Download Evolution of the lithosphere after tectonic emplacement or differential

Evolution of the lithosphere after
tectonic emplacement or differential
stretching
Ylona van Dinther
Department of Earth Sciences
Faculty of Geosciences
Utrecht University
Bachelor thesis supervised by Prof. Dr. M.J.R. Wortel and Dr. R. Govers
Utrecht, 30-08-06, Version 2.0
Contents
1. Introduction ............................................................................
2. Methods
............................................................................
3. Results and Analysis
........................................................
Geotherm evolution
........................................................ 11
Topography
.................................................................. 13
Heat flow .................................................................. 15
Lithosphere thickness ............................................... 17
Moho temperature
........................................................ 19
Melt generation ........................................................ 21
4. Applications…………………………………………………………………………….............
Eastern Mediterranean ............................................... 22
Pannonian Basin ........................................................ 22
Calabria .................................................................. 23
5. Discussion
……………………………………………………………………………..………..
6. Conclusions ............................................................................
Acknowledgements
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References
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Abstract
In the Mediterranean-Carpathian region there are three regions with an abnormal
lithosphere structure, namely Turkey, the Pannonian Basin and Calabria. A
combination of observations in these areas of e.g. high, young topography, high
heat flow, shallow low velocity features together with the occurrence of
volcanism indicate either the presence of a thin lithospheric mantle or the
absence of it. There are no models satisfactorily explaining all observations and
describing the evolution of this lithospheric structure. Two models for the
formation of an abnormal, thin lithospheric mantle are explored; tectonic
emplacement and differential stretching. The first tectonically emplaces the
upper part of the crust directly on top of hot astenosphere and the second is
lithosphere extension with different stretching factors for the crust and mantle.
The thermal evolution of these models is numerically modelled assuming a
conductive heat transfer. From these temperature data the topography, heat
flow, lithosphere thickness and Moho temperature were calculated. Finally, the
possible volcanism and melt generation were inferred by adding solidus curves.
The results are applied to first order observations done in the three regions to
verify the relevance of the models. The tectonic emplacement model predictions
are in good agreement with the observations done in Turkey and the Pannonian
Basin. The unusual high topography of Calabria can not be explained by tectonic
emplacement. All together the modelling results give a first indication that the
tectonic emplacement model is a alternative, promising vision on the
development of thin lithospheric mantle structures.
1
1.
Introduction
The Mediterranean-Carpathian region is part of the final stage of the
convergence between the African and Eurasian plates (Fig. 1). The region
comprises many complicated geological and tectonic areas such as the Pannonian
Basin in Hungary, the Anatolian plateau in Turkey and Calabria in southern Italy.
These areas are subjected to extension, accretion or roll-back and have an
abnormal lithospheric structure. The evolution of these abnormal lithospheric
structures is the focus of this thesis. First the origin of the abnormal structures is
described followed by the specific observations. After mentioning previously
published work two possible tectonic models are suggested.
Figure 1: Plate boundary evolution in the Mediterranean-Carpathian region. The large
black arrows indicate the inferred directions of the lateral migration of slab detachment
and the blue colours indicate bathymetry. From Wortel and Spakman, 2000.
In Turkey the convergence between the African (and Arabian) and European
plate led to the formation of an accretionary wedge of stacked nappes forming an
abnormally thickened continental crust as described by van Hinsbergen et al.
(2005). The region is dominated by a young, high topography (e.g. Sengor&Kid,
1979) and a shallow, hot upper mantle structure with a low shear wave velocity
anomaly (Maggi and Priestly, 2005). Next to these main features an extremely
high heat flow of 107 ± 45 mW/m2 (Metin Ilkisik, 1995) and a long-wavelength
free-air gravity anomaly are observed. Recent volcanism has a lower lithospheric
mantle signature (Maggi and Priestly, 2005). A combination of these
observations can be explained by a thin or absent lithospheric mantle. There are
no models describing the evolution of this type of lithospheric structure.
In the past 10-20 My the Apennine-Calabrian subduction zone rolled back
eastward (see Figure 1). This motion rifted the current Calabria away from
Sardinia and Corsica and opened the Tyrrhenian Sea. After this geodynamic
evolution Calabria is unusually elevated for its crustal thickness and has a normal
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heat flow. Gvirtzmann an Nur (1999 and 2001) explained this by stating that the
crust is floating freely on top of the astenosphere. They are the first and only one
to note a completely absent lithospheric mantle.
In Early Miocene time the subduction of Europe beneath Apulia lead to collision in
the Eastern Alps causing the eastward extrusion of the Carpathian–North
Pannonian lithosphere (ALCAPA block, Csontos, 1995) towards the present day
Pannonian Basin. The lateral movement was accommodated by the roll-back
towards the North-East of the Carpathian Arc. It is generally thought that this
lead to east-west extension and basin formation in from Middle to Late Miocene
time. The observations show however a thin lithosphere of 40-60 km (Bielik et
al., 2004) and a high heat flow of around 90 mW/m2 (Lenkey, 1999).
Alternatively, this abnormal thin lithospheric mantle structure can be explained
by a process other than lithosphere stretching.
In previous work on the response of the lithosphere to stretching by e.g.
McKenzie(1978) and Kusznir et al. (1990,1991,1992) the extension of the
lithosphere and the post-rift thermal subsidence are always modelled with a
uniform stretching factor. These models are still not completely sufficient to
explain all observations done on sedimentary basins. The returning assumption
of a uniform lithosphere extension is an unnecessarily simplification. It is more
realistic to use two different stretching factors for the crust and mantle, because
the partitioning of the extensional strain between the crust and the mantle is not
necessarily equal as they have varying strengths and rheological properties.
Combining the partially unexplained observations with the work done on
stretching so far two models for the origin and evolution of the thin lithospheric
mantle were proposed; tectonic emplacement and differential stretching. The
first type of tectonic emplacement has an initial structure where the upper part
of continental crust is directly emplaced on top of hot astenosphere. The second
type of differential stretching is modelled using different stretching factors for the
crust and lithospheric mantle.
The evolution of the geotherm specific for both situations is numerically modelled
assuming a conductive heat transfer regime. Secondly, these temperatures
dependent upon time and depth are used to model the evolution of the
topography, heat flow, lithosphere thickness and Moho temperature from which
the occurrence of volcanism is inferred. These results are finally applied to a few
first-order observations done in Western Turkey, the Pannonian Basin and
Calabria to show the relevance of the models, but first the general features,
accompanying perturbation of the geotherms and background for both types of
anomalous lithosphere structure are described.
Tectonic emplacement
The first type of tectonic model discussed is tectonic emplacement. Tectonic
emplacement is the tectonic sliding of the upper part of a possibly thickened
continental crust into a hypothetical gap for example created by roll-back of a
subducting slab as illustrated in Figure 2. This leads to a structure were the
continental crust is directly emplaced on top of hot astenosphere. The tectonic
sliding can be accomplished by the extrusion of a region.
3
Figure 2: Schematic illustration of the emplacement of a crust directly on top of hot
astenosphere. Left: Development of this situation. Right: Final stage. The grey shaded
area is lithospheric mantle and the yellow dotted area is astenosphere.
Figure 3 shows the change in geotherm caused by tectonic emplacement. To the
left the original geotherm in a normal lithosphere is shown and in the right figure
the normal upper crustal part is placed on top of a 1300 °C astenosphere. The
perturbation results in an enormous increase in temperatures at a shallow depth.
Figure 3: Evolution of the geotherm for tectonic emplacement. Left: Original steady-state
geotherm. Right: Geotherm after emplacement. The red line is the original geotherm and
the green line is the new, perturbed geotherm. The light grey line represents the Moho,
boundary between the crust and mantle, and the dark grey line represents the
Lithosphere-Astenosphere boundary.
Differential stretching
The second type of process is differential stretching of the crust and lithospheric
mantle. This model is an extension of the standard pure-shear-stretching model
of McKenzie (1978), which deals with homogeneous extension of the entire
lithosphere. In this type two different stretching factors (β) for the crust and
lithospheric mantle are assumed as is seen in e.g. Royden and Keen (1980) and
illustrated in Figure 4.
4
Figure 4: Schematic illustration of differential stretching of the crust and mantle. Left:
Original situation. Right: Final situation. The grey shaded area is lithospheric mantle and
the yellow dotted area is astenosphere. Unfortunately the situation where the crustal
stretching factor is larger than the mantle stretching factor is shown.
Generally the stretching factor of the mantle is larger than the stretching factor
of the crust as is noticed by Davis and Kusznir (2004). This kind of differential
stretching is interesting, because it leads to a thin lithospheric mantle.
The consequences of differential stretching for the geotherm are shown in
Figure 5. Characteristic for instantaneous, differential stretching is a kink in the
geotherm at the Moho, because the steepness of the geotherm differs for
different stretching factors. In situations with a larger stretching in the
lithospheric mantle the increase in temperature gradient is strongest in the
lithospheric mantle.
Figure 5: Evolution of the geotherm for differential stretching. Left: Geotherm before
extension. Right: Geotherm after extension. The red line is the original geotherm and the
green line is the new, perturbed geotherm. The light grey line represents the Moho,
boundary between the crust and mantle, and the dark grey line represents the
Lithosphere-Astenosphere boundary.
A comparison between the two types of model shows that tectonic emplacement
gives a larger increase in geothermal gradient and therefore higher
temperatures. Note that tectonic emplacement is more or less the same as
differential stretching with a stretching factor of 1 for the top of the crust and of
infinity for the mantle. The only difference then is the missing of the lower part
of the crust.
5
2.
Methods
The analysis of the geothermal evolution of the upper 250-km of the Earth is
done in two steps. First, the initial steady state geotherm is adjusted to fit the
two different types of initial temperature distribution leading to a non-steady
state geotherm. Secondly, the evolution of this geotherm is carried out assuming
a conductive heat transfer.
Initial steady-state geotherm
The initial steady-state geotherm is calculated assuming that the main heat
transfer mechanism is conduction, which is heat transfer through a material by
atomic or molecular interaction that occurs through solid material and at
boundary layers. The temperature distribution can be determined by solving the
heat conduction equation (Eq. 1). The units in this equation are all S.I.-units
except for the Temperature T (°C). A conductive heat transfer is appropriate for
the solid lithosphere, but below these depths this assumption is questionable,
because other heat transport mechanisms as convection become more efficient.
A convecting mantle is described by an adiabatic temperature gradient which is
the rate of increase of temperature with depth as a result of the increasing
pressure. An adiabatic regime has no heat transfer, so equation EA does not
hold. Instead a constant temperature profile not dependent upon pressure is
used below the lithosphere. Possibly a temperature increase of 0.5 °C/km could
be added afterwards, but this only leads to a 75 °C increase at a depth of 250
km and is therefore not added.
T
k
A
 2
T 
t
  cp
  cp
(1)
From this equation onwards the first assumptions made are that the upper
modelled part of the Earth is homogeneous and the materials are isotropic.
Isotropic conductivities can be safely assumed, because we are dealing with
polycrystals in the Earth. Secondly it is assumed that heat does not flow
horizontally, so T does no longer depend upon x and y, leading to a one
dimensional depth-profile. In order to model a steady-state geotherm the
temperature is also assumed not to be dependent upon time; Tt  0 . The
Ordinary Differential Equation left is solved numerically, because this allows the
use of more complicated and realistic variables. The temperature profile is now
calculated by iteration in the program Chapman, written by R. Govers after
Chapman (1986).
A thermal definition of the lithosphere is adopted in this study meaning that the
lithosphere is the outer thermal boundary layer of the Earth and its convecting
mantle. The 1300 °C isotherm is chosen as the Lithosphere-Astenosphere
boundary. This is chosen because it falls within the permissible temperature
interval required by many continental lithosphere thermal evolution models and
it bounds the upper mantle temperature fields determined from xenoliths
(Chapman, 1986)
The model extends from the surface to a depth of 250 km in order to capture the
entire lithosphere and be well below skin depth of the temperature variations.
6
This model constitutes of four layers with thicknesses taken from Chapman,1986
(Table 1). The 250 km is covered with a depth increment of 500 m, which is
enough to derive a fluent temperature profile.
The material properties used are stated in Table 1. For these properties it is
assumed that the conductivity is independent of temperature and pressure and a
non-exponential, layered radiogenic heat generation profile following the
distribution of radiogenic isotopes is appropriate. The independence of the
conductivity on temperature and pressure is justified, because the thermal
conductivity is mainly determined by the materials composition. Nowadays it is
common to use the experimental database and determine thermal conductivity
as a function of pressure and temperature directly. This is not necessary in this
case, because temperature and pressure influences are negligible at geologically
relevant pressures and temperatures in the crust (Cull, 1976; Schatz and
Simmons, 1972). The value of the conductivity of the upper crust is partially
determined by the heat flow value at the surface. The heat production of the
upper crust is calculated from the surface heat flow value according to Pollack
and Chapman (1977).
Layer
Depth bottom (km) k (W/m/K)
rho*cp (W/m3/°C)
A (μW/m3)
Upper Crust
17.5
2.429
3.0*106
1.714
Lower Crust
35.0
2.550
3.0*106
0.450
Upper Mantle
200.0
3.200
3.0*106
0.020
Astenosphere
250.0
3.200
3.0*106
0.045
Table 1: Parameters and material properties used in Chapman and Temp2d both after
Chapman (1986) and some are calculated averages by the program Chapman.
The principal constraint used is an observed surface heat flow of 75 mW/m2. This
is a reasonable value for a standard, relatively young lithosphere. The other
boundary conditions used to solve the ODE is a surface temperature of 0 °C. A
constant surface temperature is proper because of convection in the atmosphere.
Perturbation of the steady-state geotherm
In this first step the initial geotherm is determined. In order to test the ideas for
the abnormal lithospheric structure the standard temperature distribution is
adapted to these structures before it is used at input for the second evolution
step. It is assumed that these new temperature distributions are instantaneously
emplaced. This assumption is valuable since tectonic processes are much faster
than heating of the crust by conduction. However, a high temperature difference
that exists at the Moho in our models will lead to faster conduction. The first
tectonic emplacement type is adjusted by instantaneously emplacing the upper
part of the continental crust on top of the astenosphere causing instantaneous
upwelling of the astenosphere leading to a constant 1300 °C from the crust
onwards. For the second differential stretching type it is assumed that the crust
and upper mantle are both instantaneously stretched with a factor of resp. βc and
βm again causing upwelling of the astenosphere. It is also assumed that the
temperature at a depth corresponding to the initial thickness of the crust and
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mantle are fixed as is conform McKenzie (1978). The thickness of the original
structure before stretching are 17.5 km, 17.5 km and 115 km for the upper
crust, lower crust and lithospheric mantle, respectively.
Temperature evolution
The second step is to calculate the evolution of the temperature distribution by
conduction of the in step one created initial geotherm. Therefore the starting
position is the same as in the first step; the thermal conduction equation 1. The
assumption of conductive heat transfer has the same objections as in the
previous step of the model and is solved similarly because the output of
Chapman (1986) is used as input here and temperature at the bottom is
bounded to stay at 1300 °C. Assumptions made to simplify equation 1 are again
those of a homogeneous and isotropic medium and the program Temp2d,
originally written by R. Govers, then solves the isotropic thermal conduction
equation. The finite difference approach and specifically the method-of-lines
technique is used to solve the equation in which temperature is dependent upon
both time and space.
The finite difference solver is used on a rectangular grid with dimensions of 250
km. The layering within the 250 km is the same as those used in Chapman and is
opposed on the grid by changing the material properties at the nodes specifying
the depths specified in Table 1 for each input.
Material properties are assumed to be homogeneous and isotropic within each
layer as stated earlier and the thermal conductivity and volumetric heat
production assumed for each node are the same as those used in Chapman
(1986:see Table 1). The small amount of energy arising from latent heat is
neglected. The density here is assumed to be constant and independent of
temperature and pressure. For diffusion the temperature dependence of density
is not important, because the influence of density on the diffusion rate is minor
at differences in densities experienced in the crust. The specific heat for different
materials and material composition of the crust ranges between 1000-1200
J/kg/K and does not vary more than 20% (Oxburgh, 1980) making it a sound
assumption. The material properties are adjusted to the thickness of the layers
for each initial scenario.
The boundary conditions used to define the geotherm are once more the
constant temperature of 0 °C at the surface and specifically for this step a
constant temperature of 1300 °C at the maximum depth of 250 km. The second
boundary condition results from adopting a thermal definition of the lithosphere.
This implicitly states that the temperature at the base is known and can be used
as a boundary condition. A temperature value as a boundary condition is
preferred above a reduced heat flow value. This namely allows the preservation
of a constant temperature gradient in the astenosphere. Otherwise conduction is
again the dominant heat transfer mechanism till maximum depths, which is not
correct as explained earlier, so in order to realistically build in the convection in
the astenosphere a constant temperature value is chosen. Moreover, a reduced
heat flow value depends more on crustal thickness, which is very variable in this
study. The side boundaries are assumed to be insulating (Q = 0 mW/m2),
thereby excluding horizontal heat flow.
8
Topography
Next to the evolution of the geotherm features that are observable at the Earth’s
surface and dependent upon temperature are modelled in order to make model
predictions. The first observable feature modelled is the topography. This
calculation is based on the principle of isostasy; the pressure at the
compensation depth of 250 km is the same for both the standard reference
column and the column of the abnormal lithospheric structure.
The reference lithosphere has a crustal thickness of 35 km and a temperature
distribution as is formed in the tectonic emplacement situation after 50 My of
cooling. This geotherm is almost the same as the initial steady-state geotherm in
the crust and reaches 1300 °C at a depth of 150 km. According to the definition
used in this study it has a lithospheric thickness of 94 km. In this area this is
appropriate, because surrounding lithospheres are relatively young and cratons
are not near. The topography is measured relative to the sea level of this
reference lithosphere.
Both columns are assumed to have a constant density of the upper and lower
crust (respectively 2600 kg/m3 and 2900 kg/m3). This is a reliable assumption
because the temperature variations induced are small and have an even smaller
compensation effect. These temperature variations are larger and essential below
the crust, so here the density is calculated as a function of temperature for each
node. The relationship between the density and the temperature comes from the
definition of the linear thermal expansion coefficient and is given by the linear
Equation Of State in equation 2.
 (T )   0  (1    T )
(2)
ρ0 is the density of a material at a temperature of 0 °C and α is the volumetric
coefficient of thermal expansion. The values for all quantities are in Table 2. The
linear equation of state is used here, because we are dealing with isotropic
polycrystals that have a density change with temperature through alpha. This
temperature dependence of the density was neglected previously, but the error
resulting from that assumption is only minor for the resulting temperature and
would have led to a circle argumentation where temperature is dependent upon
temperature. For the large temperature differences now it is of importance for
the density.
ρwater
ρUC
ρLC
ρ0,ast
α
= 1000 kg m-3
= 2600 kg m-3
= 2950 kg m-3
= 3330 kg m-3
= 3.28*10-5 °C-1
Table 2: Values of parameters used (mostly taken from McKenzie).
The pressure of the columns is calculated according to the equation 3. The
integration is done numerically according to Simpson’s Rule.
P    g h  g 
z  250000
  ( z, T )  dz
z 0
9
(3)
It assumed that the density at a depth of 250 km is a constant 3188 kg/m3,
which is a sound assumption because temperature is constant at those depths.
This allows one to analytically calculate the elevation from the pressure
difference. Cooling and contraction of the astenosphere can cause the surface
topography to subside below the reference sea level. As this happens it is
assumed that water fills the basin created.
Heatflow, lithosphere thickness, moho-temperature
Additional temperature-dependent, observable features are calculated from the
results. The heat flow Q at the surface is calculated according to equation 4.
Q  kUC 
T
z
(4)
z 0
The lithosphere thickness is defined as the thickness when the temperature
exceeds 90% of the 1300°C isotherm.
10
3.
Results and Analysis
In this chapter the results of the thermal modelling of the geotherm are
presented in the first section. An analysis of these results resulted in the
topography, heat flux, lithosphere thickness and Moho temperature (section 3.2
– 3.5). Finally, an application of these results gives the melt generation in
section 3.6.
Each section consists of a description of the results for first the tectonic
emplacement type and secondly the differential stretching type, which are
represented parallel.
3.1
Geotherm evolution
The modelling of the thermal evolution of the perturbed, non-steady state
geotherms resulting from first tectonic emplacement and secondly differential
stretching results in respectively Figures 6a and 6b at the next page. The
thermal evolution is only shown for one pre-set parameter (resp. 15 km
emplaced continental crust and βC =2.0, βM = 6.0), because a different value of
this parameter only changes the depth of the Moho, Lithosphere-Astenosphere
Boundary and the geothermal gradient at depth, but the general shape of the
geotherm and its evolution are more or less the same.
Tectonic emplacement
The evolution curves in Figure 6a show a heating of the emplaced continental
crust and a cooling of the astenosphere material gradually forming a mantle part
of the lithosphere again. The thermal gradient of the crust is increased during
the first 10 My after which its starts to decrease to more or less the initial
steady-state geotherm. The emplacement of thinner crust upon the astenosphere
however leads to a shallower emplacement of astenosphere material and more
extensive melting of the mantle.
Differential Stretching
Essentially the same astenosphere cooling is seen in Figure 6b during the
evolution of the differential stretching situation. A distinct difference however is
that the crust is not subjected to very extensive heating in the stretching
situation, because a lithospheric mantle is still present above the hot
astenosphere. Another difference is the permanent thinning of the continental
crust, because continental crust does not grow like the lithospheric mantle does.
This thin continental crust leads to a steep gradient for the first 20 My.
Consequently the solidus is crossed earlier and longer leading to more extensive
melting of the crust. Possible melting and volcanism is discussed in section 3.6.
Comments on the geothermal evolution
The final stage of the evolution of the geotherm is an almost linear geotherm
through 0 °C and 1300 °C, because these were the boundary conditions imposed
and no heat flow was admitted to the base. The heat production from radioactive
decay of isotopes, which is concentrated towards the surface, prevents an exact
linear geotherm from occurring. This final linear stage means that a restriction
for the modelling range is required; ages larger than 100 My give no realistic
results.
11
Fig. 6. a) Top; Tectonic emplacement and b) Bottom; Differential stretching.
Geotherm evolution of a 15 km tectonically emplaced geotherm and a differentially
stretched geotherm (βC =2.0, βM = 6.0). The added horizontal grey lines represent:
light = Moho, dark = Lithosphere/Astenosphere Boundary. The purple lines are the
solidus curves of wet (left) and dry (right) lherzolite after Van de Zedde and Wortel
(2001) and the blue lines are the solidus curves for wet and dry (again left and right)
granite (left) and andesite (right).
12
The final linear geotherm is caused by the absence of a heat flux into the
lithosphere. Basal heating of the lithosphere occurs due to secondary convection
in the lower lithosphere and underlying astenosphere and due to heat transfer
from mantle plumes impinging on the base of the lithosphere. Secondary
convection as an explanation for the flattening of the cooling curves for an
oceanic lithosphere was proposed by e.g. Richter and Parsons (1975) and can be
applied to continental lithosphere as well. Secondary scale convection is
associated with instability within the lower part of lithosphere that arises as it
ages and develops a sufficiently low viscosity to develop an internal convective
flow. It is generally accepted that mantle plumes are a source of basal heating,
only the magnitude remains a question (Schubert, Turcotte and Olson, 2001).
These sources of head addition to the base of the lithosphere bring the entire
lithosphere in a thermal state of equilibrium for old continental lithosphere that is
not reached in this model, because these sources are not considered.
3.2
Topography
The results of the evolution of these temperature profiles for the elevation of the
surface are shown in Figures 7a and 7b for tectonic emplacement and differential
stretching respectively.
Tectonic emplacement
The subsidence of the surface from an initial elevation onwards as seen in Figure
7a is primarily explained by the conductive cooling of the astenosphere that
forms the heavier lithospheric mantle as is determined by the thermal definition
of the lithosphere. As the astenosphere material cools down it contracts and this
leads to an increase in density which causes subsidence by the increase in
pressure that has to be compensated. This process is analogue to the subsidence
of oceanic lithosphere. The young part of the curves ( 0-20 My) can also be
described by the relation between elevation and age of the oceanic crust;
elevation is proportional to the inverse square of age in millions of years.
The initial uplift for crustal thicknesses greater than 29 km is caused by the
difference in density between the less dense, emplaced astenosphere and the
more dense lithospheric mantle. In order to remain isostatic equilibrium this is
compensated by the inflow of more astenosphere material leading to an uplift of
the entire continental crust along with its surface relative to the reference sea
level. For smaller thicknesses the influence of the heavier-than-crustal material
in place of the thinned crust counteracts this effect.
Figure 7a also shows a kink to a faster subsidence as the surface comes below
reference sea level and water flows in to decrease the compensation effect of
removing astenosphere and increase the weight pressing on the column. If the
assumption of water filling the basin is not correct and instead it is filled with
heavier sediment the pressure is increased more and the compensation effect is
even smaller leading to a faster subsidising basin. On the other hand can an
absence of water lead to no infill and optimise the compensation effect leading to
a slower subsidence.
13
Fig. 7. a) Top; Tectonic emplacement and b) Bottom; Differential stretching. Topography
as a function of time relative to the sea level of the reference lithosphere as depicted in
Figure 6a. The thicknesses of the continental crust mentioned are tectonically emplaced
directly on top of astenosphere material. Bc an Bm are the stretching factors βCruct and
βMantle. The black line represents uniform stretching of the entire lithosphere with a
factor 2 and the yellow line represents the tectonic emplacement situation for
comparison.
14
The heating of the continental crust was neglected for the calculation of the
topography, because its effect is only minor compared to the temperature effect
of the mantle. This means that the actual initial elevation is a bit larger than
calculated here, because heated crust expands. The final depth is not influenced
by this assumption, because the temperature gradient in the crust returns to
more or less the same gradient within the 100 My.
Differential stretching
The first observation in Figure 7b is that severe stretching of the mantle relative
to a normal crust and the tectonic emplacement situation evolve similarly as one
would expect considering that tectonic emplacement is actually infinitesimal
stretching of the mantle and no stretching in the crust. This is also seen in
Figures 8b, 9b and 10b.
The initial elevation is primarily determined by the amount of stretching in the
crust; more stretching and a thinner crust give more initial subsidence caused by
the replacement of less denser crust with denser astenosphere material. The
subsequent subsidence is for both situations more or less the same as it is
caused by the same cooling of the new mantle part of the lithosphere, which
therefore has the same problems as mentioned earlier.
Comments on topography
A continuous subsidence which does not approach a final depth is observed in
Figures 7a and 7b. This is not realistic, because we observe that subsidence does
not go on forever. This on-going subsidence is explained by the continuous
decrease of the thermal gradient towards a final linear geotherm which is not
correct. The reasons for this erroneous behaviour where explained in section 3.1
and hold in this case as well.
3.3.
Heat flow
The evolution of the heat flow at the surface is calculated from the evolution of
the thermal gradient at the surface. This is shown in Figures 8a and 8b
respectively for tectonic emplacement and differential stretching. The heat flow is
a good indicator because it varies significantly and it can easily be measured at
the surface.
Tectonic emplacement
The difference in heat flow values is completely determined by the thermal
gradient at the surface. This gradient is steepened as the crust is heated from
below by the astenosphere and causes an increase in temperature. Figure 8a
shows that for thinner continental crust more heat reaches the surface and this
maximum is also reached faster as is expected when the distance the heat needs
to travel vertically is shorter. The highest maximum heat flow is almost 120
mW/m2. The initial heat flow of just below 75 mW/m2 is reached again by all
crustal thicknesses at 67 My after which the heat flow slowly decreases further.
67 My is the time it takes for the crustal part of the lithosphere to regain its
original temperature profile. This time is independent of the crustal thicknesses
since it does not vary with this. This time is also more or less confirmed by the
thermal evolution shown in Figure 6a. The slow decrease after this original
crossing point is reached might be similar to the cooling of young continental
lithosphere.
15
Figure 8a: Heat flux as a function of time for the first scenario of tectonic emplacement.
Figure 8b: Heat flux as a function of time for the second scenario of differential
stretching. Bc an Bm are the stretching factors βCruct and βMantle. The black line represents
uniform stretching of the entire lithosphere with a factor 2 and the yellow line
represents the tectonic emplacement situation with 35 km crust for comparison.
16
Differential stretching
The differential stretching type in Figure 8b shows a strong dependence of heat
flow on β for times less than about 30 My. The maximum heat flow in stretching
situations is primarily dependent on the amount of stretching the crust
experiences; more stretching means more thinning so more increase in thermal
gradient, which means a higher maximum heat flow at the surface. Extension
roughly increases the heat flux by a factor β, which is justified by the assumption
of instantaneous stretching compared to heat transfer that leads to an increase
of the thermal gradient with a factor β. Note also that the maxima with a high
mantle stretching factors have a kink. The first maximum of the kink can be
explained by an immediate increase in thermal gradient caused by stretching,
while the second maximum might be caused by heating from the thin lithospheric
mantle which is heated by the astenosphere as it wells up to shallow depths. An
absolute maximum of 140 mW/m2 is reached for stretching with a crust factor of
2.0 as is in agreement with the statement above. After more than 40 My heat
flow is larger for smaller crustal stretching factors, because these crusts remain
thicker and therefore include more radioactive isotopes and produce more heat
by radioactive decay.
Comments on heat flow
An essential difference between the two types of scenario is that the heat flow
for stretching is immediately increased while that for tectonic emplacement
requires some time for the heat to penetrate the crust. The immediate increase
in heat flow follows from the increase of the crustal gradient caused by
instantaneous emplacement as the temperature at a depth corresponding to the
initial thickness stays fixed as was assumed. Both situations reach an almost
asymptotic heat flow value, because the temperature in the upper crust is mainly
determined by the constant heat production in the upper crust.
Compared to a global average heat flow of about 60 mW/m2 these heat flow
values are high, but this is logical as we are concentrating on regions which are
original young and underwent heating from below in these models. From a
comparison with the heat flow values as deduced from the oceanic plate cooling
model it is seen that heat flow values are lower at young ages and higher at ages
until 100 My as is expected from a new, very thin, very hot oceanic crust and an
older crust with lesser radioactive isotopes.
3.4
Lithosphere thickness
The cooling of the hot astenosphere results in the formation of new lithospheric
mantle as analogue to the thickening of oceanic lithosphere, which results in an
increase in lithosphere thickness according to the thermal definition. The
evolution of the lithosphere thickness is shown in Figure 9a and 9b for
respectively tectonic emplacement and differential stretching.
Tectonic emplacement
In the tectonic emplacement type, Figure 9a, one sees that the initial lithosphere
thickness is completely determined by the crustal thickness as expected and that
the thickness increases at a decreasing rate to a common value of about 120 km
independent of the crustal thicknesses. The decreasing rate is similar to the
inverse square relationship
17
Figure 9a: Lithosphere thickness versus time for the first tectonic emplacement type.
Figure 9b: Lithosphere thickness versus time for the differentially stretched type.
Bc an Bm are the stretching factors βCruct and βMantle. The black line represents uniform
stretching of the entire lithosphere and the yellow line represents the tectonic
emplacement situation for comparison.
18
Differential stretching
The differential stretching type in Figure 9b shows that the lithosphere thickness
also increases at a decreasing rate to a common value of around 120 km as in
the tectonic emplacement type. The thicknesses for both types lie within the
same range and show no significant differences.
Comments on Lithosphere thickness
From observations it is known that continental lithosphere does not continue to
thicken with age but instead approaches a thermal equilibrium structure as
mentioned in the thermal evolution section. The results in this section show
again that no equilibrium is approached here as a result of the non-equilibrium of
the temperature distribution, which is again the result of the absence of basal
heating.
The general thickness of the continental lithosphere remains uncertain with
estimates ranging from 100 km to 300 km. According to Pollack and Chapman
(1977) a thermally stabilized lithosphere has a thickness between 100 and 120
km, sometimes 150 km under large and stable cratons. The 120 km resulting
from this study lies within this range and is plausible for relatively young
lithosphere.
3.5
Moho temperature
The temperature at the Moho illustrates whether the bottom of the lower crust
melts and the occurrence of volcanism is expected. The evolution of the
temperature at the boundary between the crust and the mantle is shown on the
next page in Figure 10a for tectonic emplacement and in Figure 10b for
differential stretching. The grey, purple and orange lines plotted in these figures
are specific isotherms that are the melting temperatures of the low temperature
melting rock granite and the lower crustal components granodioritic and tonalite
respectively.
Tectonic emplacement
In the tectonic emplacement type, Figure 10a, an almost immediate raise in
temperature at the Moho is seen after which the temperature decreases at a
decreasing rate the same way for all thicknesses leading to a higher, specific
temperature for thicker crusts as is expected. Melting of the crust resulting from
the crossing of the isotherms is discussed in section 3.6.
Differential stretching
In the differential stretching situation (Fig. 10b) the maximum temperatures at
the Moho are reached about 2 My after t = 0. This delay is caused by the fact
that the heat from the upwelling astenosphere first travels through the
lithospheric mantle. This is confirmed by the later occurrence of the maxima for
smaller mantle stretching factors. The presence of the lithospheric mantle
between the Moho and the astenosphere also leads to a considerable lower
maximum Moho temperature at younger ages for the stretching situation
compared to that of tectonic emplacement. In the end, temperatures are not
affected by this as is seen from the resemblance between the final temperatures
and the temperatures solely determined by the crustal thickness.
19
Fig. 10. a) Top; Tectonic emplacement and b) Bottom; Differential stretching.
Temperature at the Moho through time. Isotherms are added to show the melting
behaviour; grey at 700 °C for granite, dark green at 900 °C for andesite (both after
Chapman, 1986), purple at 1000 °C for granodiorite and orange at 1100 °C for tonalite
(both after Fowler, 1990). Bc an Bm are the stretching factors βCruct and βMantle. The black
line represents uniform stretching of the entire lithosphere with factor 2.0 and the
yellow line represents the tectonic emplacement situation for comparison.
20
3.6
Melt generation
The solidus and isotherm curves plotted in the figures for the geotherm evolution
(6a;6b) and the Moho temperature (10a;10b) demonstrate whether melt can be
generated and possibly be the source of volcanism. This section is meant to
acquire a first order impression of the occurrence of melting and volcanism. It is
beyond the scope of this thesis to precisely predict the melt generation. Three
types of melting can occur in these settings; crustal melting, melting of
lithospheric mantle and astenosphere melting. All three types will be discussed
separately.
1) Crustal melting
Tectonic emplacement
High temperature partial melting of deep continental crust will probably not
occur, because the 1100 °C isotherm, where tonalite liquids are produced and
andesites erupt at the surface, is only reached briefly by an unlikely thick
continental crust. The production of granodioritic melts however is possible for a
short period in small amounts in case water is present to reduce the solidus
(Fowler, 1990). The water needed to reduce this solidus can in a subduction
setting be provided by the upwards-moving volatiles coming from the descending
slab in the near subduction zone. Lower temperature melting can be expected.
The melting temperature of granite (650-850 °C; Chapman, 1986) is reached for
substantial periods of time for all thicknesses. The melting of granitic rocks is
likely for the tectonic emplacement type, because these dominantly upper crustal
rocks become exposed to the astenosphere by emplacement. The melting of
andesite (800-1100 °C; Chapman, 1986) can possibly occur if enough water is
present. From Figure 10a can generally be concluded that thicker continental
crust on top of hot astenosphere melts more, because pressures are higher at
greater depths.
Differential stretching
In case of differential stretching the only type of crustal melting that can occur is
low temperature melting of wet, granitic rocks. The production of granitic melts
can be expected if stretching factors are greater than either 2.0 for the crust or
about 10.0 for the mantle. Part of this result was also found by e.g. McKenzie
and Bickle (1988); extension of the continental lithosphere generates little melt
unless βC≥2.
2) Lithospheric mantle melting
In Figures 6a and 6b the solidus curves of lherzolite, a peridotite composed of
olivine, clino- and orthopyroxene, was added to see if melting of this mantle rock
occurs. In both the tectonic emplacement and differential stretching situations
melting of dry lherzolite can occur during the first few millions of years. The
presence of water can extend this period with tens of millions of years. In
addition to this melting adiabatic melting of the mantle might occur.
3) Astenosphere melting
Finally, melting of the astenosphere might occur by decompression melting as
the astenosphere wells up in response to lithosphere thinning and reaches
shallow levels at which much lower pressure exist. According to the
argumentation of McKenzie and Bickle (1988) mantle material will always melt
when it is brought up to depths shallower than about 40 km unless it occurs in
21
small regions where material moves slowly. Since this is not the case melting of
the astenosphere can easily occur. This melting is also likely to induce extensive
melting of the lower crust (Thompson et al., 1995).
22
4.
Application to the Mediterranean-Carpathian region
In this chapter the relevance of the models is illustrated by applying the results
to some first-order observations done in the Anatolian plateau, the Pannonian
Basin and Calabria. This first order application of the models is brief, but it gives
a first indication if these models may present a possible explanation and have
prospects for further research. It should be understood that these implications
only demonstrate relevance and do not immediately confirm that the suggested
models are correct.
4.1
Eastern Mediterranean; Western and Central Turkey
The thickened continental crust of western and central Turkey caused by the
accretion of island arcs and continental fragments and its hot upper mantle
structure is typically a setting for the tectonic emplacement model.
The model predictions in Figure 7a show that a minimum of 34 km of continental
crust is needed to produce the observed elevation of 500 – 1500 m. This
required crustal thickness is comparable to the thickness of 30 km as determined
by Saunders et al. (1998) from receiver function analysis. Secondary results for
a crustal thickness of 30-35 km give a heat flow of 75-95 mW/m2 (Figure 8a).
These values are within the range of 107±45 mW/m2 observed by Metin Ilkisik
(1995).
The resemblance between the model predictions and the observations is
reasonably good. If the model predictions are correct a subsidence of the
topography is expected.
4.2
Pannonian Basin
The Pannonian Basin is located inside the Carpathian Arc and to the east of the
Eastern Alps. It is generally thought that roll-back of the Carpathian Arc lead to
east-west extension that started 16 Ma. In this generally accepted extension
theory the thin lithosphere is formed by re-growth after large extension of the
lithospheric mantle. A problem in this theory is that it is not clear where the
largely extended lithosphere material is. Another view that avoids this spaceproblem and can possibly also explain the lithospheric structure is tectonic
emplacement. Below observations are compared to the modelling results.
The lithosphere beneath the Pannonian Basin is very thin compared to the
surrounding Carpathians and European platform. It has a crustal thickness of 2530 km and the lithospheric mantle varies from 10-35 km (Bielik et al., 2004).
The region shows a present day high heat flow value of around 90 mW/m2
(Lenkey, 1999). The topography of the Pannonian Basin has an average value of
150 m.
In the tectonic emplacement model a 60 km thick lithosphere for a 25-30 km
crust can be formed after 15-17 My of re-growing lithospheric mantle. Heat flow
values after 15-17 My would be around 95 mW/m2. The tectonic emplacement
model predicts a subsidence of 300 to 1000 m if the basin is filled with water.
Instead the basin was filled with 7 km of sediment which makes it difficult to
predict the original elevation.
23
The lithosphere structure can also be formed by differential stretching of the
lithosphere with a larger stretching factor for the mantle. This type gives similar
results to the tectonic emplacement model. A distinction between the two is
difficult, because lithosphere thickness and heat flow give the same time range
for the event, but differential stretching gives a larger subsidence. This
subsidence can not be directly compared to the real subsidence again because
the basin is filled with sediments. A comparison between the models can be
made by correcting for the sediment infill through back-stripping.
The distinct correlation between the lithosphere thickness, timing and heat flow
shows that the possibility of tectonic emplacement should be considered as a
possible alternative next to the traditional extension theory.
4.3
Calabria
According to Gvirtzmann and Nur (1999 and 2000) the continental crust of the
Calabrian Peninsula is directly underlain by astenosphere. This situation with
absence of a lithospheric mantle can be used to test the tectonic emplacement
model.
The lithosphere thickness beneath Calabria is about 25 km and the Moho is
located at a depth of 20-25 km. Other observations are a heat flow of about 50
mW/ m2 (Loddo and Mongelli, 1978) and a maximum elevation of 2000 m. Ages
of uplifted marine terraces show the uplift of Calabria started 0.7 Ma (Westway,
1993).
According to the tectonic emplacement model predictions of a 20-25 km thick
continental crust a subsidence of 400 – 1000 m is expected around 0.7 Ma. This
is far off from the observed 2000 m elevation. This difference can be explained
by the assumptions of density values for the astenosphere and mantle
lithosphere. By recalculating this initial elevation and using ρA=3200 kg/m3 and
ρML=3250 kg/m3 adopted from Gvirtzman and Nur (1999) an initial uplift of 2400
m is calculated.
The second observation of the relatively low to normal heat flow can also not be
fitted with the modelling results which expect a high heat flow. This misfit is
however directly caused by the chosen boundary conditions. Such a low heat
flow does suggest that no stretching of the crust has occurred and heat did not
penetrate the crust yet as is predicted in the model.
Observations of the Calabrian area do not directly support the tectonic
emplacement model as imposed here, but both misfits are possibly caused by the
choice of input-values, which means the model could be adjusted to this specific
situation.
24
5.
Discussion
The modelling of the thermal evolution and subsequent features of both types of
perturbed geotherm requires the simplification of a complicated reality. In order
to achieve some basic concepts acceptable assumptions have to be made as is
successfully done by for example McKenzie (1978). These assumptions introduce
uncertainties, but these are necessary to acquire a basic idea.
The most profound assumption that is physically not realistic is the assumption
of instantaneous emplacement of the initial geotherm as in the standard model
of McKenzie (1978). In a first approximation this holds well if the extensional or
emplacement phase is short (smaller than 20 My) compared to the duration of
the thermal equilibration or sagging phase. Many sedimentary basins however
show evidence that the duration of the physical stretching was comparable to the
thermal time constant saying that it may be comparable to the thermal cooling
on the same time scale. According to the rule of thumb suggested by Jarvis and
McKenzie (1980) in eq. EZ it was sufficient to assume that stretching was
instantaneous and occurred prior to any thermal equilibration in the discussed
cases of the Mediterranean-Carpathian region.
t
60my
2
, if β < 2
(EZ)
Moreover it seems unrealistic to assume no heat exchange has taken place when
a continental crust of 600 °C is emplaced 500 m. above a hot astenosphere of
1300 °C. In the millions of years it takes the tectonic process, a heat transfer
comparable to that of the difference between the 10 My-geotherm and the 0 Mygeotherm in Figure XB and XC could have taken place. This however does not
change the basic concepts acquired by making this assumption.
Another error could result from applying local isostasy to regions in which this is
not true. Local isostasy is invalid when the lithosphere is influenced by flexure
caused by a large weight placed upon the lithosphere such as a mountain range.
In case of the Pannonian foreland basin in between the Eastern Alps and the
Carpathian range this assumption might not be completely true. Local isostasy is
also not valid in regions that are pulled down by subducting slabs as for example
in Calabria. However it is suggested by Gvirtzmann and Nur (2001) that Calabria
is in isostatic equilibrium because the overriding plate is free from the drag of the
subducting Ionian slab.
25
6.
Conclusions
The modelling of the thermal evolution and topography of the thin lithospheric
mantle structure originating from the tectonic emplacement and differential
stretching models lead to the following similarities and differences in model
predictions.
In general the thermal evolution of the lithospheric mantle is comparable for
both models. The hot mantle cools down at decreasing rate to form lithospheric
mantle. The thermal evolution of the crust differs however. The crust is heated
during the first 10 Mys after tectonic emplacement, while the differentially
stretched crust only cools down.
The subsidence occurs at the same decreasing rate in both models. The
difference is made in the initial elevation. Generally, tectonic emplacement has a
higher initial elevation. Initial uplift occurs for a minimum of 29 km tectonically
emplaced crust. The initial elevation in case of differential stretching is largely
determined by the stretching factor of the crust.
Both models have a high heat flow. The timing of the maxima of the heat flow
however differs; it occurs immediately for differential stretching and after 6-12
Ma for tectonic emplacement.
Cooling of the mantle lithosphere leads to an increase in lithosphere thickness
towards a common thickness of around 120 km at 100 My for all situations. The
re-growing of the lithospheric mantle is similar in both models. The only
difference is the thickness of the crust that has changed for differential
stretching.
Low temperature melting of granite and wet andesite can be expected for
substantial periods of time for all crustal thicknesses emplaced on astenosphere.
In the case of differentially stretched lithosphere partially wet granite will only
melt if βC≥ 2.0 or βM≥10.0. Crustal melting can also be induced by possible
decompression melting of the astenosphere. Melting of the lithospheric mantle
rock lherzolite can also occur for tectonic emplacement and heavily stretched
lithosphere.
The relevance of the models is demonstrated by applying the results to three
regions. The elevation of western and central Turkey can possibly originate from
a minimum of 34 km crust directly on top of astenosphere. An alternative view
on the formation of the 60 km lithosphere beneath the Pannonian Basin is that
15-17 Ma a 25-30 km crust is tectonically emplaced on top of a hot astenosphere
that has subsequently cooled to form the lithospheric mantle. The model
predictions from tectonic emplacement are not in agreement with the
observations in Calabria.
The tectonic emplacement model is an alternative, promising vision on the
development of thin lithospheric mantle structures that deserves further study.
26
Acknowledgements
I would like to express my gratitude to my supervisors Rinus Wortel and Rob Govers, who helped
me fulfil this research in a good way. Rinus is thanked for his useful discussions and suggesting
about the course of my little research. Rob is thanked for his patience during my problems with the
pure programming part. I am grateful for Joop Hoofd´s assistance during my computer problems.
A word of thanks also goes to my roommates in O.207 Mara van Eck van der Sluijs and Hugo for
their support and some discussions. Finally I would like to thank the 61st board of the ‘Utrechtse
Geologen Vereniging’ for their support and supply of food and drinks. The last word is saved for my
parents, Marcel and Jitty van Dinther, who helped me through the final phase by supporting me
and serving me with food and drinks.
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