Download Lithosphere thermal structure and evolution of the

Physics of the Earth and Planetary Interiors 126 (2001) 249–267
Lithosphere thermal structure and evolution of the
Transylvanian Depression — insights from new geothermal
measurements and modelling results
C. Demetrescu a,∗ , S.B. Nielsen b , M. Ene a , D.Z. Şerban a ,
G. Polonic a , M. Andreescu a , A. Pop c , N. Balling b
b
a Institute of Geodynamics, 19-21 J.L. Calderon Street, R 70201 Bucharest, Romania
Department of Earth Sciences, The University of Aarhus, Finlandsgade 8, DK-8200 Aarhus N, Denmark
c ROMGAZ S.A., Mediaş, Romania
Received 20 February 2000; received in revised form 2 February 2001; accepted 1 June 2001
Abstract
The surface heat flow density pattern of the Transylvanian Depression (TD) represents a marked high amplitude short
wavelength low in a region of elevated heat flow. Detailed temperature–depth profiles obtained by continuous temperature
logging, combined with a finite element modelling of topographic and fluid flow effects, support the conclusion that the
observed thermal gradient in the TD truly represents the rate of heat loss of the subsurface. The surface heat flux is 30 mW m−2
in the centre of the depression, increasing to about 60 mW m−2 at margins.
Climatic correction is of the order of 2–10 mW m−2 depending on the investigated depth interval. The measured temperatures
are compatible with the cooling of the Earth surface during the Weichselian glaciation followed by the climate warming which
started 10–11,000 years ago.
The transient effects of Neogene sedimentation and erosion mean an overall 2–3 mW m−2 reduction in surface heat flux. A
low mantle heat flux (30 mW m−2 ) and a low crustal heat production rate in the TD (0.5 ␮W m−3 in the centre) are necessary
to explain the heat flux anomaly.
The surface heat flux pattern of the TD and the extreme low value in the centre of the depression point to an exceptionally cold
and strong lithosphere in the Transylvanian region, as shown by strength envelopes characteristic to the entire depression. The
evolution of the TD generally can be understood in terms of an exceptionally strong lithosphere block caught in a compressive
tectonic environment. © 2001 Elsevier Science B.V. All rights reserved.
Keywords: Heat flow; Climate corrections; Sedimentation effect; Compressional formation of basins; Transylvanian Depression;
Intra-Carpathian basins
1. Introduction
From the geothermal point of view the Transylvanian Depression (TD) sticks out among the intra∗ Corresponding author. Fax: +40-1-2100604.
E-mail address: [email protected] (C. Demetrescu).
Carpathian basins of the same age with its low surface
heat flux, of about 30 mW m−2 in the central part, and
increasing values to about 60 mW m−2 , towards the
edges (Demetrescu, 1978/1979; Demetrescu et al.,
1981; Demetrescu et al., 1991/1992).
The TD is a structural element with molasse sediments of Neogene age, reaching a thickness of 4000 m
0031-9201/01/$ – see front matter © 2001 Elsevier Science B.V. All rights reserved.
PII: S 0 0 3 1 - 9 2 0 1 ( 0 1 ) 0 0 2 5 9 - X
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C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
in the centre of the depression. The TD is superimposed on two older tectonic units, namely a folded
basement and its post-tectonic cover. The term “TD”
refers to the Neogene sedimentary basin while the term
“Transylvanian Basin” includes sediments back to mid
Cretaceous. The location of the TD in the Carpathian–
Pannonian system is shown in Fig. 1. The basement
is a segment of the main Thetyan suture and consists
of metamorphic nappes of inner and median Dacides,
overridden in the west–central area by the ophiolitic
suite which marks the suture (Săndulescu and Visarion, 1978; Săndulescu, 1988). The post-tectonic cover
consists of 4000 m thick upper Cretaceous, Paleogene,
and Eo-Miocene deposits. Of particular interest in this
context is the occurrence of significant amounts of
ophiolites with a low heat production rate.
The TD is a basin created in the compressional setting related to the Carpathian orogeny. No satisfactory
explanation for its formation has been forwarded so
far. Mrazec (1932) considered the TD as an old subsided block, on which sediments have been deposited
in Neogene; Dumitrescu et al. (1962), similarly,
viewed the TD as a result of the subsidence of an old
massif which was not involved in the Alpine orogeny;
Bleahu et al. (1973) suggested a back-arc origin of the
TD, related to the westward subduction in the eastern
Carpathians, while other authors (e.g. Sclater et al.,
1980) assumed a common extensional formation
mechanism for all intra-Carpathian basins belonging
to the Pannonian Basin system s.l including implicitly
the TD. Ionescu et al. (1986) suggested, in accordance with the gravity and magnetic field patterns,
Fig. 1. The study area (rectangle) and the general tectonic framework of the Transylvanian Depression: AB, location of the studied profile;
1, the major Tethyan suture — Transylvanides, Pieniny Klippes; 2, Internal Dacides — a, Tatro-Biharides; b, sub-tatrique nappes; 3,
Median Dacides and Serbo-Macedonian Massif; 4, Outer Dacides; 5, Marginal Dacides (Danubian) and Pre-Balkans; 6, Moldavides; 7,
post-tectonic cover; 8, molassic depressions and foredeep; 9, Neogene volcanites; 10, North Dobrudjan orogen; 11, east European Platform;
12, Scythian Platform; 13, Moesian Platform. Tectonics after Săndulescu (1984).
C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
the intrusion of a large magmatic body at the origin of
the basin. Royden (1988) realised the non-extensional
origin of the TD, suggesting that the subsidence was
temporarily controlled by a dynamic loading from
below during subduction, amplified by the effect of
sedimentary loading. The recent uplift is then a consequence of unloading due to the detachment of the
subducted plate. No quantitative modelling has been
attempted and none of the forwarded models can
explain the low heat flux and the fact that not much
deformation can be seen in the Neogene sediments.
During and after the last stages of thrusting in the
outer Carpathians, the intra-Carpathian basins formed
(Burchfiel and Royden, 1982), namely, the Pannonian
s.s., the Vienna, the Transcarpathian, and the Transylvanian Basins. Generally, the basins are characterised by high heat flow and thin crust and exhibit a
two-phase subsidence history. The first phase consists
in rapid subsidence during Karpathian (17.2–16.4 Ma)
and Badenian (16.4–13 Ma). The shallow water sediments are well localised within distinct fault-bounded
basins, are typically cut by synsedimentary normal
faults and exhibit rotated bedding. A maximum sediment thickness of 4 km is reached in the Vienna and
the Transcarpathian basins. This phase is poorly developed in the Pannonian basin. The second phase is a
slow, long-term subsidence due to cooling and contraction of the lithosphere. It began at the end of Badenian
(13 Ma) and has continued to the present. Sedimentary rocks deposited during this phase are generally
flat-lying and unfaulted. Their horizontal extent is
much greater than that for first-phase sediments and
the sediments onlap onto the pre-Neogene basement.
This pattern of development does not fit the TD,
which shows normal crustal thickness, low heat flow
and almost no signs of extension (Ciupagea et al.,
1970; Royden et al., 1983; Ciulavu and Dinu, 1998).
Furthermore, this basin has undergone recent uplift
and erosion, although there has been little or no folding
and deformation of the basin fill. Surface elevation of
the TD is of 400–600 m, whereas the other basins have
elevations of ∼100 m above sea level.
A joint co-operation between the Institute of Geodynamics (Bucharest) and the Geophysical Laboratory
(The University of Aarhus) on the heat flow and lithosphere evolution of the TD started in 1996 as a contribution to the EUROPROBE-PANCARDI Subproject
“Paleo heat flow and fluid flow in the Transylvanian
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and Pannonian Depression”. It included detailed heat
flow investigations, with new measurements and
assessment of corrections for disturbances of the shallow geothermal field caused by ground water flow,
topography, paleoclimate, sedimentation and erosion,
as well as quantitative modelling of mechanisms of
the formation of such a basin in the compressional
setting of the Carpathians.
Parts of the complex joint investigations have been
presented in detail elsewhere (Şerban et al., 2001;
Andreescu et al., 2001).
In this paper, we shall present the new geothermal measurements and discuss, based on 2-D finite
element models, the topography, ground water flow,
paleoclimate, and sedimentation/erosion effects on
the subsurface temperature field and surface heat flux.
Methods presented by Şerban et al. (2001) and Andreescu et al. (2001) were used in the context of a much
better constrained information on the surface heat flux
pattern and also of refined structural information. We
shall show that of these, only the cooling of the Earth’s
surface during the Weichselian glaciation and the climate warming which started 10–11,000 years ago have
a significant effect on the surface heat flux, depressing
it regionally by 7–8 mW m−2 . Then we analyse the
heat flux budget of the TD and show that the present
pattern of the surface heat flux is mainly the result
of the distribution of heat generation in the crust and
demonstrate that the TD lithosphere is abnormally cold
and mechanically strong. This observation is then related to the thermo-mechanical model of the formation
of the TD presented elsewhere (Nielsen et al., 2001).
2. New geothermal measurements
In 1996 and 1997, new temperature measurements
in 22 boreholes in unsampled areas of the TD, made
by continuous logging, were added to the 12 previous
ones (Demetrescu et al., 1981) made in the 1970s
by discrete logging. The location of all boreholes
is shown in Fig. 2. Geographical co-ordinates and
other details concerning the new measurements are
summarised in Table 1.
The logging system, built at the University of
Aarhus, consists of a quartz thermometer, a winch
driven by an electric engine, a cable counter for
depth measurements, a PC for data acquisition, and
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Fig. 2. Surface heat flux distribution in the TD. Continuously temperature logged borehole of this study (full circle); discrete temperature
logged borehole (Demetrescu et al., 1981, 1991/1992) (cross); identification numbers (Table 1) (1–22); borehole 1 Chiheru (C); borehole
1 Nicolesti (Demetrescu et al., 1991/1992) (N).
an electric power generator. A one-conductor cable
1.5 km long links the probe to the surface parts of the
system. The principle of measurement is counting of
the frequency of a quartz oscillator. The temperature
coefficient is around 1 kHz K−1 and the frequency
of the oscillator is about 9 MHz. The frequency is
divided prior to transmission and then is multiplicated to 28 MHz. The sensitivity of the measurement
is in the mK range. Measurements are taken every
1 s during the thermometer descent with a speed of
3–6 m min−1 . A deconvolution procedure is applied
to account for the continuous movement of the probe
(Conaway and Beck, 1977; Nielsen, 1986). The probe
was calibrated against reference thermometers in a
water bath. The precision of temperature measurements is of ±0.005 K. Old temperature measurements,
taken by the stop and go technique at 20 m intervals,
by means of a thermistor probe with a sensitivity of
about 10 mK, were assigned an overall precision of
±0.05 K (Demetrescu et al., 1981).
All boreholes were in a steady-state thermal regime
achieved after a long (several months, years) period
of rest after drilling (case of discrete logging) or after
the shut down (case of continuous logging). Some of
the wells measured by continuous logging in the TD
are dry in the upper section because they communicate through perforations in the casing with the former
producing formation. The level of the water column
in the well is the piezometric one for the sedimentary
unit opened by the well and does not correspond to
the shallow ground water level; it varies from several
meters to several hundred meters (Table 1).
C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
253
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C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
Continuous temperature logging (Conaway, 1977;
Conaway and Beck, 1977; Beck, 1982; Nielsen and
Demetrescu, 2001) generally gives a detailed temperature profile with better possibility of detection of
disturbing phenomena. In the dry section of the well
the time constant of the temperature probe is large
(several minutes) due to poor heat transfer conditions.
This means that the process of continuous logging
blurs the temperature profile in the dry section of the
well and that, therefore, possible disturbing effects are
not conspicuous. In order to reduce this problem, and
to make the upper several hundred of meters of continuous temperature logs useful, we developed a method
for deconvolution of temperature logs obtained in dry
boreholes. The deconvolution model is a generalisation of the previously used deconvolution model.
Let x and y be the temperature of the air column and
the recorded temperature, respectively. In conventional
theory (Conaway and Beck, 1977; Nielsen, 1986), x
and y is related through the convolution of x with h
according to
1
t
y = h ∗ x,
h = exp −
H (t)
(1)
τ
τ
where h is the probe impulse response, τ the probe
time constant, t the time, H(t) the Heavyside’s step
function, and ‘∗’ denotes convolution.
The differential formulation of Eq. (1) is as follows:
heat transfer between the sensitive tip and the probe
housing and between the sensitive tip and the air, respectively. Factors α and β appear when it is assumed
that the coefficients of surface heat transfer (Carlslaw
and Jaeger, 1959) between liner and air and between
air and probe housing are identical. Here, a and b are
probe radius and liner radius, respectively, ρ 1 c1 and
ρ 2 c2 the heat capacities per unit volume of air and
probe, respectively.
The model of Eq. (3) assumes that the rate of heat
transfer is proportional to temperature differences.
Furthermore, it is assumed that the virgin temperature, x, of the liner or casing is unaffected by the
presence of the probe.
An example for the well 110 Prod (no. 18 in Fig. 2)
is shown in Fig. 3. The observed temperature profile
shows jumps at the logging stops in the air column
and at the air–water interface. It is apparent that the
deconvolution reduces jumps in relation to logging
stops and at the air–water interface.
The 22 new temperature–depth profiles are shown
in Fig. 4. The first 50–100 m of data cannot be used in
deriving thermal gradients because they are disturbed
by the transition from high surface air temperatures
of the summer to the quasi-equilibrium temperature
∂y
1
= − (y − x)
(2)
∂t
τ
which states that the rate of temperature change
experienced by the sensitive tip is proportional to the
temperature difference between the tip and the surroundings. Eq. (2) is now generalised by introducing
the temperature of the probe housing, z, the temperature of the air column, v, and let x be the virgin
temperature of the liner or casing. The relationship
between v, x, y, and z is now stated as
∂v
1
α
∂z
β
= − (v−x) − (v−z),
= − (z − v),
∂t
τ1
τ1
∂t
τ1
1
α
∂y
= − (y − z) − (y − v)
(3)
∂t
τ2
τ3
with α = a/b and β = (ρ1 c1 /ρ2 c2 ) (b2 − a 2 )/ab.
Time constant τ 1 governs the heat transfer between
air column and probe housing and between air column and liner. Time constants τ 2 and τ 3 govern the
Fig. 3. Measured and deconvolved temperature profile for the
borehole 110 Prod.
C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
Fig. 4. Temperature profiles for the boreholes of the present study.
Readings in the dry section of the wells processed by deconvolution
(see text). Identification numbers: 1–22.
in the borehole (see also Fig. 3). Thermal gradients
are in the range of 22–40 mK m−1 , similar to those
previously observed (Demetrescu et al., 1981). Possible disturbances by topography and ground water flow
will be discussed in the next section, and the effects of
paleoclimate changes will be discussed in Section 5.
3. Topography and ground water flow effects
At present the mean topographic surface is at
400–600 m above sea level owing to an inversion
of the basin starting at 5.6 Ma. A network of rivers
created a characteristic topography with elongated
valleys and parallel hill crests. The height difference
between crests and valleys is 150–250 m and the
wavelength of these structures (valley to valley) is
10–30 km. The mean air surface temperature in TD is
about 8◦ C (Tistea et al., 1979). Generally, the ground
surface temperature is 1–2◦ C higher than the air temperature (Blackwell et al., 1980). For this study, a
255
ground temperature of 9.5◦ C, at the valley bottom
and a temperature lapse rate of 6.5 mK m−1 for the
temperature variation with elevation are considered.
Şerban et al. (2001) studied in detail the effects of
topography and topographically driven groundwater
flow on the subsurface temperature in the TD solving
the differential equations of conductive and advective
heat transfer. A 2-D finite element model reproducing
the characteristic wavelength and valley to hill amplitude of 20 km and 220 m, respectively, was used
to reproduce the topography. The depth-dependent
thermal conductivity and diffusivity of sediments was
computed as a function of matrix thermal conductivity, fluid conductivity and porosity (see Section 4).
Based on permeability values (Davis, 1969) for marls,
sandstones and sand, the main sediment in the TD,
the interval 10−14 to 10−17 m2 was investigated. Measurements of sediments of Pannonian age in the upper
50–60 m indicate permeabilities in the same range
(Alexandru Danchiv, 1999; personal communication).
The linear relationship of Bethke (1985) between
permeability and porosity was used to describe the
permeability.
The model calculations show that topography
induces a decrease in heat flux values at the top of
the hill and an increase at the bottom of the valley.
The groundwater flow induced by topography accentuates the topographic effect in the first ∼1000 m of a
borehole, but reduces it at greater depth. The reduced
values of permeability for marls and sandstones generally make the topographically driven groundwater
flow insignificant in TD.
Apart from surficial effects of topography induced
water flow, in sedimentary basins regional scale water
flow at deeper levels could significantly alter the surface heat flux pattern (Powell et al., 1988). However,
in case of TD, in spite of the basin being surrounded
by higher topography, the heat flux pattern, with low
values in the centre, is obviously not compatible with
a regional discharge system. Correcting the surface
heat flux values for the effects of such a system would
enhance the heat flux minimum.
The predicted perturbations of the geothermal field
in the TD caused by topography and fluid flow amount
to maximum differences of about 2–3 mW m−2 between the hill crests and valleys. However, such a
systematic difference between observed temperature
gradients at crests and valleys is not evident in the
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continuously logged boreholes of the TD presented
in this study. The continuously logged boreholes of
the TD were divided into two groups according to the
elevation of the boreholes above the nearest main valley and the average gradient for the two groups was
derived. The difference of 1.31 mK m−1 which was
found is hardly statistically significant. For an average thermal conductivity of 1.5 W m−1 K−1 it would
mean a maximum heat flux correction of 1 mW m−2
for each well, an upward correction for wells at crests
and a downward correction for wells in the valleys.
This correction is well within the usual error in heat
flow determinations, and is therefore, assessed to be
insignificant and has not been applied.
4. Updated heat flow map of the
Transylvanian Depression
The temperature data from the 22 boreholes of this
paper have been processed to obtain the depth variation of the geothermal gradient. Given the sensitivity
of temperature measurements, in the mK range, and
the fact that the boreholes were in a steady-state thermal regime achieved after a long period of rest after the
shut down, we appreciate that the vertical geothermal
gradient was determined within 3% of its undisturbed
value. In Table 1, the mean gradient for each borehole,
computed from values corresponding to 10 m sections
of the temperature logs, and the standard deviation
about the mean are given. The latter is merely a measure of the depth variation of the gradient, caused,
in a steady-state conductive heat transfer regime, by
depth variation of thermal conductivity of rocks.
No cores for conductivity measurements were
available for the boreholes presented in this study.
Generally, very few rock samples are recovered by
the companies drilling the holes, as the sedimentary
structure in the TD has been known for a long time.
For the same reason, only electric logs are routinely
run, which prevented us to apply methods devised to
assess thermal conductivity from well logs (Demongodin et al., 1991; Vasseur et al., 1995). The only
available data on thermal conductivity of rocks from
the TD were published by Demetrescu (1978/1979).
They were measured using a divided bar technique on
41 cores in the depth interval 0–3100 m from various
boreholes drilled in the TD. These values are shown
Fig. 5. Measured thermal conductivity on sedimentary rocks from
TD (Demetrescu, 1978/1979) for the depth interval of 0–1500 m.
The thermal conductivity–depth profile (full line, labelled λ) and
the porosity–depth profile (broken line, labelled Φ) of the present
study are also shown.
as symbols in Fig. 5 for the 0–1500 m depth interval.
We computed the depth-dependent thermal conductivity as a function of matrix thermal conductivity
λm , fluid thermal conductivity λw and porosity Φ,
assuming a porosity decrease with depth according to
an exponential law
Φ
λ = λ1−Φ
m λw
(4)
Φ = Φ0 e−z/ l ,
(5)
where Φ 0 is the porosity at surface and l the porosity
decay length.
The optimum parameters of the depth-dependent
variation of conductivity were chosen as to match
available measured thermal conductivities in TD
shown in Fig. 5. The resulting thermal conductivity–
depth profile, as well as the porosity–depth profile
corresponding to the chosen conductivity–depth profile, are shown in the same figure. The parameters
defining these curves are λm = 2.6 W m−1 K−1 ; λw =
0.6 W m−1 K−1 ; Φ0 = 0.65; l = 1.4 km. The standard deviation of the fit is of ±0.77 W m−1 K−1 . This
C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
rather large value reflects however the limits of our
approach in defining a model for the depth-dependent
conductivity. We are aware of the fact that we can
only characterise the thermal conductivity structure
of the sediments in the depth range of temperature
measurements by mean values. In Section 6, information on the lithology of Neogene sediments along
a line crossing the TD from NW to SE (profile AB in
Fig. 1) is added in a more sophisticated model of thermal conductivity (Andreescu et al., 2001), which was
used to assess the thermal effects of sedimentation.
For the purpose of this section, the simpler model
described above is considered to be appropriate. For
each borehole, the mean value of the thermal conductivity for the investigated depth interval, computed
from values corresponding to 10 m sections of the
conductivity–depth curve of Fig. 5, is listed in Table 1.
The average gradient for each borehole was combined with the average thermal conductivity of rocks
in the same depth range resulting in surface heat flux
values listed in Table 1. Same results were obtained
when deriving the surface heat flux by averaging, for
one borehole, values obtained from different sections
of the hole according to the thermal conductivity and
temperature structure. The 10 m intervals were considered in the calculations. The standard deviation about
the arithmetic mean value is also given in Table 1.
Because the fine structure of thermal conductivity
of sediments is smoothed out in our approach, the vertical variation of the geothermal gradient, illustrated
by the standard deviations of Table 1, is not compensated, in a conductive steady-state heat transfer
regime, by a corresponding variation of conductivity.
Consequently, the heat flux values derived in this
study are affected by rather large uncertainties of
10–30%. It is, however, remarkable that, in spite of
this, the lateral distribution of geothermal gradients
and surface heat flux values is consistent at the basin
scale as illustrated in next paragraph.
The 22 new surface heat flux values, together with
the 12 values previously reported (Demetrescu et al.,
1981, 1991/1992), have been used to construct an
updated surface heat flux map of the TD. Having
in view that the surrounding tectonic units are characterised by much higher surface heat flux values,
namely 80–120 mW m−2 in the east and north, in
the Neogene volcanic area, and 70–80 mW m−2 in
the west, in the Apuseni Mountains and Pannonian
257
Depression (Demetrescu et al., 1991/1992; Mirel
Ene, unpublished data), a value of 60 mW m−2 was
assumed at the depression limits, in order to obtain
a contour map. For the southern Carpathians, in lack
of any determined surface heat flux value and taking
into account the completely different tectonic and
thermal histories of the TD and tectonic units south
of the TD, namely the southern Carpathians and the
Moesian Platform, values of 60–70 mW m−2 , indicated by statistics of heat flow versus tectonic age
(Polyak and Smirnov, 1968; Chapman and Pollack,
1975; Jessop, 1990), were considered appropriate by
Demetrescu and Andreescu (1994) in drawing heat
flow contours in this tectonic unit.
The surface heat flux distribution is presented in
Fig. 2. It shows a finer structure than the previous map
(Demetrescu et al., 1981, 1991/1992), but in general
confirms the heat flux low characteristic to the TD.
As the heat flux minimum characterises the centre of
the depression, where the sedimentary deposits are the
thickest, we analyse the sedimentation/erosion effects
and discuss the surface heat flux pattern in terms of
the heat flux budget and crustal heat production in
Section 6.
5. Climate correction to the derived
surface heat flux
It can be seen from Fig. 4 that almost all temperature
profiles in the TD show a gently increasing gradient
with depth. At shallow depths (0–80 m) the curvature
generally is quite strong owing to the rapid cooling of
the probe from summer temperatures, often in excess
of 30◦ C, to the average shallow subsurface temperature of about 10◦ C. The deconvolution of the temperature logs does not compensate sufficiently for the
shallow strong curvature (Fig. 3), and this part of the
temperature log is hence discarded.
Fig. 6 shows the reduced temperature with respect
to a linear temperature–depth variation for a few boreholes in order to highlight this effect. Also shown are
two temperature profiles obtained by discrete logging
(locations marked by C and N in Fig. 2), taken in
water filled boreholes, demonstrating that the curvature of the profiles is not an artefact produced by the
continuous logging and deconvolution process.
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Fig. 6. Reduced temperature with respect to a linear variation for selected wells (14, 13, 15, 16, N, C). Thick lines represent data and thin
lines climate inversion model results.
The gently increasing gradient occurs in spite of the
fact that thermal conductivity increases with depth,
which should produce a decreasing gradient. Direct
modelling (Şerban et al., 2001) shows that likely
temperature variations associated with the last glaciation can produce a significantly increasing gradient
with depth for the present conductivity–depth function. In view of the above discussion with regards to
water flow and topography we, therefore, believe that
the most likely explanation of the gently increasing
gradient is climatic. In pursuit of this idea the six
temperature profiles of Fig. 6 and other seven continuously logged borehole temperature profile were inverted for the past ground surface temperature history
(GSTH). Least squares inversion in the formulation
of Tarantola and Valette (1982) was applied following the procedure of Nielsen and Beck (1989). As
the depth-dependent thermal conductivity is assumed
to be known, the inverse problem is linear. The variable parameters were assigned rather large confidence
intervals, so they can be modified by the data. The uncertainties in data values cover both observational and
model errors. Data from the top few tens of meters,
where transition from high summer temperature and
the quasi-equilibrium temperatures in the boreholes
takes place, are discarded by assigning very large
uncertainties. For details, see Şerban et al. (2001).
The variable parameters include the ground surface
temperature history and the background heat flux.
Modelled temperature profiles in a reduced temperature representation for the six wells chosen as
examples are shown in Fig. 6. The derived GSTHs for
all 13 studied wells show a tendency towards lower
ground temperatures prior to about 10,000 years.
This indicates that the curvature in the temperature
profiles is consistent with the climatic cooling of the
last glaciation (Weichselian) and the climate warming
following the Weichselian period.
Table 1 lists the heat flux values derived in the process of climatic inversion for the continuously logged
boreholes and for the other two boreholes, logged by
means of the stop and go technique. Heat flow values are corrected by an amount in the range of −4 to
4 mW m−2 . In Fig. 7(a) and (c), these corrections are
plotted against the length of the dry section and, respectively, the investigated depth, in order to assess a
possible influence of these parameters. No trend at all
is visible in these plots. Note also that in six cases of
the 13 boreholes (no. 3, 9, 10, 17, N, C) temperature
data used in inversion were obtained in water. While
C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
Fig. 7. Climate correction as a function of the length of the dry
section of boreholes (a and b) and of the investigated depth (c
and d), in case of individual inversion (left) and joint inversion
(right) of temperature profiles. Numbers: identification number of
the borehole.
plot 7a demonstrates that using for inversion data taken
in the dry section of a borehole is a proper approach,
we could expect that the climate correction depend on
the investigated depth interval, which is not the case
with plot 7c. The scatter in corrections may not be
real and points to the inexpediency of applying inverse
climate analysis to boreholes on an individual basis.
A better regional correction could be obtained
by joint inversion of all boreholes in the area.
This approach effects a stacking of the different
temperature–depth profiles, which significantly improves the ratio of climatic temperature signal to noise
from borehole temperatures and thermal conductivity
depth variations not captured by the model. A joint
inversion of temperatures from nine of the boreholes
of this study, deeper than 700 m and with no obvious
259
disturbances, was performed (Şerban et al., 2001),
obtaining a single GSTH for the area and a better
estimate of the surface heat flux corrected for palaeoclimatic changes. The results in case of a constant a
priori temperature are presented in Fig. 8 and Table 1.
All heat flux values are corrected upwards by an
amount in the range 2–10 mW m−2 . The corrections
were plotted against the length of the dry section and
against the investigated depth in Fig. 7(b) and (d),
respectively. Again, no influence of a possible drawback of the deconvolution procedure for temperatures
taken in the dry section of the well is visible. On
the other hand, the climate correction does show a
dependence on the investigated depth interval, being
larger for shorter boreholes.
Using the regression line of Fig. 7(d) to evaluate
climate corrections for all 22 + 12 boreholes available for the TD results in a climate corrected surface
heat flux map (not shown), very similar to the one
in Fig. 2, but shifted upwards by 7–8 mW m−2 . This
is to be expected for a small area like the TD. The
Transylvanian surface heat flux would still represent a
minimum, in map view, with respect to the surrounding areas, affected by the same climatic changes as
the TD. The only consequence, discussed in Sections
6 and 7, would be the increase of the mantle heat
flux by 7–8 mW m−2 , the increase of temperature,
by variable amounts, in the mantle lithosphere and
the decrease, by 20–30 km, of the thermally defined
lithosphere thickness.
6. The Transylvanian heat flux budget
The surface heat flux pattern reflects the integrated
thermal structure of the lithosphere due to tectonothermal events, heat production in the crust, and heat
conducted through the lithosphere from below
(Pollack and Chapman, 1977). However, in a sedimentary basin, the deposition of cold sediments
reduces the surface heat flux, while heat production in
sediments tends to increase the heat flux. The effects
of uplift and erosion are the opposite.
In this section, we examine the heat flux budget
of the Transylvanian Basin, investigating the thermal
effects of sedimentation and erosion on the thermal
structure history, as well as the contribution of the
crustal heat production and mantle heat flux to the
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Fig. 8. Joint inversion results for a priori constant temperature. GSTH, SD ratio, misfit for each borehole (T) and histogram of data
coverage are shown. Numbers in the rightmost plot: identification number of the borehole.
observed surface heat flux. Our study is restricted to a
WNW–ESE geologic profile 240 km long crossing the
depression in its central part (Fig. 1). A 2-D model of
the basin thermal evolution, which takes into account
both the lithosphere structure and the time changes
in sedimentation rate and in the lithology of the sediments, is solved numerically using the finite element
method. The basin evolution is carried out in discrete
time steps, in which the material is added or eliminated with a constant rate for a certain sedimentary,
respectively erosion period (for details on the numerical technique, see Andreescu et al. (2001)).
The present-day thermal structure of the Transylvanian crust in a 1-D approach has been analysed in
previous papers (Demetrescu et al., 1981; Visarion
and Veliciu, 1981; Demetrescu, 1982; Crânganu and
Deming, 1996).
The general structure of the crust and basin along
the cross-section is presented in Fig. 9. The data of
Visarion et al. (1973), Săndulescu and Visarion (1978),
and Ionescu et al. (1986), based on borehole data and
reflection seismic lines, were used to obtain an image
for the upper part of the crust and that of Rădulescu
et al. (1976), Cornea et al. (1978); Rădulescu (1979),
based on refraction studies, for the Conrad and Moho
discontinuities. The sediment structure and composi-
tion is taken from regional studies (Ciupagea et al.,
1970; Ştefănescu et al., 1985) and available borehole
information.
The structure of the crust reveals a depth to the
Moho discontinuity of 29 km in the centre of the basin,
where a thinning of the upper crust to about 6–8 km is
also clearly distinguished. The crust becomes thicker
at the ends of the profile (40–45 km in the east, towards the eastern Carpathians, and 32 km towards the
western margin, under the Apuseni Mountains). Of
particular importance for our discussion is the Transylvanian nappe, obducted over the inner and median
Dacides in the western and central parts of the basin
and the thin (2–6 km) metamorphic part of the upper
crust overlain by ophiolites.
The Neogene sedimentary formations, beginning
with the Hida molasse (Karpatian age), are deposited
over the post-tectonic cover. Around 16.4 Ma ago,
the basin underwent a short (0.2 Ma) salinity stage,
when salt was deposited over volcanic tuff layers. In
Middle Miocene (16.2–11 Ma ago), the basin experienced a rapid subsidence (about 4 km in 5 Ma), and
the deposited rocks were mostly sandstones, marls,
and marly clays mixed with sands. The last 5.6 Ma of
the evolution represents an uplift stage with creation
of erosional valleys.
C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
261
Fig. 9. Crustal cross-section along the profile AB of Fig. 1. Upper panel: heat flux variation along the profile; data point along profile
(full circle); surface heat flux from map isolines of Fig. 2 (open circle).
In calculating the thermal effects of sedimentation
and erosion along the profile, a special care has been
taken in accounting for the lateral variation of thermal
properties of sediments. Volume fractions of the main
sediment types deposited during the formation of the
Transylvanian Basin (90–5.6 Ma), obtained from reference wells located on the profile, were used to define
the surface porosity, decay length of porosity, density,
heat capacity, thermal conductivity, and heat generation rate. A smooth variation of the material parameter
values within each individual formation was invoked
by linearly interpolating the parameter values given
for the selected wells.
Modelling results show that the surface heat flux
evolution between 90 Ma and present is characterised
by pronounced variations in Neogene, as it can be
seen in Fig. 10. The evolution of the surface heat
flux is shown for two points on the profile, where the
sedimentation and erosion rates are those illustrated
in the upper panel of Fig. 10. For the point located
at 100 km on the profile, a maximum reduction of
about 10 mW m−2 is reached during the Badenian–
Sarmatian times (16.2–11 Ma), when the sedimentation rate is maximum. However, the thermal relaxation
following this rapid sedimentation episode played an
important role in diminishing this heat flux deficit by
about 7–8 mW m−2 . At present, the overall effect of
the Neogene evolution of the basin in the surface heat
flux is of only about 2.5 mW m−2 . The basin is still in
a transient thermal state, but the surface heat flux variation is very low (∼0.4 mW m−2 per Ma). The lateral
variation of the Neogene subsidence and uplift effect
is of only about 2 mW m−2 between the middle of the
TD (4000 m of sediments) and the margins (500 m).
It cannot explain the observed surface heat flux variation, of about 30 mW m−2 , from centre to margins.
The main factor controlling the lateral variation of
the surface heat flux in the TD is the crustal heat production rate. A very simple model with depleted and
thinned upper crust in the middle of the basin can satisfactorily reproduce the present day surface heat flux
distribution if included in the 2-D sedimentation model
described above. This model shows a good agreement
with observations (Fig. 11). We assigned a low heat
production rate of 0.5 ␮W m−3 to the upper crust beneath the central part of the basin, corresponding to
the presence of ophiolites and thinned metamorphic
inner and median Dacides and a more normal upper
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Fig. 10. Surface heat flux evolution at the basin centre for two locations on the profile AB of Fig. 1 (lower panel). Sedimentation/erosion
rate (upper panel).
Fig. 11. Upper panel: model of lateral variation of heat generation rate: B, basement; C, upper/lower crust limit; M, Moho. Sediments: H = 0.8 ␮W m−3 . Upper crust: intensely hatched areas, H = 2.0 ␮W m−3 ; light hatched area, H = 0.5 ␮W m−3 . Lower crust:
H = 0.1 ␮W m−3 . Lower panel: surface heat flux in case of constant mantle heat flux (30 mW m−2 ), lateral variation of heat generation
rate of the upper panel, and sedimentation. Surface heat flux data (full circle); surface heat flux from map isolines of Fig. 2 (open circle).
C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
crustal heat production rate of 2 ␮W m−3 to the upper
crust of the basin margins. Sediments have an average
heat production rate of about 0.8 ␮W m−3 , and the
lower crust a heat production rate of 0.1 ␮W m−3 . For
the mantle heat flux, a value of 30 mW m−2 was required by the model. Various combinations of mantle
heat flux and crustal heat production rate are possible,
but an increase in either of the values will have to be
compensated by a decrease of the other. A laterally
variable mantle heat flux cannot be the origin of the
short wavelength high amplitude surface heat flow
anomaly: it produces only a uniform shift of the heat
flux (Andreescu et al., 2001). The upper crustal contribution to the heat flux increases by 10–18 mW m−2
from the centre to the margins.
It is obvious that using in modelling climatically
corrected surface heat flux values instead of measured
ones would not change any of the results presented
in this section, because, in such a small area, the correction would be uniform, as shown in Section 5. All
values, including the mantle heat flux and the initial
surface heat flux of the model, will be shifted upwards
by 7–8 mW m−2 , but the lateral variations and the
relative time variations resulting from modelling will
remain unchanged. We prefer the formulation based
on determined values rather than on corrected ones,
having in view, on one hand, uncertainties associated
with calculating such a correction (Beck, 1977; Powell
et al., 1988), and, on the other hand, the facts that past
climate changes would have influenced a much larger
area than the TD, for which, however, no information
regarding the palaeoclimate effect is available, and that
currently available surface heat flux maps which include the TD (e.g. the heat flow map of Europe (Hurtig
et al., 1991/1992)) are based on measured values.
The differences of 5–15 mW m−2 between measured and calculated heat flux values at the ends of
the profile can be explained by the influence of high
Pannonian heat flux in the Apuseni Mountains at the
western end and, respectively, by the presence of the
Neogene igneous masses at the eastern end. They are
not taken into account in our modelling because of
the no horizontal heat flux condition at the model vertical margins. The relatively low mantle heat flux of
30 mW m−2 can be related to the heat flux depression
produced in front of the subduction known to having
taken place at the eastern limit of the TD, in the eastern Carpathians (Demetrescu and Andreescu, 1994).
263
7. The compressional formation of the
Transylvanian Depression
Our detailed heat flow investigations reported in this
paper demonstrated that TD is abnormally cold, with
a surface heat flux in the centre of about 30 mW m−2 .
Corrections for the effects of rapid Neogene sedimentation and paleoclimate may increase this value
to about 38 mW m−2 , which is still low for a geodynamically active continental region. The tomographic
model of Fan et al. (1998) also indicates a colder
lithosphere in the TD than in surrounding areas.
The corresponding modelled thermal structure
of the lithosphere along the profile AB of Fig. 1
is presented in Fig. 12. It shows a cold Moho of
350–450◦ C, and a cold lithosphere (∼850◦ C at a
depth of 100 km), with somewhat lower temperatures
in the centre as compared to margins. Based on this
thermal structure and on mechanical properties of
the three lithospheric layers of the model (granite,
diabase/basalt and olivine, respectively) we estimated
the strength envelopes along the profile, according
to principles and material properties of Ranalli and
Murphy (1987) and Ranalli (1991). The results are
shown in Fig. 12 for four locations along the profile (40, 85, 110, 200 km) and also for two locations
in the adjacent tectonic units west and east of the
TD (Apuseni Mountains and the Neogene volcanic
chain of the eastern Carpathians, respectively). The
strength was limited to 600 MPa (Bassi, 1995). The
well known limitations of this kind of modelling (e.g.
Ranalli, 1991; Cloetingh and Burov, 1996; Fernandez
and Ranalli, 1997), imposed by large uncertainties
on the values of rheological parameters extrapolated from laboratory experiments, uncertainties in
the lithosphere temperature field, and by adopting
a simplified vertical structure of lithosphere, make
our calculations rather schematic. But they, however,
demonstrate in a first approximation that the entire
basin lithosphere is mechanically strong as compared
to the surroundings characterised by much higher
surface heat flux. This is also supported by the fact
that generally, except at basin margins, the Neogene
sediments are not disturbed by the compressive events
affecting the areas surrounding the TD (intense thrusting and strike-slip movements and mountain building in the eastern and southern Carpathians). Using
corrected surface heat flux values in modelling, the
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C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
Fig. 12. Thermal and mechanical structure of the lithosphere along the profile AB. For comparison, strength envelopes for adjacent tectonic
units are shown. Isotherms (thin line); crustal structure (thick broken line) (B, basement; C, upper/lower crustal limit; M, Moho); strength
envelope (thick line).
thermal and mechanical structure of the lithosphere,
instead of measured ones, would not produce significant changes in the above discussed image of the lithosphere strength in the TD and adjacent tectonic units.
The TD subsided and was filled with molasse
sediments derived from erosion of thrust-induced
topography in the Carpathians during three compressional phases identified in the surroundings, namely
the old Styrian (20–17 Ma), young Styrian (16–15 Ma)
and Moldavian (13–11.3 Ma). Structural data from
the surroundings of the TD, supporting the existence
of the compressional regime, have been recently published (Huismans et al., 1997; Ciulavu and Dinu,
1998). Ciulavu and Dinu (1998) present a detailed
review of the literature on the Transylvanian Basin
and compile existing and new structural evidence.
They interpret the Transylvanian Basin as a “squeezed
block” in Pliocene between the south Carpathians and
the northern part of the east Carpathians. Only gentle
compressional deformation of the Neogene sediments
is seen in the form of folding related to salt tectonics.
Polonic (1996) also determined weak (4–6%) extensional episodes at base Senonian and base Badenian.
The strong lithospheric block surrounded by weaker
lithosphere in the eastern and southern Carpathians
and in the Apuseni Mountains, evolving in a compressional setting and having lateral dimensions comparable to its thickness, called for a new model of
compressional basin formation presented elsewhere
(Nielsen et al., 2001). According to this model, when
a relatively strong lithospheric block exists amidst
lithosphere of more normal strength, compressional
forces act on the lithosphere causing permanent shortening and thickening of the relatively weak surroundings. The requirement of the continuity of the upper
mantle strong layer results in dragging down the
upper mantle of the strong zone, which in turn creates
space for sediments at the surface. A stress relaxation
results in basin inversion.
8. Conclusions
New geothermal measurements were carried out by
a continuous logging technique to depths between 400
and 1400 m in 22 boreholes in the TD. These data add
to 12 previous surface heat flux determinations and
confirm the heat flow low which is characteristic to
the TD. A new, more detailed surface heat flux map
of the TD has been derived. It shows surface heat flux
values of about 30 mW m−2 in the centre, increasing
to 60 mW m−2 toward the margins.
High quality temperature data, acquired with a
borehole thermometer sensitive in the mK range, and
geothermal gradients well within 3% of their natural value are reported. Models of the average depth
distribution of thermal conductivity in the TD, based
on measured conductivities on cores from several
C. Demetrescu et al. / Physics of the Earth and Planetary Interiors 126 (2001) 249–267
boreholes in the TD were derived. Local variations
cannot be caught by these models and, consequently,
the standard deviation associated to the new surface
heat flux data ranges between ±10 and ±30%.
The new temperature measurements were taken
in former production wells in a steady-state thermal
regime achieved after a long period of rest after the
shut down. Some of the wells measured by continuous
logging were dry in the upper several hundred meters
because they communicate through perforations in the
casing with the former producing formation, a piezometric level of the water column being established.
To reduce the problem of the large time constant of
the temperature probe in the dry section of boreholes,
and to make the upper several hundred of meters
of temperature log useful, a method for deconvolution of temperature logs has been developed. A few
stops in the air column and the temperature jump at
the air/water interface are used to derive corrected
temperatures by means of an inverse analysis of the
transient temperature records during the stops.
Finite element modelling and temperature data
analysis show that the topography and topographically driven ground water flow corrections are not
significant. The climate correction is of the order of
2–10 mW m−2 , depending on the investigated depth
range. Individual and joint inversions of temperature
profiles from sets of boreholes show that measured
temperatures are compatible with the cooling of
the Earth surface during the Weichselian glaciation
and with the climate warming beginning 10–11,000
years ago.
Modelling the evolution of the basin in the interval
90–0 Ma shows pronounced variations of the surface
heat flux in Neogene, with a maximum reduction
of about 10 mW m−2 reached during Badenian and
Sarmatian (16.2–11 Ma). The subsequent thermal relaxation and the uplift of the basin 5.6 Ma ago tend
to restore the pre-Neogene thermal structure, so, at
present, the overall effect of the Neogene evolution of the basin is a reduction of heat flux of only
2.5 mW m−2 , as compared to the steady-state value.
The main factor controlling the lateral variation
of the surface heat flux in the TD is the crustal heat
production rate. A very simple model with depleted
and thinned upper crust in the middle of the basin
can satisfactorily reproduce the present day surface
heat flow distribution. A low heat production rate
265
of 0.5 ␮W m−3 has been assigned to the upper crust
beneath the central part of the basin, corresponding
to the presence of ophiolites and thinned inner and
median Dacides overlain by ophiolites, and a more
normal value, of 2 ␮W m−3 , to the upper crust of the
basin margins. Also, a relatively low mantle heat flux
of 30 mW m−2 is necessary to explain the general low
values of the surface heat flux in the TD as compared
to the surrounding areas.
The surface heat flux pattern of the TD and the extremely low value in the centre of the depression point
to an exceptionally cold and strong lithosphere in the
Transylvanian region, as shown by strength envelopes
characteristic to the entire depression.
The evolution of the TD generally can be understood in terms of an exceptionally strong lithospheric
block caught in a compressional tectonic environment.
Acknowledgements
The paper is part of the Subproject 12 “Geothermal
field and fluid flow in the Pannonian and Transylvanian
Basins” of EUROPROBE-PANCARDI ESF research
programme. This study was supported by the Danish
Natural Science Research Council Grant no. 9601354
and by the projects IG-02.3/1996, 1997 of the Institute
of Geodynamics, Bucharest. Partial support for presenting the paper at the IUGG 1999 General Assembly
was granted to CD from the IGCP Project no. 428. The
Romanian National Agency of Mineral Resources and
ROMGAZ S.A. are thanked for providing access to the
wells. Niels Breiner and Kenneth Madsen are thanked
for help with continuous temperature logging and
Venera Dobricã for assistance with drawing figures.
Suggestions for improving the manuscript, by Jean
Claude Mareschal, Ilmo Kukkonen, and an anonymous reviewer, are gratefully acknowledged.
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