[Pasquale Et Al. 2012]_Heat Flow in the Western Po Basin and Surrounding Orogenic Belts

Geophys. J. Int. (2012) doi: 10.1111/j.1365-246X.2012.05486.x

GJI

Geo

dyna

mic

san

dte

cton

ics

Heat flow in the Western Po Basin and the surrounding orogenic belts

V. Pasquale, P. Chiozzi, M. Verdoya and G. GolaDipartimento per lo Studio del Territorio e delle sue Risorse, Settore di Geofisica, Universita di Genova, Viale Benedetto XV 5, I−16132 Genova, Italy.E-mail: [email protected]

Accepted 2012 March 28. Received 2012 March 28; in original form 2011 June 22

S U M M A R YOn the basis of lithostratigraphic data from petroleum wells, geophysical logs and laboratorymeasurements, we revised surface heat-flow values from previous studies and obtained newdata for the western sector of the Po Basin (Italy). The available bottom-hole temperatures werecorrected for mud circulation. The in situ thermal conductivity was estimated by taking intoaccount the combined effects of mineral composition, anisotropy, temperature and porosity.The radiogenic heat of the basin rocks was evaluated from natural γ -ray logs and with labo-ratory γ -ray spectrometry. After correction for sedimentation and palaeoclimate influences,the inferred surface heat flow of the basin is 73 ± 4 mW m−2 in the undeformed Po foredeepand decreases to 64 ± 6 mW m−2 in the Alps and Apennines buried units, affected by activethrust. Numerical and analytical calculations indicate that such a decrease is likely caused bythrusting. Local heat flow minima (<60 mW m−2) may be related to groundwater movement.These data together with those available for the surrounding orogenic belts allow redrawingthe heat-flow pattern of northwestern Italy.

Key words: Downhole methods; Heat flow; Sedimentary basin processes; Heat generationand transport.

1 I N T RO D U C T I O N

The Po Basin is a foreland basin laying between the Alps and Apen-nines orogenic belts. It has formed since the Palaeocene within ageneral tectonic context of convergence. After the region experi-enced Mesozoic extensional events, which led to the formation of awide carbonatic platform, the tectonic regime turned into compres-sive. This caused the formation of south-verging thrusts during theNeo-Alpine phase in the Southern Alps and north-verging thrustssince the Oligocene in the northern Apennines (Hill & Hayward1988). The basin can be thus considered as a result of continen-tal lithosphere flexure in response to the load increase caused bystacking of the Alps and Apennines thrust sheets (Ingersoll & Busby1995). The Mesozoic carbonatic platform is at present deeply buriedbeneath the clastic sediments deposited during Tertiary times (Pieri& Groppi 1981).

The western sector of the Po Basin, investigated in this paper,is shown in Fig. 1. The principal structural characteristics are wellknown from seismic studies and drillings carried out for hydrocar-bon exploration and from results of international scientific projects,such as EGT (see Blundell et al. 1992) and ECORS-CROP (seeNicolas et al. 1990; Roure et al. 1990). The following tectonic unitscan be recognized.

(1) Southern Alps Buried unit (SAB), related to the post-collisional (Oligo-Miocene) deformation of the Alps belt, and con-sisting of thrust sheets, laying beneath a Plio-Quaternary cover,

whose regional detachment surface corresponds to the top of theEocene deposits.

(2) Undeformed Po Foredeep unit (UPF), filled by terrigenoussediments deposited since Upper Eocene–Oligocene times and char-acterized by an impressive contribution of dismantling debris of theAlps and, only in recent times, of the Apennines.

(3) Northwestern Apennines Buried unit (NAB), consisting ofthrusts bonded by arcuate external fronts, developed west since theOligocene and east since the Messinian, and underlying the Plio-Quaternary cover.

The foregoing structural arrangement makes difficult the study ofthe thermal evolution of the basin. To this regard, the knowledge ofthe surface heat flow can give a basic contribution. Heat flow is theonly observable geophysical quantity related to the thermal effectsof the regional tectonics. This parameter is particularly difficult todeal with, because it may include both transient components relatedto the tectonic history and geological noise originated by a varietyof shallow and deep processes, which can be of different origin (e.g.Pasquale et al. 2010).

The only papers tackling the surface heat-flow pattern in the basinare those by Pasquale et al. (1986a,b) and Pasquale & Verdoya(1990). These papers were based on temperatures from petroleumexploration wells and thermal conductivity determinations of coresamples recovered from deep drillings. Heat-flow data (more than100) were incorporated in the database by Cermak et al. (1992).Fig. 1 shows their location. The inferred heat flow was generally

C© 2012 The Authors 1Geophysical Journal International C© 2012 RAS

Geophysical Journal International

2 V. Pasquale et al.

Figure 1. Location of heat-flow sites from previous studies in the Western Po Basin (see text). Sites whose temperature data were reprocessed are indicatedwith the code number of Table 1. NAB, Northwestern Apennines Buried unit; SAB, Southern Alps Buried unit; UPF, Undeformed Po Foredeep unit. Fullsquares are wells at which GR logs were available.

low, being on average about 50 mW m−2. Attempts to account forsuch a low value with the thermal effects of the tectonic evolutionwere made by Pasquale et al. (1993), but a thorough analysis to testthe quality of thermal parameters and the data used for the heat flowinference was never carried out.

In this paper, we use recently implemented techniques to revisethe data presented in previous works and to process new temperaturedata. Attention is paid to anisotropy effects, porosity variation withdepth and the temperature-dependence of thermal conductivity, andthe heat flow is calculated with the approach based on the thermalresistance. The thermal regime of the basin stems from the super-position of different heat-transport mechanisms and geodynamicprocesses, which have changed the structure and consequently thephysical properties. Thus, besides sedimentation, palaeoclimaticand thrust faulting effects, radiogenic heat and heat refraction due tothermal conductivity are analysed. Finally, we use the new heat-flowdata together with those available for the orogenic belts surround-ing the basin to depict the surface heat-flow pattern of northwesternItaly.

2 T H E R M A L DATA

Several hundred of geophysical well-log headers, available at theItalian Economic Development Ministry (Mining National Officeof Hydrocarbons and Georesources), report on temperature data ofthe basin. To improve the data quality, we selected only wells pre-senting at least two bottom-hole temperatures (BHT) and detailedinformation on lithostratigraphy. 39 wells were chosen; some ofthem (seventeen) were already analysed in previous studies (Fig. 1and Table 1). As a whole, 194 temperature data were available,

including 82 BHTs with their relative shut-in time te, that is, thetime elapsed between cessation of the mud circulation and the BHTmeasurement, and 29 time-series of BHTs measured at different te

at the same depth. Moreover, nine temperatures measured duringdrill stem tests (DST) were available. The latter data correspond tothe reservoir fluid temperature, supposed to be in equilibrium withthe surrounding rocks (formation temperature).

The estimation of surface heat flow from raw temperatures andthermal properties involves some basic processing. While drilling,the mud circulating in the well perturbs the formation temperature.This causes the deepest parts of the well to get cooled, while theascending mud carries some of the heat from the deep formationsand increases the temperature in the upper section of the well. Thus,with the exception of DST information, raw data must be processedto infer the formation temperature. To this regard, the duration ofmud circulation, the temperature difference between the formationand the mud, the nature of the heat exchange processes that occurduring drilling, the possible loss of circulation, the well radius,the thermal properties of the mud–rock system and the drillingtechnology used play an important role. Another major problem isthe estimation of the thermal properties at in situ conditions. Thefollowing sections describe how we inferred thermal properties andcorrected temperature data for mud circulation.

2.1 Thermal properties

Pasquale et al. (2011) implemented a technique, based on labora-tory measurements and mineralogical analyses, for estimating thein situ thermal properties from the rock mineral composition or,alternatively, from lithostratigraphic data available from drilling

C© 2012 The Authors, GJI

Geophysical Journal International C© 2012 RAS

Heat flow in the Western Po Basin 3

Tab

le1.

Pre

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p

Wel

lcod

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me

(N)

(E)

(m)

(m)

(WK

−1m

−1)

(mW

m−2

)(m

Wm

−2)

(WK

−1m

−1)

(μW

m−3

)(m

Wm

−2)

(mW

m−2

)(m

Wm

−2)

(mW

m−2

)

1-G

erol

a45

◦46

’09

◦27

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937

542.

2955

572.

870.

5360

6263

632-

Lis

anza

45◦

45’

09◦

37’

338

3282

––

–2.

610.

9654

5656

563-

Cas

cina

Riv

iero

45◦

40’

09◦

54’

248

3784

––

–2.

360.

8058

6060

604-

Fran

ciac

orta

45◦

36’

09◦

60’

244

3329

––

–2.

660.

6269

7172

725-

Cas

tano

45◦

32’

08◦

46’

175

6728

––

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181.

1256

5966

666-

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ari

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32’

09◦

57’

142

6840

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590.

7359

6368

687-

Turb

igo

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32’

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46’

167

6630

1.97

5259

2.24

1.05

5760

6868

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071

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30’

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50’

116

7267

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350.

9353

5764

6410

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45◦

30’

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34’

106

6470

2.11

4956

2.72

0.91

6060

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45◦

29’

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44’

152

6226

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510.

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261.

1155

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1.69

3237

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5253

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22’

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03’

107

5009

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471.

1457

6075

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01’

104

5538

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421.

0857

6077

7717

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ana

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19’

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19’

146

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19’

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4452

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5960

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rana

45◦

06’

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40’

101

5705

1.85

4252

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3841

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25–

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1.63

2649

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45’

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260

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reports. These authors also report on density and porosity of rocksamples of the basin. We applied their approach and used their datato infer thermal conductivity and volumetric heat capacity at eachinvestigated well.

The in situ thermal conductivity kin was obtained with the geo-metric mixing model

kin = k(1−φ)m kφ

w, (1)

Figure 2. Vertical thermal conductivity kin versus depth of Belvedere (a) and Sali Vercellese (b) wells as inferred from lithostratigraphic information. H,Holocene; E–PC, Eocene–Palaeocene.




where km and kw are the matrix and water thermal conductivity,respectively. Porosity φ was assumed to decrease with depth z as

φ = φo exp (−bz) , (2)

where b is the compaction factor and φo is the surface porosity.By expressing depth in kilometres, values adopted for φo and b are0.180 and 0.396 km−1 in carbonate rocks, 0.298 and 0.461 km−1

in marls and silty marls, 0.284 and 0.216 km−1 in sandstones andcalcarenites, and 0.293 and 0.379 km−1 in shales and siltstones, re-spectively. Carbonate rocks, marls and sandstones were consideredas isotropic, whereas thermal anisotropy of the clay-rich lithologies(siltstones and shales) was taken into account. In anisotropic rocks,the vertical matrix conductivity, which decreases with depth due

Figure 3. Cooper & Jones slope of 29 time-series of BHTs measured atdifferent shut-in times as function of depth. The best-fitting curve is givenby eq. (12).

to the orientation of the clay and mica platelets during burial, wasestimated by using the relation

km = 2.899 − 0.251z. (3)

The water thermal conductivity kw was assumed to change withtemperature as suggested by Deming & Chapman (1988), whereasthe temperature dependence of the matrix conductivity was evalu-ated with the expression by Sekiguchi (1984). The total uncertaintyon thermal conductivity, which takes into account the errors incorrection for anisotropy, temperature and porosity, is 10 per cent(Pasquale et al. 2011).

Fig. 2 shows as an example the stratigraphic column and thethermal conductivity profile modelled for two wells. The verticalthermal conductivity was calculated at the middle-point of 20 mintervals. In the uppermost kilometres, the compaction effect islarger than that due to the temperature and, for the same lithotype,this causes an increase of conductivity with depth. Both wells showthat the maximum values of conductivity occur in mudstones andsandstones. Horizons of silty shales are present at different depthsand exhibit minima of conductivity. In these horizons, due to thepresence of thermally anisotropic sheet silicates, conductivity isconstant or decreases with depth.

The in situ volumetric heat capacity (ρc)in was computed as theweighted average of the volumetric heat capacity of the matrix (ρc)m

and the volumetric heat capacity of water (ρc)w in the voids

(ρc)in = (1 − φ) (ρc)m + φ (ρc)w . (4)

The specific heat of water as a function of temperature varies ac-cording to the relation by Somerton (1992). As long as the pressure ishigh enough to keep the water in a liquid phase, the volumetric heatcapacity of water under subsurface (high pressure) conditions wasestimated with good accuracy without including pressure depen-dence. Since the thermal cubic expansion coefficient is very smallfor rocks, density was considered as constant over the temperaturerange expected within the sedimentary basin, so that the volumetricheat capacity of the matrix increases in accordance with the riseof the specific heat as a function of temperature. The temperaturedependence of specific heat for any mineral matrix was computedby means of the equation by Hantschel & Kauerauf (2009).

Figure 4. Temperature correction (T∞ − BHT) versus depth for different shut-in times (te). rb is the well radius, tc is the mud circulation time.




2.2 Formation temperature

Depending on the information available in the well reports, weapplied different kinds of correction to reduce BHT data. In wellswhere temperature time-series at a given depth were available, weused the method by Cooper & Jones (1959) that assumes a long holeof small diameter has been drilled quickly and filled with a fluidcooler than the formation. The temperature of the mud approachesthat of the formation as heat flows radially inwards from the wallsof the well. For any BHT measurement, one has

BHT = Tm + (T∞ − Tm) [1 − F (α, τ )] , (5)

where T∞ is the formation temperature and Tm is the temperature ofthe drilling mud. The function F(α,τ ) can be expressed as (Bullard1947)

F (α, τ ) = 4α

π 2

∫ ∞

0

exp(−τu2

)u� (u)

du, (6)

with �(u) = [u J0 (u) − α J1 (u)]2 + [uY0 (u) − αY1 (u)]2, Jn andYn are Bessel function of order n of the first and second kind,respectively, α = 2 (ρc)in/(ρc)f , τ = (kin te)/[(ρc)in r 2

b ], te is shut-intime and rb is well radius. The drilling mud is a mixture of water

and clay that produces a colloidal suspension (Leblanc et al. 1981;Bear 1988), so its volumetric heat capacity (ρc)f was assumed tovary linearly as function of the amount of clay. Then, (ρc)f can beinferred from data on the mud density ρf

(ρc)f = (ρc)w + (ρf − ρw) cc, (7)

where ρw and cw are the density and the specific heat of water and cc

is the specific heat of clay. For the few wells whose rb was unknown,we applied the relation proposed by Pasquale et al. (2008) for thePo Basin

T∞ = BHT + (18.9z − 2.7z2

)ln

(1 + tc

te

), (8)

where tc is the circulating mud time and z is the depth. For te <10 hran additional temperature correction of 2 ◦C is necessary. The timetc can be estimated by means of the relation

tc = 1.7 + 0.05z + 0.10z2, (9)

where time is expressed in hours and depth in kilometres.When temperature time-series were not available, a technique

that enables to correct a single BHT was applied. By introducing

Figure 5. BHTs after mud correction and DST temperatures versus depth. The least-squares regression curve is shown.




the Bullard method (Bullard 1947; Funnell et al. 1996) in the tem-perature recovery one obtains (Zschocke 2005)

T∞ = BHT + Q

4πkin

[E1

(r 2

b

4κinte

)− E1

(r 2

b

4κin (tc + te)

)], (10)

where E1 is the exponential integral. Q is the heat supplied dur-ing thermal relaxation of the well, which is given by (Kutasov1999)

Q = 2πkin (T∞ − Tm) qD (tD) , (11)

Figure 6. Temperature versus thermal resistance (R) for wells of Table 1. Eq. (14) is shown for each well.




where qD(tD) = 1/ ln(1 + D

√tD

), D = π/2 + 1/(

√tD + b), b =

2/(2√

π − π ), tD = κintc/r 2b and κ in is the rock thermal diffusivity.

The term (T∞ − Tm), which controls the radial heat flow, wasestimated from the relation

T∞ − Tm = −3.0z2 + 34.6z − 30.6, (12)

deduced from slope data of eq. (5) versus depth z, expressed inkilometres (Fig. 3). Fig. 4 shows T∞ − BHT as function of depthfor different values of tc and rb. For the average shut-in time of theanalysed data set (about 10 hr), a maximum correction of about10 ◦C occurs at 5.5 km depth.

BHTs after correction for mud effect and DST temperatures ofthe selected wells (Table 1) are plotted as a function of depth inFig. 5. The largest temperature (188 ◦C) was found at 6670 m depth,whereas the maximum measurement depth is 7250 m. Despite therigorous treatment of thermal data, the uncertainty of inferred for-mation temperature may be still relatively large (say 3 ◦C). However,the large depth of measurement smoothes the bias on the averagethermal gradients, which can be estimated with relatively high pre-cision (Deming 1994). By assuming a mean ground surface tem-

perature of 12.5 ◦C, the least-squares fit to all the data set yields anaverage geothermal gradient of 23.6 mK m−1.

3 H E AT F L OW

The combination of the temperature data and the thermal conduc-tivity estimated from the lithostratigraphic columns of each wellallows the calculation of the surface heat flow. The classical ap-proach of the thermal resistance method was applied. The thermalresistance R along the vertical between the surface and the depth dis

R = zd∑

z=0

(1

kin

), (13)

where kin is the estimated in situ vertical thermal conductivity atany depth interval z = 20 m. The subsurface temperature in ahorizontally layered, isotropic medium is related to the thermalresistance as

Td = To + qo R, (14)

Figure 6. (Continued.)




Figure 6. (Continued.)

where Td is the temperature at depth z = d, To is the ground surfacetemperature and qo is the surface heat flow. To was assumed todecrease with elevation at a rate of 6 mK m−1.

Fig. 6 presents the formation temperatures against R for eachwell. The slope of the linear fit to data gives the surface heat flow.The obtained heat-flow values (Table 1) do not take account ofradiogenic heat, sedimentation and climatic change effects, andthus further corrections must be applied.

Eq. (14) is based on the assumption that the radiogenic heat rateH does not affect the temperature–depth distribution. In terms ofheat flow, simple calculations show that the error introduced bythis assumption is −H L/2, that is the product of the well half-depth L by the average radiogenic heat rate (Rybach & Bodmer1983). We evaluated H by means of laboratory measurements andγ -ray logs available from eight wells (Fig. 1). In laboratory, it wasdetermined from uranium, thorium and potassium concentrationsmeasured with a γ -ray spectrometer on rock samples recoveredfrom some wells (see Chiozzi et al. 2002 for details on methodand uncertainty). After correction for well diameter, drilling muddensity and logging tool eccentricity, γ -ray logs allow estimating H

in μW m−3 by means of the relationship (Bucher & Rybach 1996):

H = 0.0158 (GR − 0.8) , (15)

where GR is the log reading in API units. Eq. (15) holds for GRvalues lower than 350 API units and gives H within an acceptableerror (<10 per cent).

Fig. 7 depicts the GR logs and the radiogenic heat rate profilesfor two example wells, whereas the average values of H of the mainlithotypes of the basin are shown in Table 2. Since no sample ofschist and orthogneiss, forming the basin basement, was available,their radiogenic heat was taken from the literature (Pasquale et al.2001). The weighted average of the radiogenic heat rate H of eachwell is listed in Table 1. The overall average radiogenic heat rate is1.03 ± 0.17 μW m−3, implying a correction of 0.52 mW m−2 km–1.

Sedimentation in the basin is of utmost importance. Its ther-mal effect can be approached with different methods, for example,by using a sudden deposition or a constant sedimentation model,which of course may give different results (e.g. Beardsmore &Cull 2001). Because several compressive tectonic phases, involv-ing shortening and overthrusting, have taken place in the basin, the




Figure 7. Radiogenic heat rate (H) derived from GR log of Chiari (a) and Carpaneto (b) wells (see Fig. 1 for location).

sedimentation rate of the different deposition cycles as well as ero-sion or lacking of sedimentation and compaction are difficult toquantify. Therefore, in order to evaluate the thermal effect of sedi-mentation, we used the simplified approach by Von Herzen & Uyeda(1963), which is based on the assumption of a constant sedimenta-tion rate (see the Appendix).

We modelled only the thermal effect of the most recent andimportant deposition cycle (Plio-Quaternary), considered as a singleevent, which took place on the Miocene formations acting as abasement. The model predicts that the decrease of surface heat flowis significant when sedimentation rate is >10−4 m yr−1 (Fig. 8).The average correction to heat−flow data is 22 per cent, but in the




Table 2. Radiogenic heat rate (H) of basin main lithotypesfrom GR log and γ−ray spectrometry (GRS).

Lithotype H (μW m−3)

GR log GRS

Sand 0.74 (0.13)Sandstone 1.05 (0.02)Siltstone 1.13 (0.12)Shale/silty shale 1.33 (0.24)Marl/silty marl 0.92 (0.14) 1.30 (0.14)Argillaceous sandstone 1.39 (0.23)Argillaceous limestone 0.63 (0.14)Mudstone/wackestone 0.45 (0.22) 0.34 (0.25)Dolostone 0.46 (0.32)Radiolarite 0.43 (0.09)Dacite 0.58 (0.21)Acid tuff 2.19 (0.15)Basaltic tuff 0.47 (0.05)Schist 2.27 (0.14)Orthogneiss 2.92 (0.33)

Figure 8. Ratio of observed heat flow qob to undisturbed heat flow qun ver-sus sedimentation time. Labels indicate the rate of sedimentation. Thermaldiffusivity κ = 30 m2 yr−1 was assumed (see the Appendix).

UPF unit, at Castelnovo and Bosco Rosso wells, where the coverthickness is 5400 and 6500 m, respectively, correction is as large as55 per cent (Table 1).

The palaeoclimate effect was evaluated with the depth-dependentcorrection curve proposed by Majorowicz & Wybraniec (2011)for south–southwestern Europe. The palaeoclimatic correctionas a response to five glacial cycles since 600 kyr ago withglacial–interglacial surface temperature amplitude of 7 ◦C wascalculated for a model with homogeneous thermal conductivity(2 W m−1K−1), diffusivity (28.4 m2 yr−1) and basal heat flow(60 mW m−2). Such past temperature changes caused a heat flowreduction, which smoothes with depth. The correction is about5 mW m−2 in the less deep wells (about 1200 m) and becomesnegligible at depth larger than 2000 m.

Table 1 lists the average thermal conductivity of the analysedwells, together with the heat-flow values, observed and after correc-

tion for radiogenic heat, sedimentation and palaeoclimate. Thermalconductivity and heat flow from previous studies (only the sedimen-tation correction was applied) are also shown for comparison. Theaverage corrected heat flow of the earlier data set is 48 ± 8, against67±7 mW m−2 of the new data set. The increase of the average heatflow (about 39 per cent) is due, above all, to the improved estimateof thermal conductivity, which accounts for by 24 per cent. The ra-diogenic heat correction (not applied earlier) affects the heat flowby only 4 per cent, whereas the remaining 11 per cent is given by thepalaeoclimate effect and the different technique of BHT correction.The sedimentary correction is nearly the same of that applied in theprevious study by Pasquale & Verdoya (1990).

By assuming that uncertainties on formation temperature, surfacetemperature, well depth and thermal conductivity are 3 ◦C, 0.2 ◦C,2.0 m and 0.25 W m−1K−1, respectively, for an average depth of3500 m and a thermal conductivity of 2.5 W m−1K−1, the bias oneach heat−flow value is ±10 per cent. The uncertainty on ther-mal conductivity is by far dominating the other contributions (seeClauser & Villinger 1990).

A heat-flow map of the study area is presented in Fig. 9. Contour-ing of surface heat−flow data was carried out with the kriging tech-nique. In the regions surrounding the basin, the map was adjustedby taking into account the data set from previous works (Table 3)reviewed by Pasquale (1985). In general, the heat flow of the basinis about 70 mW m−2 and presents local minima (<60 mW m−2).Heat flow tends to increase towards the Alps to 80–100 mW m−2. Innorthwesternmost part of the Apennines and the Molasse foredeep,it is on average 70 mW m−2.

4 I N T E R P R E TAT I O N

The heat flow in the Western Po Basin depicted by new and repro-cessed thermal data is typically within the 54–78 mW m−2 range.The goodness of the linear fit to temperature data (Fig. 6) sug-gests the thermal regime to be prevalently dominated by conduc-tion. Heat-flow minima (<60 mW m−2) might be evidence of local,minor convective heat transport. Further calculations would be nec-essary to quantify the influence of groundwater on the temperaturefield (e.g. Clauser & Villinger 1990).

Table 4 summarises the average heat flow in the basin tectonicunits and surrounding areas (see also Fig. 9). Heat flow variesacross the basin, being lower in the northern and southern units(66 ± 6 mW m−2 in SAB; 62 ± 7 mW m−2 in NAB) and larger(73 ± 4 mW m−2) in the UPF unit. Since data were corrected forsedimentation and climate change effects and regional groundwaterflow has low probability to occur, because of the thick, practicallyimpermeable sediment cover, such differences could be accountedfor by other geologic processes. Lateral change in the basal heatflow does not seem a possible explanation, as heat-flow variationhas a relatively short wavelength. Thus, we performed an additionalanalysis to discern the possible causes of such a heat-flow pattern.

The thermal structure along a cross-section that includes the sed-imentary sequences and the crystalline basement of the basin wasmodelled with a 2-D, finite element approach, under the assump-tion of steady-state conduction (Fig. 10). A constant temperature(12.5 ◦C) and an incoming vertical heat flow were assumed as sur-face and lower boundary conditions, respectively. For each elementin the finite-element mesh, changes of thermal conductivity dueto temperature, burial depth and anisotropy were taken into ac-count. The radiogenic heat rate in the sedimentary layers and inthe basement was assumed to be uniform. Values of 1.0, 1.2, 0.7




Figure 9. Contour map of heat flow (isolines in mW m−2). Position of heat-flow sites (dots) is shown (see Tables 1 and 3). (1) Molasse Foredeep; (2) DeformedForeland and Flysch Belt; (3) Penninic Zone; (4) main buried thrust fronts of the deformed in Southern Alps and Northwestern Apennines units. NAB, SABand UPF units as in Fig. 1.

Table 3. Surface heat-flow values measured in areas surrounding the basin (see Fig. 9 for site location). L, lakes; T, tunnels; B, boreholes. Numberof heat flow determinations in lakes within brackets.

Site code/name Latitude (N) Longitude (E) Elevation (m) Heat flow (mW m−2) Reference

1L Leman (3) 46◦ 27’ 06◦ 35’ 372 76 ± 6 Finckh (1981, 1983)2L Como (11) 46◦ 00’ 09◦ 15’ 198 86 ± 14 Haenel (1974); Finckh (1981, 1983)3L Mergozzo 45◦ 57’ 08◦ 28’ 196 86 Haenel (1974)4L Lugano (5) 45◦ 56’ 08◦ 57’ 271 67 ± 9 Finckh (1981, 1983)5L Maggiore (7) 45◦ 55’ 08◦ 35’ 194 91 ± 8 Haenel (1974); Finckh (1981, 1983)6L Orta (2) 45◦ 49’ 08◦ 24’ 290 86±1 Haenel (1974)7L Iseo (4) 45◦ 43’ 10◦ 04’ 185 92 ± 15 Haenel (1974); Finckh (1981, 1983)8L Garda (8) 45◦ 40’ 10◦ 43’ 65 100 ± 6 Finckh (1981); Haenel & Zoth (1982)9L Sirio 45◦ 29’ 07◦ 53’ 271 66 Haenel (1974)10L Viverone 45◦ 25’ 08◦ 02’ 230 65 Haenel (1974)1T Gotthard 46◦ 40’ 08◦ 36’ 1150 67 Clark & Niblett (1956)2T Guspisbach (shaft) 46◦ 36’ 08◦ 34’ 1685 71 Rybach et al. (1977)3T Loetschberg 46◦ 25’ 07◦ 42’ 1230 80 Clark & Niblett (1956)4T Simplon 46◦ 15’ 08◦ 07’ 700 92 Clark & Niblett (1956)5T Mont Blanc 45◦ 51’ 06◦ 53’ 1300 83 Bossolasco & Palau (1965)1B Biaschina 46◦ 25’ 08◦ 51’ 455 80 Haenel (1971); Bodmer & Rybach (1984)2B Chezallet 45◦ 44’ 07◦ 53’ 657 82 Haenel (1974)3B Marsaglia 44◦ 43’ 09◦ 23’ 305 78 Verdoya et al. (2007)4B Santuario 44◦ 21’ 08◦ 26’ 100 76 Pasquale et al. (2001)5B Pontremoli 44◦ 22’ 09◦ 52’ 236 68 Pasquale et al. (1993)

and 2.6 μW m−3 were chosen for Plio-Quaternary, Oligo-Miocene,Middle Triassic–Eocene and crystalline basement, respectively.

Several simulations have been carried out by varying the basalheat flow. The heat flow of the Undeformed Po Foredeed unit, whichis, in principle, the less affected by tectonothermal processes, wastaken as a benchmark. The calculated surface heat flow is relatively

uniform along the cross-section, and the isotherm pattern indicatesthat the effects due to heat refraction, associated with lateral vari-ability in thermal proprieties, are negligible. In the UPF unit, theminimum misfit (±2 mWm−2) between measured and modelledheat flow is obtained for a uniform basal heat flow of 49 mW m−2.In the SAB and NAB units, the observed surface heat flow is lower




Table 4. Average surface heat flow for the tectonic units of Fig. 9.

Tectonic unit No. of sites Heat flow (mW m−2)

Deformed Foreland and Flysch Belt 7 75 ± 8Penninic Zone 5 79 ± 9Southern Alps 39 84 ± 12Northwestern Apennines 3 70 ± 7Western Po Basin 38 67 ± 7

UPF 11 73 ± 4SAB 16 66 ± 6NAB 11 62 ± 7

than the modelled one by 9–12 mWm−2. In summary, numericalcalculations indicate that heat-flow differences among the differenttectonic units cannot be explained by lateral change in structureand thermal properties. This can argue that the origin of such aheat−flow pattern lies in the tectonothermal processes that havetaken place in the basin and the surrounding areas.

The evolution of the basin is controlled by the relative conver-gent motion of the Adriatic microplate and European Plate, andthe surface heat flow could be somehow affected by this process.Plate collision implies lithospheric shortening and overthrustingyielding a thermal perturbation, which can be still present. Forthe northeastern part of the Apennines, Pasquale et al. (1993) in-vestigated the thermal effects of overthrusting with the model byBrewer (1981) and demonstrated that the surface heat flow is sen-

sibly reduced because of the recent thrust. Here, we apply the samekind of approach (see the Appendix) and evaluate the order-of-magnitude of the thermal field perturbations in the SAB and NABunits.

Data about geometry and timing of the deformation of the basinburied thrusts show that the youngest compressional tectonic eventhas taken place from the Serravallian to the Late Messinian inthe SAB unit and from the Messinian to the Pleistocene in theNAB unit (Fantoni et al. 2004; Costa 2003). The main detach-ment surfaces seem to lie within Neogenic terrains and next to theMesozoic–Tertiary boundary, but some surfaces extend to greaterdepth and could be located at the Mesozoic basement and sometimesin the crystalline basement too. Due to the difficulty in defining thereal depth of the detachment level, we assumed a slab with thicknessof 4 km in the SAB unit, and a thicker slab (5 km) in the NAB unit,including the cover and some upper levels of the basement.

Fig. 11 depicts the surface heat-flow variation with time, obtainedby means of eqs (A3) and (A4) for a slab that slips horizontally atconstant rates of 0.5, 1 and 2 cm yr−1. The surface heat flow of73 mW m−2 of the UPF unit is assumed as the equilibrium value(Table 4). The heat flow initially decreases by about 35–45 per cent,then it increases until the end of thrusting, reaching values of about60–70 mW m−2 for a slip rate of 0.5 cm yr−1 and 85–95 mW m−2

for a slip rate of 2 cm yr−1. Subsequently, it decreases only dueto thermal relaxation. The observed heat-flow values in the NABand SAB units are better matched for slip rates of 1 cm yr−1.

Figure 10. Above: observed (full circles) and modelled (continuous line) surface heat flow along cross-section AB (see Fig. 9). Average observed heat flowis indicated (thin broken line). Below: upper crust structure (after Fantoni & Franciosi, 2010) and modelled isotherms. (1) Terrigenous, (2) carbonate and (3)crystalline basement units. Digit within the circle is the well code of Table 1.




Figure 11. Heat-flow variation arising from overthrusting at different rates of slip in the SAB and NAB units. Model parameters: g = 9.8 m s−2, κ = 30 m2 yr−1,k = 2.5 W m−1 K−1, cp = 1 kJ kg−1 K−1, ρ = 2600 kg m−3 and f = 0.6 (see the Appendix).

Overthrusting has also affected the Alps chain, but the thermaleffects should be negligible on the present-day heat flow (Rybachet al. 1977).

5 C O N C LU S I O N S

Accurate processing of new temperature data and a revision ofprevious thermal information provide new estimations of surfaceheat flow in the Western Po Basin. The uncertainties on formationthickness, surface temperature and formation temperature are ofminor importance in the determination of the heat flow, compared tothe bias in thermal conductivity. Therefore, particular care was takenin evaluating the basin rock properties under any possible conditionof burial depth, temperature and anisotropy. To this purpose, we usedan approach that allows the inference of in situ thermal parameterson the basis of litostratigraphic data. Heat flow was corrected forthe cooling effects of the Plio-Quaternary sedimentation and thepalaeoclimate change, and for the radiogenic heat.

The revised heat-flow values are larger than those reported by pre-vious studies. This is due to better estimates of thermal conductivity,the use of more accurate techniques to infer the formation tempera-tures and, secondarily, to the radiogenic heat and the palaeoclimate,which were not taken into account in early studies. The new datashow that the heat flow in the basin ranges from 54 to 78 mW m−2,with lower values occurring in the units with recent overthrusting.The thermal regime seems chiefly conductive, with the exceptionof small areas bounded by heat flow <60 mW m−2, which might beascribed to local groundwater movement.

The obtained heat-flow information together with data availablefor the surrounding orogenic belts allows us to draw a new heat-flow map of northwestern Italy. The tectonic difference between thebasin and the surrounding orogenic belts seems to be reflected by theheat-flow lateral variation. In the Alps the surface heat flow is withinthe 65–100 mW m−2 range, and in northwesternmost part of theApennines and the Molasse foredeep it is on average 70 mW m−2.

A C K N OW L E D G M E N T S

This work was carried out within the framework of the MIUR-2008project ‘Geothermal resources of the Mesozoic basement of the PoBasin: groundwater flow and heat transport’.

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A P P E N D I X : T I M E - D E P E N D E N TD I S T U R B A N C E S O F T H E S U R FA C EH E AT F L OW

Sedimentation

The method by Von Herzen & Uyeda (1963) assumes a horizontaldepositional surface at depth z = 0 which is in thermal equilibriumwith the underlying basement with constant thermal diffusivity andunbiased (initial) heat flow qun. Deposition of sediments with thesame thermal properties of the underlying basement begins at timet = 0 and continues at a constant rate V . The heat transfer equationis

∂2T

∂z2− V

κ

dT

dz= 1

κ

∂T

∂t, (A1)

whose solution in terms of surface heat flow is

qob = qun

[1 − erf (X ) − 2X√

πe−X + 2X 2erfc (X )

], (A2)

where κ is the thermal diffusivity, X = 0.5 V (t/κ)1/2, qob is the biased(observed) heat flow, and erfc denotes the complementary errorfunction. Such a sedimentation model gives the order of magnitudeof the heat flow bias caused by sedimentation.

Overthrusting

The model by Brewer (1981) demands assumptions on the physicalparameters of medium and the definition of the thickness h of thethrusting slab, the rate of slip u and the coefficient of friction f atthe base of the slab. The surface heat flow at any time t after start ofthrusting qo(t) is given by the sum of the surface heat flow arisingfrom thermal relaxation:

qr = qo

⎛⎝1 − he− h2

4π t√πκt

⎞⎠ (A3)

and from the contribution of frictional heating

q f = kuτ

ρcPκerfc

(h

2√

κt

)t ≤ t1

q f = kuτ

ρcPκ

[erfc

(h

2√

κt

)− erfc

(h

2√

κ (t − t1)

)]t > t1

(A4)

where ρ, k, κ and cp are density, thermal conductivity, thermaldiffusivity and specific heat of the rock, respectively, and qo is thesurface heat flow before thrusting. The rate of heating is proportionalto the stress across the fault: τ = fρgh, where g is acceleration dueto gravity and t1 is the time of thrusting. This model holds if heattransport by solid-state convection dominates conductive flow awayfrom the thrust.



[Pasquale Et Al. 2012]_Heat Flow in the Western Po Basin and Surrounding Orogenic Belts

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