniform grid spacing dx and dy
(not equal). The finite difference code verifies that the resulting grid
spacing is neither too coarse nor too dense for the given flexural
wavelengths. Finite difference output is subsequently back-converted
to a geographic grid for plotting.
5. Results
5.1. Lower Evaporite deposition
!
A2 w
A2 w
D 2 þD 2 þ
Ax
Ay
!
A DA w A DA w
A2 D A2 w
þðv−1Þ
þ
þ 2ð1−vÞ
2
2
2
2
AxAy AxAy
Ax Ay
Ay Ax
2
113
ET 3
12ð1−m2 Þ
We solve Eq. (1) for zero vertical displacement and displacement
gradient boundary conditions using a newly developed finite
difference package (“GFLEX”). Solutions consists of vertical displacements w, horizontal displacement derivatives ∂w/∂x and ∂w/∂y (used
to compute change in surface slopes), and second derivatives ∂2w/∂x2,
∂2w/∂x∂y and ∂2w/∂y2 (for computing flexure-induced stresses and
changes in erosion/sedimentation). In Appendix A we discuss the
matrix equations and experiments that were used to validate the code.
The first scenarios that we model represent the deposition of all
dominantly marine evaporites. We include in this model the Lower
Evaporites and Salt units from the seismic profiles of the deep westMediterranean basin, collectively labeled as “Lower Evaporite” on Sicily,
and the Mobile unit of Messinian evaporites in the east Mediterranean
(Levantine) basin. We assume that no major sea level fluctuations occurred during evaporite deposition and that the Mediterranean level remained at Atlantic values. We realize, however, that the exact timing of the
onset of evaporite formation in the deep basins is still uncertain, but we
want to emphasize that this will have little influence on our model predictions. The surface load distribution q in this case is determined by the
evaporite that displaces water volume, constituting a net load increase of
ðρs −ρw Þgh
where ρs represents the average density of the Lower Evaporites., ρw
is the average seawater density (taken to be 1024kg/m3), and h is the
thickness of the evaporite layer. As water flows into the flexural moat
next to the load, the density of the infilling material ρi is equal to the
density of seawater. In Sicily, the Lower Evaporites (Decima and Wezel,
1973) consist primarily of gypsum (2200–2600kg/m3; Telford et al.
(1976), limestones (1550–2750kg/m3; Carmichael, 1989), and halites
(2100–2200kg/m3; Carmichael, 1989). In our model, we assume a
single density for the Lower Evaporite deposit of 2250kg/m3.
One of the main outstanding issues in MSC research is the original
thickness of the Messinian deposits. This uncertainty derives from the
inability of seismic and drilling methods to penetrate into the
underlying sedimentary sequences. In most locations, MSC sequences
must be thick and common estimates are between 1500 and 3500m
(e.g., Rouchy and Caruso, 2006) for deep basins. One complication
with these estimates is that it is unclear to which extent thicknesses
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Fig. 2. Two scenarios for the distribution of the evaporites, illustrated for a north–south cross section from the Rhone (left) to Algeria (right); a) Uniform Load scenario: deposition of a
uniform thickness layer, if possible, and b) Deep Load scenario; deposition beneath a critical water depth. c) Lower Evaporite thickness (km) resulting from the uniform load scenario
in which we assume a layer thickness of 1500 m. d) Lower Evaporite thickness (km) resulting from the Deep Load scenario, in which evaporites displace seawater at depths greater
than 1500 m. A significantly greater load results in the eastern (2500 m) than in the western basin (~ 1250 m). e) Investigated evolution of Mediterranean sea level (gray dashed line)
and evaporite thickness (solid black line). We show results of isostasy calculations at two moments (indicated by arrows); towards the end of the deposition of the massive halite, and
after sea level drawdown. See text for further explanation.
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Similarly, the response to a 3km thick uniform evaporite layer is
approximately twice the response of the 1.5km layer that we show
below. The reason why this is only approximately accurate is that the
horizontal distribution of the load changes somewhat in marginal
regions when we change load thickness in both deposition scenarios.
Fig. 2 (continued).
have been affected by tectonics, salt diapirism and gravity sliding, for
which there is also ample evidence.
A related, but slightly dissimilar, issue is the lateral thickness
variation of the evaporites. Available data hint at thinner sequences in
marginal areas than in the deepest part of the basins. Here too, it is
unclear whether thickness estimates reflect the primary sediment
thicknesses, or that marginal instabilities resulted in a major redeposition into deeper parts of the basin. We will come back to this
point after showing below that deeper basinal portions undergo a
more pronounced subsidence than marginal parts.
In view of these uncertainties, we defer to two simple assumptions
for the load (Fig. 2). In our first scenario, we suppose that evaporite
deposition is more like a sedimentary process resulting in blanketing
by a uniform thickness layer. We assume that this layer is 1500m thick
in parts of the basin where, initially, the water column is greater or
equal to 1500m high. The load distribution resulting from this
“Uniform Load” scenario is shown in Fig. 2c. In the second scenario,
we suppose that evaporite deposition is a chemical process that occurs
at sea water depths greater than some critical value (1500m). This
“Deep Load” scenario leads to a variable load thickness, which
depends on the initial bathymetry that we reconstructed in Fig. 1.
The resulting load thicknesses are shown in Fig. 2d.
Besides the spatial distribution of the evaporite load, it is also
important to consider the buildup of the evaporites with time (Fig. 2e)
in relation to the sea level evolution. We think that it is unlikely that a
sea level drawdown occurred much earlier than deposition of most of
the Lower Evaporites, for two reasons; (1) concentrating on salt
budgets only, Krijgsman and Meijer (2008) show that an early
drawdown, while maintaining Atlantic inflow, would have resulted
in unrealistically thick massive halite; (2) concentrating on the
isostatic consequences of a sea level drawdown in the Alboran region,
Govers (submitted for publication) demonstrates that it results in a
complete blocking of Atlantic exchange, thus leaving an unrealistically
thin massive halite if occurring much earlier. Both studies conclude
that a restriction of the oceanic exchange towards the end of the
massive halite is most likely. This is not new; Debenedetti (1982) was
the first to advocate the resulting idea that most of the MSC evaporite
column was deposited during normal sea level conditions.
We recognize that these choices for evaporite density, thickness and
history are not tightly constrained by observations. This introduces a
margin of error and it is important to estimate the influence of these
uncertainties on our model results. This is not very difficult; one
consequence of the linearity of differential Eq. (1) is that, for instance, a
20% smaller density results in a 20% smaller surface deformation.
5.1.1. Uniform evaporite loading
Fig. 3a shows the results of the uniform evaporite scenario of Fig. 2a
and c. Flexure theory indicates that wide loads subside to the depth of
local isostatic equilibrium (e.g., Watts, 2001). The central part of the
western basin indeed subsides to the local isostatic depth of 845m.
Peripheral anti-bulges develop at the fringe of this flat-lying center;
constructive interference in the 3D geometry results in subsidence up to
928m. This result is consistent with flexure theory (see Appendix B),
which predicts that such anti-bulges should develop at a horizontal
distance of 88km from the rim of the load when the horizontal extent of
the load exceeds the flexural wavelength (175km). The maximum
subsidence in the eastern basin south of Greece and Turkey also exceeds
the Airy isostatic value, because the north–south extent of the load
(450km) surpasses the flexural wavelength of the oceanic lithosphere
(260km), resulting in a flexural anti-bulge. West of 20° E, the maximum
subsidence is close the local isostatic value around latitude 35° N. Further
to the NNW, the width of the evaporite load drops below the flexural
wavelength, resulting in less-than-Airy maximum subsidence to the
NNW of 40° N as expected from flexure theory. Flexural bulge uplift of
several decameters is predicted along both the eastern and the western
basin. Cumulative contributions by both basins in the central part of the
Mediterranean product a maximum of 54m in western Sardinia.
Many surface processes result in a critically balanced angle of
repose in soil and river slopes. An “externally” driven change (by deep
processes such as flexure) of these slopes can therefore evoke a
pronounced response, like a reorganization of river drainage systems
or slope instabilities. Fig. 3b displays the flexural slope change, i.e., the
magnitude of the horizontal uplift gradient. Deepening of the
basement mostly increases slope angles of Messinian margins,
particularly in the western basin. Here, a basinward slope increase
by up to 9‰ is predicted. Slope change in the eastern basin is less
pronounced than in the west owing to its thicker EET (Fig. 1b), which
results in longer wavelength bending. The predicted slope angle
changes are probably enough to drive extensive sub-aqueous gravity
flows (e.g., Canals et al., 2004). This would lead to a redistribution of
the evaporite load towards deeper parts of the basin, to amplification
of the flexural gradients and to further gravity flows. To study these
positive feedback loops is beyond the focus of this paper. Qualitatively,
we expect however that current evaporite thicknesses, representing
autochtonous and allochtonous units, are thicker in particularly the
deepest western Mediterranean basin than near the margins as a
result of gravity sliding.
To (very) first order, changes in surface slope curvature are
proportional to changes in erosion rate or sedimentation rate. The
basis for this proportionality is Culling's (1960) model which, when
combined with a mass conservation constraint, results in a linear
diffusion equation of topography. The proportionality constant is the
sediment transport coefficient, and its value ranges over many orders
of magnitude, depending mostly on climate and material type. We do
not believe that a linear description is particularly accurate or useful
for prediction, but we do think that the sign of the flexural curvature
indicates where erosion increases or where sedimentation increases.
One could question whether these model predictions have any
geological relevance, since we cannot distinguish between increase
in erosion and decrease in sedimentation. We actually think they have,
because the geological context of a region will help further
interpretation. For instance, in a sedimentary basin one would be
looking for a decrease in sedimentation (rate).
Fig. 3c displays the change in surface curvature due to flexure. In
our context of regional isostasy we are particularly interested in
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Fig. 3. a) Subsidence (negative, highest subsidence 928 m) and marginal uplift (densely striped regions, ≤54 m) resulting from the deposition of a 1500 m uniform thickness evaporite
layer. Contour lines are shown at every 100 m. The dashed line shows the location of the profile line, for which the flexure is shown in the inset panel at the upper right. b) Slope
change (‰) predicted by the uniform Lower Evaporite scenario. c) Predicted changes in erosion and sedimentation rate due to flexure by a uniform evaporite load. Stippled areas
signify an increase in sedimentation (here equivalent to a decrease in erosion). Darker, hatched regions indicate an increase in erosion (here equivalent to a decrease in
sedimentation). d) Effective stress exceeding 10 MPa. Stress is tensional in regions with plusses, and compressive in horizontally hatched areas.
erosion and sedimentation changes on a basin scale (5km and larger)
on time scales that are longer than the time to develop lithospheric
flexure (few thousand years). The sediment transport coefficient
corresponding to these dimensions is approximately 104m2/yr. Taking
a topography change of 100m (or more) in a million years as a
significant figure, a flexural curvature of 10− 7/m, or larger, suggests a
relevant change in erosion or sedimentation rate. Erosion increase (or
sedimentation decrease) is expected wherever the curvature is more
negative than − 10− 7/m. Sedimentation increases (or erosion decrease)
where curvature is greater than 10− 7/m. We do not make a distinction
between continental and marine conditions. Erosion is predicted to
increase along the basin margins, and sedimentation within the
basins. Whether the predicted rate changes are significant or not
depends on the actual transport coefficient; for example, a value of
104m2s− 1 results in a maximum erosion rate increase by 1mm/yr. A
one order-of-magnitude increase of the transport coefficient makes
1cm/yr.
In upper crustal and sedimentary layers, flexure causes tension
where the basement arches up, and compression where it flexes
down. Failure will occur if these stresses are higher than the rock or
sediment strength. Fig. 3d shows the effective stresses resulting from
the uniform thickness evaporite load. Effective stress σE is the square
root of the second invariant of the deviatoric stress tensor σ′, with a
sign that derives from the first invariant:
qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi
σ E uF 1=2σijV σijV
F ¼ SIGN σijV
(summation convention). As our numerical model is based on an
elastic plate, these stresses represent the potential for deformation;
e.g. high enough tensile stresses may cause normal faulting near the
surface. Whether or not such faulting actually occurs however
depends on the rock strength at any location. The results indicate
that we should expect significant normal faulting along basin margins
in early Messinian and older layers as a result of evaporite loading.
Shortening is predicted in deep portions of the basins. The transition
from tension to compression is relatively sharp. Its location is
controlled the location of the edge of the load.
5.1.2. Deep evaporite loading
Fig. 4 shows the results of the Deep Evaporite scenario for the first
stage of the MSC (Fig. 2b and d). Differences between Figs. 3a and 4a
mostly reflect the significant increase in evaporite load of the eastern
basin in that both the subsidence in, and uplift around, the eastern
basin is higher. Slope changes (Fig. 4b) are consequently also increased
relative to the uniform loading case (Fig. 3b), so that the most
extensive slope instabilities are expected in the eastern rather than in
the western basin.
A very similar picture as previously arises for the change in erosion/
sedimentation rate (Fig. 4c). This relative insensitivity to the magnitude
of the vertical deformation is a consequence of the fact that we chose to
only display sign information, i.e., whether erosion or sedimentation
increases at any location. This sign is mostly dependent on the location
of the evaporite load, and not on its magnitude; the load distribution is
very similar in the two scenarios considered here. We note a similar
insensitivity to the predicted stress regime (Fig. 4d; c.f. Fig. 3d). The
reason for the similarity is the same; whether stresses are tensional or
compressive depends on the horizontal evaporite distribution.
5.2. Late Messinian desiccation
The classical study of Norman and Chase (1986) was the first to
address the solid Earth response to the desiccation of the Mediterranean Sea. In their and subsequent studies (Gargani, 2004a) the
contribution of Lower Evaporites to the basin bathymetry was
ignored. For historical reasons we will follow this approach also,
allowing us to make a comparison with this earlier work. Next, we will
consider the superposition of all Messinian loads to produce what we
consider the best possible prediction for the late Messinian desiccation phase given our current level of knowledge.
Meijer and Krijgsman (2005) showed that evaporation results in
desiccation within a few kyr after closure of seawater gateways. This
result means that the Mediterranean Sea should be considered either
full or empty from the (limited time resolution) perspective of
geological observations. We therefore consider water unloading as an
instantaneous event, which is consistent with our modeling approach
thus far.
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Fig. 4. a) Subsidence (negative, highest subsidence 1510 m) and marginal uplift (densely striped regions, ≤60 m) resulting from the deposition of evaporites beneath 1500 m water
depth (Deep Load scenario, Fig. 2b, d). Contour lines are shown at every 100 m. b) Slope angle changes (‰) due to flexure following loading by Lower Evaporites according to the Deep
Load scenario. The most significant slope instabilities are predicted in the eastern basin. c) Predicted changes in erosion and sedimentation rate due to evaporite loading according to
the Deep Load scenario. Stippled areas signify an increase in sedimentation (here equivalent to a decrease in erosion). Darker, hatched regions indicate an increase in erosion (here
equivalent to a decrease in sedimentation). d) Effective stress in excess of 10 MPa due to the Lower Evaporite loading by the Deep Load scenario. Stress is tensional in regions with
plusses, and compressive in horizontally hatched areas.
5.2.1. Flexure due to removal of the pre-MSC water column
For the purpose of comparing with previous 2D studies, in this
section we examine the results of a 3D calculation in which we
removed the pre-MSC water column. The height of this water column
(Fig. 5) is calculated for the reconstructed basin geometry following
Meijer and Krijgsman (2005). Taking present-day values for the
freshwater budget and assuming that the net evaporation per unit
area of sea surface is constant, we seek the equilibrium level to which
the sea surface will be drawn down upon complete closure of the
strait. Lower sea level entails a reduction in sea-surface area and hence
a decrease in total evaporation. The equilibrium level is reached when
the evaporative water loss exactly equals freshwater input. The
calculation differentiates between the western and eastern parts of
the Mediterranean. For the case of our reconstruction the equilibrium
sea level of the western Mediterranean basin amounts to almost
− 2800m while the eastern basin is predicted to desiccate completely.
The weight of the removed water column is calculated using a
seawater density of 1024kg/m3. Fig. 6 displays the resulting flexural
response of significant uplift in the basins and some decameters of
subsidence along the margins. The inset panel shows the flexural
response along the profile that was also studied by Norman and Chase
(1986) and Gargani (2004a). Vertical deformation in our model is
substantially less than computed by these authors, even if we use the
same flexural rigidity that they do. An implicit, but strong, assumption
in these earlier two-dimensional studies is that the surface load
extends out to infinity in the direction perpendicular to the studied
profile. Even though the water load along the profile is similar to
theirs, the total load is appreciably less in our three-dimensional
calculation due to the curved basin shape.
5.2.2. Response to desiccation during the late Messinian
We now move on to what we consider a more realistic scenario for
the Upper Evaporite phase. Lower Evaporite deposition onto the initial
bathymetry was followed by regional isostatic subsidence of the
scenario for the Lower Evaporites. The eventual water column
therefore is the initial bathymetry (Fig. 1) minus the Lower Evaporite
thickness (Fig. 2) plus the flexural subsidence (Figs. 3 or 4). The
remaining accommodation space is generally less than the pre-MSC
bathymetry. In this section we determine the regional isostatic
response from instantaneous removal of this water column. We
neglect loading by Upper Evaporites, which probably are thin.
Fig. 5b displays the height of the water column (based on the
Uniform Load evaporite thickness) that we assume to have been
evaporated. Like before, the removed water column height was
calculated based on water flux balance. Abyssal parts of the western
Mediterranean lost more than 2100m of equivalent water weight. In the
eastern Mediterranean, a water load close to 3400m disappeared.
The predicted flexure is shown in Fig. 7a. Compared to Fig. 6, the
predicted vertical deformation is less; this is due to the reduction in
height of the evaporated water column relative to that of the pre-MSC (c.f.
Fig. 5b and a). The overall pattern of uplift and subsidence is not so much
affected. The most pronounced uplift is predicted in the central portions
of the eastern basin, where most of the water load has disappeared.
Fig. 7b displays the change in basement gradient. The change is
most pronounced in marginal areas, where a decrease in the existing
margin slope will result, contributing to an increase in slope deposit
stability. Whether the net stability of surface deposits increases or
decreases remains however unclear, because other factors are also
likely to simultaneously affect stability; the release of water pressure
will have a destabilizing effect, and the increase in internal friction due
to water release will tend to stabilize slope deposits.
Desiccation itself may have had a regional imprint on climate (e.g.,
Fauquette et al., 2006). Model predicted changes in erosion and
sedimentation (Fig. 7c) do not include these changes and therefore
should be taken with a grain of salt. In lieu of quantitative assessments
of this climate change, and on its imprint on surface processes, we
consider the results of this graph as one component that may, or may
not, have significantly contributed to surface mass redistribution.
Predicted stress changes (Fig. 7d) are not very susceptible to
climatic changes. Well-dated, regional mapping of tectonic regime in
(otherwise quiescent) marginal areas therefore holds the promise of
independent verification of MSC scenarios; if the desiccation scenario
that we investigate here is correct, the net effect of the unloading
should be that marginal regions show evidence of reverse faulting
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Fig. 5. a) Height of the pre-Messinian seawater column that is removed in the flexure calculation. b) Height of the seawater column (km) that evaporated at the beginning of the late
MSC. This seawater column remained after deposition and isostatic uplift of the Lower Evaporites following the Uniform Load scenario.
(e.g., DeCelles and Cavazza, 1995). The reason for concentrating the
effort on marginal regions is that these are more accessible to direct
observation than deep portions of the Mediterranean basin.
The above Fig. (7a–d) give the expected response to the Zanclean
flooding at the beginning of the Pliocene when taken with a minus
sign.
Our method of estimating the evaporated water load builds on
bathymetry, Lower Evaporite thickness and flexure, all of which have
uncertainties. One potential pitfall of adding various contributions is
that errors add up too, and that (already significant) error bars on the
desiccated water load increase even further. To quantify the sensitivity
of our results to Lower Evaporite thickness, we recomputed in Fig. 7e
the flexure, now assuming that the Lower Evaporites are only
500m thick. This reduces the Lower Evaporite flexure (by one third),
and leaves a taller water column to be desiccated. The figure
demonstrates that the magnitudes change, but not the locations of
uplift and subsidence.
6. Discussion
The results of this paper are based on model calculations of the
flexural down- and upwarping of the lithosphere and crust, i.e.
regional isostasy. The theory underlying such modeling is well
documented (Watts, 2001). Given a surface load, it predicts a rich
variety of possible geological footprints; vertical motions, unconformities, regions where erosion and sedimentation rates change, and
tectonic stress changes. Some of these imprints have been observed,
but others have yet received limited or no attention. Our study is by no
means intended to give a comprehensive explanation of all observations related to the MSC. We intend to estimate the solid earth
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Fig. 6. Uplift (highest uplift 1355 m) and marginal subsidence (negative; densely striped; ≤51 m) resulting from the removal of the pre-Messinian water load (Fig. 5a). Contour lines
are shown at every 100 m. The dashed line shows the location of the profile line, along which the vertical deformation is shown in the inset panel. The dashed line shows the flexure
that was calculated by Norman and Chase (1986) based on a 2D model.
(isostasy) response to MSC events, which will be one contribution only
to the observed phenomena.
Key to ongoing discussions on the MSC is the dissimilar response of
marginal and deep basins — see the recent overview by Rouchy and
Caruso (2006). The physics dictate that adjacent marginal and deep
basins respond synchronously, albeit possibly differently, to MSC
events. Synchronicity is caused by the fact that basement responds
elastically to loading or unloading, even if loading itself is not
simultaneous in peripheral and deep basins. The critical concept here
is the “radius of influence” (R), which defines the horizontal extent of
the region where a load has a marked (say, greater than 10%) imprint.
The influence radius follows directly from flexure theory:
ET 3
R ¼ −lnð0:1Þ
2
3g ð1−v Þðρa −ρi Þ
1= 4
For typical Earth parameters, a water infilling load and effective
elastic thickness T given in km R≌12.4T3/4km, amounting to R = 41km
for a basin with T = 5km and R = 178km for T = 35km. Given the limited
distance between neighbor basins, a simultaneously reaction to a
surface load is thus most likely.
In summary, the expected imprints at basin margins from the
scenarios that we investigate here are
Lower MSC; Loading from deposition of the Lower Evaporites at
current sea level leads to slope instability and thus to gravity flows.
Simultaneously, marginal uplift increases erosion and river
incision, thereby further increasing sediment fluxes towards the
deep basin. The relative scarcity of Lower Evaporite sequences
along the margins, and resedimented evaporites and turbidites in
the deep basins, may be a result of these phenomena. Synchronous
normal faulting of Lower Evaporite and older sediments and rocks
occurs on the margins.
Upper MSC; enhanced erosion of the freshly exposed continental
shelf and slope (sensu Ryan and Cita, 1978). The flexural effect of
unloading is opposite in that it stabilizes margin slope deposits.
Subsidence and riverbed sedimentation occurs on the margins,
uplift closer to the basin center, marginal sedimentation or
decrease in erosion, and reverse faulting of Upper Evaporite and
older sediments and rocks.
Zanclean flooding; destabilization of margin slope deposits, uplift
and river down-cutting, subsidence closer to the basin center,
marginal erosion, and normal faulting of Lower Pliocene and older
sediments and rocks.
This sequence of events likely leads to a stratigraphy like in Fig. 8 in
undisturbed localities. The development, or not, of detrital bodies will
depend on local factors at the margin (sediment starvation, erodability, slope angle and so on), which is why this Figure does not
represent all margins.
Fig. 1 represents our best estimate of the paleogeography,
paleobathymetry and effective elastic thickness during the MSC. In
the section on the paleogeographic reconstruction we indicated that
some parts are fairly well constrained, and others are not. Similar
uncertainties arise in the EET map. To illustrate the sensitivity to the
reconstruction and the assumed EET, we show below results of
identical calculations as presented above, except that we performed
them in the present-day geometry, with the Current EET model of
Pérez-Gussinyé and Watts (2005). These results will be shown below,
when we discuss regional predictions.
The investigated scenarios are not proven or even generally
accepted, and we consider our model predictions as a means to
move towards observational verification. There is some evidence
already, particularly along the better-studied margins in the western
Mediterranean.
6.1. Gulf of Lions
Paleogeographic uncertainties for the Gulf of Lions region are
relatively small, and the size of error bars on modeled vertical
displacements is mostly determined by uncertainties in the Lower
Evaporite load distribution (Fig. 9). Curves a, b and c show responses
to the Lower Evaporite scenarios. Differences between curves a and c
result from variations in the assumed flexural rigidity. Curves d and e
display the computed response to MSC desiccation, which for anyone
interested in the total uplift and subsidence due to the MSC should be
added to b and a, respectively.
The Gulf of Lions continental margin has largely been tectonically
quiet after the Aquitanian (Early Miocene), except for a latest Miocene
to early Pliocene extensional phase (Mauffret et al., 2001). Most of the
faults occur in post-Aquitanian sediments and are sealed by the
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Fig. 7. a) Uplift (highest uplift 1137 m) and subsidence (negative values; densely striped; up to 39 m) due to desiccation. Contour lines are shown at every 100 m. The dashed line
represents the location of the profile, for which the predicted flexure is shown in the upper right panel. b) Slope angle changes (‰) due to flexure after desiccation in the late
Messinian. The expected consequence of these changes mostly is stabilization of surface deposits. c) Predicted changes in erosion and sedimentation rate due flexure following
desiccation. Stippled areas signify an increase in sedimentation (here equivalent to a decrease in erosion). Darker, hatched regions indicate an increase in erosion (here equivalent to a
decrease in sedimentation). d) Effective stress in excess of 10 MPa due to the late Messinian desiccation. Stress is tensional in regions with plusses, and compressive in horizontally
hatched areas. e) Result of sensitivity test to an alternate scenario of desiccation following deposition of a uniform 500 m thick Lower Evaporites. The figure shows the resulting uplift
(highest uplift 1282 m) and subsidence (negative values; densely striped; up to 47 m) Contour lines are shown at every 100 m. The dashed line represents the location of the profile,
for which the predicted flexure is shown in the upper right panel.
Messinian unconformity (Gorini et al., 2005). This deformation phase
is restricted to the western part of the Gulf of Lions (Lofi et al., 2005).
We propose that these tectonics were caused by the loading of the
basin by Lower Evaporite sequences (Figs. 3d and 9).
The Messinian Erosional Surface (MES) has been documented
along several Mediterranean margins (Ryan and Cita, 1978), including
the Gulf of Lions continental margin. Onshore, the MES is recognized
as deep and narrow rivers incisions (Clauzon, 1973). Offshore beneath
the shelf, a prominent seismic reflector with a rugged topography is
interpreted as an erosional surface. Here, post-Messinian deposits lie
unconformably on top of pre-Messinian sediments (Lofi et al., 2003;
Lofi et al., 2005). This reflector is typically referred to as the MES. The
topology of the MES on the shelf represents an extensive erosion
network (Guennoc et al., 2000).
The seismic reflector can be traced from the shelf margin (Fig. 6B in
Lofi et al., 2005) towards the SE into the continental slope region, where
it changes character; it becomes a smooth surface separating conformable Miocene and Plio–Quaternary sediments. Some chaotic seismic
units occur directly above a similar reflector. These units are unique to
the slope. They are conformably overlain by early Pliocene deposits on
the slope, and by Massive Salt and Upper Evaporites along the foot of the
slope. One interpretation of this reflector is that it represents the same
MES (Lofi et al., 2005), which puts the entire erosional phase before the
deposition of the Massive Salt. An alternative interpretation of this
reflector is that it represents a décollement below the Lower Evaporites,
representing gravitational sliding well after the erosional phase that was
recorded further upslope.
Close to the current coastline, seismic and borehole observations of
hard rock incision convincingly demonstrate river activity during the
MSC. Flexural uplift due to Lower Evaporite loading may explain this
observation. Similar evidence for fluvial incision does not exist for the
rest of the offshore erosion network, leaving room to interpret it as
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Fig. 8. Schematic stratigraphy that we expect from our flexure models (inspired by Sage
et al., 2005). Deposition of the Lower Evaporites (LM; includes Massive Halite in our
scenario) causes marginal uplift and erosion, leading to (chaotic?) mass flow deposits
LC. Upper MSC desiccation produces the Upper Evaporites (UM), and simultaneously
exposes the margin to erosion, resulting in erosion/mass flow unit UC. This last erosive
phase entirely removes prior MSC deposits on the uppermost margin; here, the total
erosion is most prominent because the margin had already undergone an erosive phase,
and because Lower MSC tensile faulting primed this part of the margin for enhanced
erosion. PQ stands for Pliocene–Quaternary sediments.
entirely due to river activity or due to submarine mass flows. The
investigated scenario of Lower Evaporite deposition at current sea
level implies a submarine origin of the erosion network. Both the
Fig. 10. Implications of MSC events along a profile through Valencia Trough. The inset in
the upper right shows the location and orientation of the profile line. The shaded relief
map shows a topography swath along the profile (thick horizontal line) in the
reconstructed geometry, and the approximate location of Barcelona city. Thin lines
indicate the azimuth of the meridian and latitude line. Solid lines in the lower panel
show flexural uplift and subsidence for a) the uniform evaporite deposition scenario, b)
the deep evaporite deposition scenario, c) uniform evaporite load in Today geometry, d)
desiccation following deep evaporite scenario, and e) desiccation following uniform
evaporite scenario. The dashed line shows horizontal stress changes from the uniform
evaporite deposition scenario a) (compression negative, tension positive).
shoreline and near shore river incisions and the offshore erosion
network agree with predicted uplift and slope angle changes based on
this scenario for the first phase of the MSC.
6.2. Valencia trough
Fig. 9. Implications of MSC events along a profile through the Gulf of Lions in southeast
France. The inset in the upper left shows the location and orientation of the profile line.
The shaded relief map shows a topography swath along the profile (thick horizontal
line) in the reconstructed geometry, and the approximate location of the town of Agde.
Thin lines indicate the azimuth of the meridian and latitude line. Solid lines in the lower
panel show flexural uplift and subsidence for a) the uniform evaporite deposition
scenario, b) the deep evaporite deposition scenario, c) uniform evaporite load in Today
geometry, d) desiccation following deep evaporite scenario, and e) desiccation
following uniform evaporite scenario. The dashed line shows horizontal stress changes
from the uniform evaporite deposition scenario a) (compression negative, tension
positive).
In the likely absence of rifting tectonics after 8Ma, uncertainties
in the Messinian paleogeography of the Valencia Trough region are
relatively small. However, Alonso et al. (1990) have shown that
post-MSC sedimentation did reshape the NE Iberian margin,
something we did not account for in the present study. Uncertainties in the Lower Evaporite load distribution result in substantial
variations in the predicted subsidence and uplift (Fig. 10). Based on
the synthesis of seismic reflection lines, Maillard et al. (2006)
conclude that the Massive Halite only occur to the NE of the
Valencia Trough, and not in the Trough itself. In keeping with our
classification of the Massive Halites as part of the Lower Evaporites,
this indicates that our Deep Load scenario may be more representative for the MSC than the Uniform Load scenario. The model
predicted differences between these two scenarios are significant;
subsidence (Figs. 3a and 4a) and unconformity angles (Figs. 3b
and 4b) are substantially reduced in the Deep Load scenario relative
to the Uniform Load scenario. The observation from seismic
reflection data that unconformity angles are low or absent therefore
also supports the Deep Load scenario.
The extensive erosive imprint of the Zanclean flooding on Valencia
Trough is well documented (Stampfli and Höcker, 1989; Escutia and
Maldonado, 1992; Maillard et al., 2006). Another seismic reflector was
found beneath the Upper Evaporites by Escutia and Maldonado
(1992). Maillard et al. (2006) trace the reflector beneath the Upper
Evaporites in the central portion of the Valencia Trough and interpret
R. Govers et al. / Tectonophysics 463 (2009) 109–129
125
prominent (Fig. 11), and all major features of Fig. 8 are expected to
be present.
Following the interpretation of Sage et al. (2005) that the C and C’
units represent detrital continental deposits above the Massive Salt
layer, we think that these sediments reflect the erosive phase due to
Lower Evaporite deposition (Figs. 3c and 4c, unit LC in Fig. 8). Our
Lower Evaporite scenario predicts that there should be an
unconformity between Upper and Lower Evaporites (Figs. 3b and
4b) that increases from the upper slope the deep basin. Interpreting
the reflector between C (C’) and T (Upper Evaporite) units as the
onset of the Upper MSC, we note that such unconformity is not
imaged. This discrepancy is caused by an important limitation of our
models, which do not account for the leveling effect of surface
processes.
Onshore, the well documented Messinian normal faulting and
uplift (Casula et al., 2001) has not yet been accurately dated. Our
interpretation is that these represent peripheral uplift due to basin
loading, i.e., the onshore response to Lower Evaporite deposition.
Fig. 11. Implications of MSC events along a profile offshore W-Sardinia. The inset in the
upper left shows the location and orientation of the profile line. The shaded relief map
shows a topography swath along the profile (thick horizontal line) in the reconstructed
geometry, and the approximate location of the town of Oristano. Thin lines indicate the
azimuth of the meridian and latitude line. Solid lines in the lower panel show flexural
uplift and subsidence for a) the uniform evaporite deposition scenario, b) the deep
evaporite deposition scenario, c) uniform evaporite load in Today geometry, d)
desiccation following deep evaporite scenario, and e) desiccation following uniform
evaporite scenario. The dashed line shows horizontal stress changes from the uniform
evaporite deposition scenario a) (compression negative, tension positive).
it as an erosion surface. One alternative to this interpretation is that it
represents a tectonic décollement surface, and that its topology
controlled the location of river gullies during the later desiccation
phase.
6.3. Sardinia
The western Sardinia (and Corsica) margin is one of the best
locations to study the MSC. First, the margin has been tectonically
quiet since the Oligocene. Second, the thickness of Plio–Quaternary sediments is limited relative to other W Mediterranean
margins. Third, the expected imprints along the margin are very
Fig. 12. Implications of MSC events along a profile through western Cyprus. The inset in
the upper right shows the location and orientation of the profile line in the
reconstructed geometry, and the approximate location of the town of Polemi. Thin
lines indicate the azimuth of the meridian and latitude line. Solid lines in the lower
panel show flexural uplift and subsidence for a) the uniform evaporite deposition
scenario, b) the deep evaporite deposition scenario, c) uniform evaporite load in Today
geometry, d) desiccation following deep evaporite scenario, and e) desiccation
following uniform evaporite scenario. The dashed line shows horizontal stress changes
from the uniform evaporite deposition scenario a) (compression negative, tension
positive).
126
R. Govers et al. / Tectonophysics 463 (2009) 109–129
6.4. Cyprus
Lower Evaporites of the Pissouri and Polemi basins consist of
massive lower gypsum, a breccia showing evidence of reworking of
the lower gypsum, and upper gypsum layers interbedded with
carbonates and marls. Upper Evaporites are represented by a
chaotic succession of carbonates, marls and conglomerates. Orszag-Sperber et al. (1989) recognize two major phases of erosion,
chaotic deposits, re-sedimentation and syn-sedimentary deformation in and around Cyprus. The breccia unit of the Lower Evaporites
represents the first phase, which may be coeval to the Massive
Halite (Rouchy et al., 2001). Extension was “more or less
simultaneous” with deposition of this lower unit (Robertson et al.,
1995). The second phase is represented by the Upper Evaporites,
which show ample evidence of paleosols (Rouchy et al., 2001)
related to desiccation. Overall, these observations and inferences
are consistent with flexurally driven erosion and N–S horizontal
extension during the first phase of the MSC (Lower Evaporites;
Fig. 12), followed by erosion due to desiccation.
7. Conclusions
Surface processes of the Messinian Salinity Crisis result in loading
(evaporite deposition) and unloading (desiccation) of the solid Earth.
We construct a new paleo-bathymetric reconstruction of the
Mediterranean. We develop a new finite difference solver to compute
the flexural response in this 3D basin geometry. We demonstrate that
the vertical deformation was substantially overestimated in previous
2D studies.
Vertical deformation of hundreds of meters is predicted, and that
uplift and subsidence are highly variable, albeit synchronous, along
basin margins. Both loading and unloading are expected to cause
significant basement tilting that would be observable as tectonic
unconformities at margins of the Mediterranean basin.
The investigated scenario of Lower Evaporite deposition at current
sea level predicts marginal uplift, unconformities, hill slope instability,
river incision and erosion, and normal faulting. All these phenomena
are observed, in particular the erosion surface.
Desiccation causes significant flexural uplift of the deep basins and
margin subsidence. The most pervasive imprint of this phase is
erosion of freshly exposed margins.
Our study of the tectonic contribution of regional isostasy leads us
to propose that two erosion surfaces of different origins are found in
the Mediterranean (Fig. 8); the first occurs towards the end of the
Lower Evaporites and is caused by flexure, the second is represented
by the Upper Evaporites and is caused by desiccation.
8. Uncited reference
DeCelles and Cavazza, 1995
Acknowledgements
Computational resources for this work were provided by the
Netherlands Research Center for Integrated Solid Earth Science (ISES
3.2.5; High End Scientific Computation Resources). Figures in this
paper were created using GMT version 4.1.2 (Wessel and Smith,
1991).
boundary conditions at nodes i b 1, i > NX, j b 1 and j > NY.
Replacement of differentials in
A2
Ax2
D
!
A2 w
A2 w
A2
þ
D
þ 2
2
2
Ax
Ay
Ay
D
A2 w
A2 w
þD 2
2
Ax
Ay
!
by first order differences in the usual way results in
Di−1;j
þ
Δx4
Di−1;j þ Di;j−1
Di−1;j þ Di;j Di;j þ Di−1;j
Di−1;j þ Di;jþ1
−2wi−1;j
þ
þ
þ wi−1;jþ1
þwi−1;j−1
4
Δx2 Δy2 Δx2 Δy2
Δx
Δx2 Δy2
Di;j−1
Di;j−1 þ Di;j Di;j þ Di;j−1
þ
þwi;j−2
−2wi;j−1
þ
4
Δx2 Δy2
Δy4
Δy
Diþ1;j þ 4Di;j þ Di−1;j Di;jþ1 þ 4Di;j þ Di;j−1
8Di;j
þ
þ 2 2
þwi;j
4
4
Δx
Δx Δy
Δy
Di;jþ1 þ Di;j Di;j þ Di;jþ1
Di;jþ1
−2wi;jþ1
þ
þ wi;jþ2
4
2
2
4
Δy
Δx Δy
Δy
Diþ1;j þ Di;j−1
Diþ1;j þ Di;j Diþ1;j þ Di;j
Diþ1;j þ Di;jþ1
þwiþ1;j−1
−2wiþ1;j
þ
þ wiþ1;jþ1
2
2
4
2
2
2
2
Δx Δy
Δx
Δx Δy
Δx Δy
Diþ1;j
þwiþ2;j
Δx4
wi−2;j
where wi,j and Di,j represent vertical displacement and flexural rigidity
at node i,j. First, second, third and fourth order horizontal derivatives
of the vertical displacement field are thus explicitly defined upon
discretization. Similarly, the part of differential Eq. (1) connected to
lateral variations in the flexural rigidity
!
A2 D A2 w A2 D A2 w
A2 D A2 w
ðm−1Þ
þ 2
þ 2ð1−mÞ
2
2
2
AxAy AxAy
Ax Ay
Ay Ax
results in matrix contributions
ð1−mÞ Diþ1;jþ1 −Di−1;jþ1 −Diþ1;j−1 þ Di−1;j−1
8Δx2 Δy2
ð1−mÞ
8 −Di;jþ1 þ 2Di;j −Di;j−1
þwi−1;j
8Δx2 Δy2
ð1−mÞ −Diþ1;jþ1 þ Di−1;jþ1 þ Diþ1;j−1 −Di−1;j−1
þwi−1;jþ1
8Δx2 Δy2
ð1−mÞ
8 −Diþ1;j þ 2Di;j −Di−1;j
þwi;j−1
8Δx2 Δy2
ð1−mÞ
16 Diþ1;j þ Di−1;j þ Di;jþ1 þ Di;j−1 −4Di;j
þwi;j
8Δx2 Δy2
ð1−mÞ
8 −Diþ1;j þ 2Di;j −Di−1;j
þwi;jþ1
8Δx2 Δy2
ð1−mÞ −Diþ1;jþ1 þ Di−1;jþ1 þ Diþ1;j−1 −Di−1;j−1
þwiþ1;j−1
8Δx2 Δy2
ð1−mÞ
8 −Di;jþ1 þ 2Di;j −Di;j−1
þwiþ1;j
8Δx2 Δy2
ð1−mÞ Diþ1;jþ1 −Di−1;jþ1 −Diþ1;j−1 þ Di−1;j−1
þwiþ1;jþ1
8Δx2 Δy2
wi−1;j−1
Terms due to in-plane traction boundary conditions
Pxx
A2 w
A2 w
A2 w
þ Pyy 2 þ Pxy
2
AxAy
Ax
Ay
result in matrix contributions
Appendix A. Numerical solution of the flexure equation
We solve Eq. (1) using a finite difference approach. Here, we
document the resulting matrix equations. We assume uniform grid
spacing Δx and Δy in horizontal x- and y-direction, so that node
coordinates xi = x0 + (i − 1)Δx and yi = y0(j − 1)Δy. We seek solutions of
the differential equation on a numerical grid 1 ≤ i ≤ NX, 1 ≤ j ≤ NY, with
Pxy
Pxx
Pxy
þw
−wi−1;jþ1
4ΔxΔy i−1;j Δx2 4ΔxΔy
Pyy
Pxx
Pyy
Pyy
þwi;j−1 2 −2wi;j
þ
þ wi;jþ1 2
Δy
Δx2 Δy2
Δy
Pxy
Pxy
Pxx
þ wiþ1;j 2 þ wiþ1;jþ1
−wiþ1;j−1
4ΔxΔy
4ΔxΔy
Δx
wi−1;j−1
R. Govers et al. / Tectonophysics 463 (2009) 109–129
The finite difference equations are assembled into a linear matrix
equation
Aw ¼ B
in which A represents the banded matrix, w the vector of nodal
displacements, and B the load vector. The matrix equation is solved for
nodal displacements either with a direct solver (LAPACK version 3.1,
http://www.netlib.org/lapack) for small problems, or with sparse
solvers (PETSc version 2.3.1, Balay et al., 2001; http://www.mcs.anl.
gov/petsc) for larger problems, typically with a Gram–Schmidt Krylov
subspace approach and LU preconditioner.
The GFLEX software was verified to reproduce two-dimensional
and three-dimensional analytical solutions. Expected limit behavior of
cases with lateral variations in flexural rigidity was also demonstrated.
Appendix B. Flexural anti-bulges
Flexural bulges are a well-known feature in Earth Sciences, for
instance as outer rises near subduction zones. Purpose of this Appendix
is to introduce a similar, yet less known feature of flexure; the anti-bulge.
We will do so by examining a well-established analytical solution to the
two-dimensional flexure equation for a homogeneous, isotropic elastic
plate in the absence of in-plate tractions;
D
d4 w
þ gΔρw ¼ qðxÞ
dx4
The Green function (or line load response)
GðxÞ ¼
ke−kx
ðcoskx þ sinkxÞ
2gΔρ
ðxz0Þ
represents a damped harmonic function with wavenumber k, defined
as
gΔρ 1=4
ku
4D
(Turcotte and Schubert, 2002), which is related to the flexural
wavelength λ = 2π/k.
Fig. B1. Two-dimensional flexure due to box loads of various widths; from left to right,
W/λ = 0.1, W/λ = 5, W/λ = 10.
127
We are interested in the solution to a “box“ load, with width W and
magnitude Q0, that is centered around x = 0;
8
< 0
qðxÞ ¼ Q0
:
0
x a
x a
x a
b−∞; −W=2 >
½−W=2; W=2
bW=2; ∞ >
Watts (2001) shows how convolution of the Green function and
the load gives the desired solution:
8
>
>
<
Q0 −kðx−W=2Þ
coskðx−W=2Þ−e−kðxþW=2Þ coskðx þ W=2Þ
xa½0; W=2
e
2gΔρ
wðxÞ ¼
Q0 >
−kðW=2−xÞ
−kðW=2þxÞ
>
coskðW=2−xÞ−e
coskðW=2 þ xÞ
xabW=2; ∞ >
2−e
:
2gΔρ
The solution is symmetric around x = 0. Fig. B1 shows the ratio of
the vertical deformation w and the local Airy isostatic deformation
wAiry as function of non-dimensional horizontal distance x/λ for
different non-dimensional load widths W/λ. The load is assumed to
act downward, in the direction of gravity acceleration. A narrow box
load of W = 0.1λ results in a response that is very similar to that of a
line load, with a maximum subsidence directly under the load (x = 0)
that is less than the Airy isostatic value. A normal flexural bulge
develops beyond the load, which is a consequence of the mechanical
requirement of subsidence in the center and zero bending of the plate
away from the load.
The response for a very wide load of W = 10λ shows Airy isostatic
subsidence beneath most of the load, except near the load edge (in this
case at x = 5). Here, a flexural anti-bulge develops with a subsidence
that exceeds the local isostatic value. It represents the anti-symmetric
equivalent of the flexural bulge, resulting from a zero deflection
requirement beneath the load center and a deflection gradient near
the edge. Both the bulge and anti-bulge take on their extreme values
at a horizontal distance of λ/2 from the edge of the load. It is clear
from the figure that a flexural anti-bulge should be expected when the
load width is equal or greater than the flexural wavelength of the
elastic plate.
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